Microfacies of Carbonate Rocks
¨ Erik Flugel
Microfacies of Carbonate Rocks Analysis, Interpretation and Application Second Edition
With a contribution by Axel Munnecke
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ISBN 978-3-642-03795-5 e-ISBN 978-3-642-03796-2 DOI 10.1007/10.1007/978-3-642-03796-2 Springer Heidelberg Dordrecht London New York Library of Congress Control Number: 2009935385 c Springer-Verlag Berlin Heidelberg 2010 This work is subject to copyright. All rights are reserved, whether the whole or part of the material is concerned, specifically the rights of translation, reprinting, reuse of illustrations, recitation, broadcasting, reproduction on microfilm or in any other way, and storage in data banks. Duplication of this publication or parts thereof is permitted only under the provisions of the German Copyright Law of September 9, 1965, in its current version, and permission for use must always be obtained from Springer. Violations are liable to prosecution under the German Copyright Law. The use of general descriptive names, registered names, trademarks, etc. in this publication does not imply, even in the absence of a specific statement, that such names are exempt from the relevant protective laws and regulations and therefore free for general use. Camera-ready by Erentraud Fl¨ugel-Kahler, Erlangen Cover design: WMXDesign, Heidelberg Printed on acid-free paper Springer is part of Springer Science+Business Media (www.springer.com)
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Preface to the Second Edition
A second editon of this book - yes or no - or not yet? Are changes necessary or would it be enough if minor errors in the text were corrected? So I asked for help and I got answers from friends and colleagues. I wish to thank W.Ch. Dullo (Kiel), A. Freiwald (Erlangen), H.-G. Herbig (Cologne), W. Piller (Graz), W. Schlager (Amsterdam), R. J. Stanton (Thousand Oaks, USA) for their excellent and friendly advice. They were in agreement that not enough has changed in the scientific field of ‘Microfacies of Carbonate Rocks’ since 2004 when the first edition was published for a new edition to be necessary, but they all believed that an update of the references on the CD going with the volume would be useful. Therefore, the reference section has been updated to April 2009. Most of the new references are since 2002, but a few are earlier. The reference file now contains more than 16,000 references. In connection with the keywords used throughout the book these references can help the reader locate special subjects or overviews of interest. When Erik wrote this book during his last years, his primary goal was to provide students and all those interested in and working with microfacies a helpful und useful resource that contained many plates and figures.
Most of the pictures in this book were collected during Erik Flügel’s scientific work, some are from papers that were published in the journal Facies, then published by the Institute of Paleontology (now Geozentrum Nordbayern), University Erlangen-Nürnberg, and last not least, some were graciously provided by colleagues. The author had hoped to include an additional chapter with plates that would provide the reader with an expanded range of information about microfacies. This task has been accomplished by A. Munnecke (Erlangen). Chapter 20 contains plates and figures that have been used successfully in the Course on Microfacies held every year or two at the University in Erlangen. I wish to thank all of my colleagues at the institute in Erlangen, and especially A. Munnecke for the compilation of Chapter 20. Ch. Schulbert was most helpful in solving software problems with the text and CD. I thank Ch. Bendall (Springer Verlag Heidelberg) for overseeing the printing of this new edition. Erlangen, June 2009 Erentraud Flügel-Kahler
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Preface
The objective of this book is to provide a synthesis of the methods used in microfacies studies of carbonate rocks and to show how the application of microfacies studies has contributed to new developments in carbonate geology. In contrast with other textbooks on carbonate sedimentology this book focuses on those compositional and textural constituents of carbonates that reflect the depositional and diagenetic history and determine the practical usefulness of carbonate rocks. The chapters are written in such a way, that each one can be used as text in upper level undergraduate and graduate courses. The topics of the book also apply to research workers and exploration geologists, looking for current information on developments in the use of microfacies analysis. Since microfacies studies are based on thin sections, instructive plates showing thin-section photographs accompanied by thorough and detailed explanations form a central part of this book. All plates in the book contain a short summary of the topic. An –> sign leads the reader to the figures on the plate. The description of the microphotographs are printed in a smaller type. Care has been taken to add arrows and/or letters (usually the initials of the subject) so that the maximum information can be extracted from the figures. Rather than being a revised version of ‘Microfacies Analysis of Limestones’ (Flügel 1982) ‘Microfacies of Carbonate Rocks’ is a new book, based on a new concept and offering practical advice on the description and interpretation of microfacies data as well as the application of these data to basin analysis. Microfacies analysis has the advantages over traditional sedimentological approaches of being interdisciplinary, and integrating sedimentological, paleontological and geochemical aspects. ‘Microfacies of Carbonate Rocks’: • analyses both the depositional and the diagenetic history of carbonate rocks, • describes carbonate sedimentation in various marine and non-marine environments, and considers both tropical warm-water carbonates and non-tropical cool-water carbonates,
• presents diagnostic features and highlights the significance of microfacies criteria, • stresses the biological controls of carbonate sedimentation and provides an overview on the most common fossils found in thin sections of limestones, • discusses the relationships between diagenetic processes, porosity and dolomitization, • demonstrates the importance of microfacies for establishing and evaluating sequence stratigraphic frameworks and depositional models, • underlines the potential of microfacies in differentiating paleoclimate changes and tracing platform-basin relationships, and • demonstrates the value of microfacies analysis in evaluating reservoir rocks and limestone resources, as well as its usefulness in archaeological provenance studies. Structure of the Book Microfacies of Carbonate Rocks starts with and introductory chaper (Chap.1) and an overview of modern carbonate deposition (Chap. 2) followed by 17 chapters that have been grouped into 3 major parts. Microfacies Analysis (Chap. 3 to Chap. 10) summarizes the methods used in microfacies studies followed by discussions on descriptive modes and the implications of qualitative and quantitative thin-section criteria. Microfacies Interpretation (Chap. 11 to Chap. 16) demonstrates the significance of microfacies studies in evaluating paleoenvironment and depositional systems and, finally, Practical Use of Microfacies (Chap. 17 to Chap. 19) demonstrates the importance of applied microfacies studies in geological exploration for hydrocarbons and ores, provides examples of the relationships between carbonate rock resources and their facies and physical properties, and also illustrates the value of microfacies studies to archaeologists. Important references are listed at the end of chapters or sections under the heading ‘Basics’. The code
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numbers K... are keywords leading to references on specific fields of interest (see CD), e.g. K021 (coldwater carbonates), K078 (micrite), or K200 (hydrocarbon reservoir rocks). The book is also accompanied by a CD containing • an alphabetical list of about 16,000 references (updated in this version up to April 2009) on carbonate rocks (see also Appendix) as word document, and • visual comparison charts for percentage estimation.
Synopsis of the Book‘s Contents Chapter 1: New Perspectives in Microfacies. Microfacies studies, which were originally restricted to the scale of thin sections, provide an invaluable source of information on the depositional constraints and environmental controls of carbonates, as well as on the properties of carbonate rocks. Microfacies studies assist in understanding sequence stratigraphic patterns and are of economic importance both in reservoir studies and in the evaluation of limestone resources. Chapter 2: Modern Depositional Environments. Knowledge of modern carbonates is a prerequisite for understanding ancient carbonate rocks. Modern carbonates are formed both on land and in the sea, in shallow- and in deep marine settings, and in tropical and non-tropical regions, but the present is only part a key to the Past. Microfacies Analysis Chapter 3: Methodology. Which methods can be used in the field? Which sampling strategy should be applied and how many samples are required? Which laboratory techniques are useful in microfacies analysis? Which other techniques should be combined with microfacies studies? Chapter 4: Microfacies Data: Matrix and Grains. This chapter offers practical advice on how to handle microfacies data, and deals with how to describe and interpret thin-section characteristics. Matrix types and grain categories are discussed with regard to their diagnostic criteria, origin and significance. Chapter 5: Microfacies Data: Fabrics. Typical depositional and diagenetic fabrics in limestones reflect the history of the rock. Microfacies criteria indicating breaks and changes in sedimentation (discontinuity surfaces) are of specific interest in refining sequence stratigraphic boundaries. Variously sized fissures, microcracks and breccias can be used in deciphering synand post-depositional destructive processes.
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Chapter 6: Quantitative Microfacies Analysis. Whilst previous chapters focused on qualitative criteria this chapter deals with quantitative data including grain size analysis, frequency analysis and multivariate studies. Constituent analysis and the distribution patterns of specific grain types are valuable tools in the reconstruction of paleoenvironmental controls and depositional settings. Chapter 7: Diagenesis, Porosity and Dolomitization. Understanding diagenetic processes and their products is of high economic importance. The diagenetic microfacies of a rock reflects changes in the course of its lithification history. The main topics discussed in this chapter are porosity types, carbonate cements, diagenetic textures including compaction and pressure solution, and dolomitization/dedolomitization and dolomite textures. The last part of the chapter deals with thinsection criteria for metamorphic carbonates and marbles. Chapter 8: Classification – chosing a name for your sample. A classification is simply a tool for organizing information, and should not be the only source of conclusions. Whilst a name based on texture and composition can not replace a well-defined microfacies type, rock names are essential for the categorization of samples. Textural classifications proposed by Dunham and by Folk have proven to be the most practical. Specific concepts must be adhered to in the naming of reef limestones, non-marine carbonates, recrystallized carbonates rocks and mixed siliciclastic-carbonate rocks. Chapter 9: Limestones are Biological Sediments. In contrast to siliciclastic rocks, both the formation and the destruction of most limestones is directly or indirectly influenced and controlled by biological processes. This chapter stresses the biological controls on carbonate sedimentation. Microbes, encrusting organisms, and macro- and microborers can yield useful information on paleoenvironment, depositional constraints and carbonate production. Chapter 10: Fossils in Thin Sections. It Is Not That Difficult. The recognition of fossils in thin sections is not so difficult once the diagnostic criteria for the main groups have been understood, particularly for algae and foraminifera, sessile invertebrates, and organisms with shells. This chapter provides an overview of the most common fossils found in thin sections of limestones. The text concentrates on identification criteria, environmental and temporal distribution, and on the significance of the fossils. Numerous instructive plates are included to aid in the recognition and differentiation skeletal grains in thin sections.
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Microfacies Interpretation Chapter 11: Summarizing Microfacies Criteria: Microfacies Types. How can microfacies data be combined in sensitive and practicable microfacies types? Which criteria should be used, which grain types are of particular importance and, how many microfacies types are reliable? The creation of microfacies types is illustrated by means of examples. Chapter 12: Recognizing Paleoenvironmental Conditions. Carbonate sediments are particularly sensitive to environmental changes. Microfacies and organisms are excellent paleoenvironmental proxies as they reflect hydrodynamic conditions, the impact of storms, substrate conditions, light, oxygenation, seawater temperature and salinity. Significant differences between the compositions of skeletal grain associations for warm-water and cold-water carbonates provide a useful tool for estimating paleoclimatic changes. How deep was the sea? Microfacies studies provide an answer. Chapter 13: Integrated Facies Analysis. Understanding the formation and diagenesis of carbonate rocks requires the combination of microfacies with mineralogical and geochemical data. The chapter deals with acid-insoluble residues and authigenic minerals in carbonate rocks, discusses the value of minor elements and stable isotopes in tracing the depositional and diagenetic history of limestones, and deals with the potential of organic matter in carbonate rocks for facies analysis. Chapter 14: Depositional Models, Facies Zones and Standard Microfacies. Microfacies are essential for defining depositional models and recognizing facies zones. Facies models assist in understanding depositional history. Changes in sedimentological and biological criteria across shelf-slope-basin transects form the basis of generalized models for carbonate platforms, ramps and shelves. Facies belts are reflected by their biotic zonation patterns and the distribution of Standard Microfacies Types (SMF Types). The latter are virtual categories that summarize microfacies with identical criteria. Which criteria are used in differentiating the SMF Types of platform and ramp carbonates? What are the problems involved in the SMF concept? Revised and refined SMF types are a meaningful tool in tracing facies belts, but must be used with care. Common microfacies of carbonate ramps (Ramp Microfacies Types) show only partial correspondence to the SMF Types of rimmed platforms. Chapter 15: Basin Analysis: Recognizing Depositional Settings. Which diagnostic criteria characterize lime-
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stones of different carbonate systems? Case studies for non-marine and marine carbonate rocks demonstrate how to translate microfacies into ancient depositional settings. Non-marine settings can be successfully reconstructed for pedogenic carbonates, paleokarst deposits and ancient speleothems, travertine deposits, and lacustrine carbonates, these can be characterized by specific microfacies types. Marine settings can be differentiated into peritidal carbonates, platforms and ramps, platform-slope-basin transects, and pelagic deep-marine carbonates. Grain Composition Logs are particularly effective in tracing platform-basin relations. Chapter 16: Recognizing Depositional Constraints and Processes. Selected case studies are used to demonstrate the value of microfacies data in interpreting depositional controls. • How can microfacies be used in sequence stratigraphy? Cyclic depositional patterns and sequence stratigraphic constraints are documented by microfacies data that assist in recognizing sequence boundaries, parasequences, high-frequency sea-level changes, and systems tracts. • Which criteria characterize reef limestones? Major reef types differ in biota, matrix, sediment, and cements. Which methods should be employed in reconstructing former platforms and reefs that are only recorded by eroded relicts deposited on slopes and in basins? Clast analysis can solve this puzzle. • Which criteria define ancient cold-water carbonates? Ancient cool-water shelf and reef carbonates are typified by specific biotic, compositional and diagenetic features. • Which facies criteria are diagnostic of ancient ventand seep carbonates? Case studies provide answers. • How to handle mixed carbonate-siliciclastic sediments and interpret limestone-marl successions? • Constraints on carbonate deposition exhibit secular variations, which are discussed in the last section.
Practical Use of Microfacies Chapter 17: Reservoir Rocks and Host Rocks. Carbonates are the most important reservoir rocks for hydrocarbons as well as forming important host rocks for ores. Limestones and dolomites contain more than 50% of the world’s oil and gas reserves. Reservoir potential differs for carbonates formed in different depositional settings and depends on the interplay of depositional processes and diagenetic history. The microfacies of cores and cuttings assist in the translation of lithological data into petrophysical information. Facies-based outcrop-analogue studies indicate the scale of porosity and
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permeability variations within carbonate bodies. Microfacies analysis also assists in the genetic interpretation of carbonate-hosted base metal deposits controlled by specific facies patterns. Chapter 18: Carbonate Resources, Facies Control and Rock Properties. Carbonate rocks are important raw materials for chemical and construction industries and are high on the list of extracted mineral resources, both in terms of quantity and of value. Both exploration and exploitation can be enhanced by taking into account the relationships between depositional and diagenetic facies that control technologically relevant chemical and physical parameters, as well as the weathering and decay properties of carbonate rocks. Conservation and preservation of works of art and building stones should start with thin-section studies of the textural and diagenetic criteria that describe the porosity and permeability of the material. Chapter 19: Archaeometry. Microfacies analysis, combined with geochemical data has considerable potential in provenance analysis of archaeological materials. Thin sections reveal the source of building stones and of material used for mosaics and works of arts. The microfacies of temper grains in ancient pottery helps in understanding the source and production areas for ceramics. Last but not least, microfacies studies can throw new light at the love affair between Antony and Cleopatra.... Acknowledgments Thanks are due to many people who have provided photographs, information and advice: Gernot Arp (Göttingen), Martina Bachmann (Bremen), Benoit Beauchamp (Calgary), Thilo Bechstädt (Heidelberg), Michaela Bernecker (Erlangen), Joachim Blau (Giessen), Florian Böhm (Kiel), Thomas Brachert (Mainz), Ioan Bucur (Cluj-Napoca), Werner Buggisch (Erlangen),Thomas Clausing (Halle), Wolf-Christian Dullo (Kiel), Paul Enos (Lawrence, Kansas), Gerd Flajs (Aachen), Christof Flügel (München), Helmut Flügel (Graz), Beate Fohrer (Erlangen), Holger Forke (Berlin), André Freiwald (Erlangen), Robert van Geldern
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(Erlangen), Markus Geiger (Bremen), Gisela Gerdes (Oldenburg), Eberhard Gischler (Frankfurt), Dirk von Gosen (Erlangen), Jürgen Grötsch (Damascus). HansGeorg Herbig (Köln), Richard Höfling (Erlangen), Bernhard Hubmann (Graz), Andi Imran (Makassar), Michael Joachimski (Erlangen), Josef Kazmierczak (Warszawa), Martin Keller (Erlangen), Stephan Kempe (Darmstadt), Helmut Keupp (Berlin), Wolfgang Kiessling (Berlin), Roman Koch (Erlangen), Karl Krainer (Innsbruck), Jochen Kuss (Bremen), Michael Link (Erlangen), Heinz Lorenz (Erlangen), Ulrich Michel (Nürnberg), Axel Munnecke (Erlangen), Fritz Neuweiler (Göttingen), Alexander Nützel (Erlangen), Joachim Reitner (Göttingen), Jürgen Remane (Neuchâtel), Elias Samankassou (Genéve), Diethard Sanders (Innsbruck), Chris Schulbert (Erlangen), Baba Senowbari-Daryan (Erlangen), Robert J. Stanton (Thousand Oaks, California), Torsten Steiger (Bad Blankenburg), Thomas Steuber (Bochum), Harald Tragelehn (Köln), Jörg Trappe (Bonn), Dragica Turnsek (Ljubljana), Andreas Wetzel (Tübingen). I am indebted to Birgit Leipner-Mata and Marieluise Neufert (Institute of Paleontology Erlangen) for laboratory work and photography. Chris Schulbert (Institute of Paleontology) was a valuable help with all computer problems. I am sincerely grateful to my friend Johann Georg Haditsch (Graz) for critical reading the text of the book. Special thanks go to Karen Christenson (NürnbergKraftshof) who took care of linguistic problems and pitfalls. Editing, layout, drawings and the preparation of plates and figures have been carried out by my wife Erentraud Flügel-Kahler. I am most grateful for her encouragement and constant help. Finally, I am obliged to all institutions who gave permission to use published material and to the staff of Springer Verlag, especially to Dr. Wolfgang Engel for his constant encouragement and to Dr. J. Witschel for assistance with the book.
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Contents 1 1.1 1.2
New Perspectives in Microfacies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The Microfacies Concept . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . New Perspectives . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
1 1 1
2 2.1 2.1.1 2.1.2
Carbonate Depositional Environments . . . . . . . . . . . . . . . . . . . . . . . . . . Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonates are Born not Made . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . The ‘Sorby Principle’: Limestones are Predominantly Biogenic Sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Modern Carbonates: Obligatory Reading . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Sediments Originate on Land and in the Sea . . . . . . . . . . . . . . . Classification of Marine Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . Boundary Levels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Vertical and Horizontal Zonations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Vertical Zonations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Horizontal Zonations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Review of Modern Carbonate Depositional Environments . . . . . . . . . . . . . Non-Marine Carbonate Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pedogenic Carbonates, Paleosols, and Caliche/Calcretes . . . . . . . . . . . . . . Palustrine Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cave Carbonates, Speleothems and Karst . . . . . . . . . . . . . . . . . . . . . . . . . . Glacial Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Travertine, Calcareous Tufa and Calcareous Sinter . . . . . . . . . . . . . . . . . . . Lacustrine Carbonates: Lakes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluvial Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Transitional Marginal-Marine Environments: Shorelines and Peritidal Sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Beach (Foreshore), Barriers and Coastal Lagoons . . . . . . . . . . . . . . . . . . . . Peritidal Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Shallow-Marine Sedimentary Environments: ‘Shallow’ and ‘Deep’ . . . . . . Pericontinental vs Epicontinental Shallow Seas . . . . . . . . . . . . . . . . . . . . . Carbonate Shelves, Ramps and Platforms . . . . . . . . . . . . . . . . . . . . . . . . . . Shelf Margins . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reefs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tropical and Non-Tropical Carbonates: Different in Composition, Controls and Significance . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Latitudinal Zonation and Diagnostic Criteria of Tropical and Non-Tropical Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tropical and Subtropical Shallow-Marine Carbonates . . . . . . . . . . . . . . . . Non-Tropical Shelf and Reef Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . Deep-Marine Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Settings . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sedimentation Processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pelagic Sedimentation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Resedimentation (‘Allochthonous Carbonates’) . . . . . . . . . . . . . . . . . . . . .
7 7 7
2.1.3 2.2 2.3 2.3.1 2.3.2 2.3.2.1 2.3.2.2 2.4 2.4.1 2.4.1.1 2.4.1.2 2.4.1.3 2.4.1.5 2.4.1.6 2.4.1.7 2.4.1.8 2.4.2 2.4.2.1 2.4.2.2 2.4.3 2.4.3.1 2.4.3.2 2.4.3.3 2.4.3.4 2.4.4 2.4.4.1 2.4.4.2 2.4.4.3 2.4.5 2.4.5.1 2.4.5.2 2.4.5.3 2.4.5.4
7 7 8 8 8 9 9 10 10 10 10 13 13 14 14 18 21 24 24 24 25 26 26 28 29 30 31 34 37 46 46 47 47 49
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2.4.5.5 2.4.5.6 2.4.5.7 2.4.6
3 3.1 3.1.1 3.1.1.1 3.1.1.2 3.1.1.3 3.1.1.4 3.1.2 3.1.2.1 3.1.2.2 3.1.2.3 3.2 3.2.1 3.2.2 3.2.3 3.2.3.1 3.2.3.2 3.2.3.3 3.2.4 3.2.5 4 4.1 4.1.1 4.1.2 4.1.3 4.1.4 4.1.5 4.1.6 4.2 4.2.1 4.2.2 4.2.3 4.2.4 4.2.4.1 4.2.4.2 4.2.5 4.2.6 4.2.7 4.2.8 4.2.8.1 4.2.8.2 4.3 4.3.1 4.3.2
Contents
Carbonate Plankton and Carbonate Oozes . . . . . . . . . . . . . . . . . . . . . . . . . . Preservation Potential and Dissolution Levels . . . . . . . . . . . . . . . . . . . . . . . Carbonate Slopes, Periplatform Carbonates and Caronate Aprons . . . . . . . Seep and Vent Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Analysis Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Field Work and Sampling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Field Observations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Lithology, Texture and Rock Colors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bedding and Stratification, Sedimentary Structures and Diagenetic Features . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fossils and Biogenic Structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Field Logs and Compositional Logs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sampling . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Search Sampling and Statistical Sampling . . . . . . . . . . . . . . . . . . . . . . . . . . How Many Samples? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Practical Advice for Microfacies Sampling . . . . . . . . . . . . . . . . . . . . . . . . . Laboratory Work: Techniques . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Slices, Peels and Thin Sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Casts, Etching and Staining . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microscopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Petrographic Microscopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stereoscan Microscopy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fluorescence, Cathodoluminescence and Fluid Inclusion Microscopy . . . . Mineralogy and Geochemistry . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Trace Elements and Stable Isotope Analysis . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Data: Matrix and Grains . . . . . . . . . . . . . . . . . . . . . . . . . . . Fine-Grained Carbonate Matrix: Micrite, Microspar, Calcisiltite . . . . . . . . Micrite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Modes of Formation of Micrite and Other Fine-Grained Matrix Types . . . Microspar . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calcisiltite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Practical Aids in Describing and Interpreting Fine-Grained Limestones . . Significance of Fine-Grained Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Grains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bioclasts (Skeletal Grains) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Peloids: Just a Term of Ignorance? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cortoids – Carbonate Grains Characterized by Micrite Envelope . . . . . . . . Oncoids and Rhodoids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Oncoids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Rhodoids and Macroids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ooids . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pisoids and Vadoids – Simply ‘Larger Ooids’ or Carbonate Grains on their Own? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Aggregate Grains: Grapestones, Lumps and Other Composite Grains . . . . Resediments: Intra-, Extra- and Lithoclasts – Insiders and Foreigners . . . . Intraclasts: Origin and Facies-Diagnostic Types . . . . . . . . . . . . . . . . . . . . . Extraclasts: Strange Foreigners . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Morphometry of Carbonate Grains . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Intentions and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Morphometric Data for Carbonate Grains . . . . . . . . . . . . .
49 49 50 51
53 53 53 53 55 59 60 61 62 62 63 64 64 65 66 66 66 67 70 70 73 73 74 80 94 95 98 98 100 101 110 118 121 124 137 142 157 163 166 167 172 173 174 175
Contents
XIII
5.2.4.1 5.2.4.2 5.2.5 5.3 5.3.1 5.3.1.1 5.3.1.2 5.3.1.3 5.3.1.4 5.3.2 5.3.2.1 5.3.2.2 5.3.2.3 5.3.3 5.3.3.1 5.3.3.2 5.3.3.3 5.3.3.4 5.3.3.5
Microfacies Data: Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Depositional and Diagenetic Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geopetal Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Biofabrics and Grain Orientation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bedding and Lamination Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Burrowing and Bioturbation Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Birdseyes, Fenestral Fabrics and Stromatactis . . . . . . . . . . . . . . . . . . . . . . . Birdseyes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fenestral Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Stromatactis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Birdseyes, Fenestral Fabrics and Stromatactis . . . . . . . . . . Nodular Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discontinuity Surfaces: From Microfacies to Sequence Stratigraphy . . . . . Classification of Discontinuities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Major Criteria of Discontinuities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Criteria and Significance of Exposure Surfaces . . . . . . . . . . . . Microfacies Criteria and Significance of Condensation Surfaces and Hardgrounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hardgrounds . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Condensation Surfaces and Condensed Sections . . . . . . . . . . . . . . . . . . . . . Discontinuities and Sequence Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . Syn- and Postdepositional Features: Fissures, Veins, Breccias . . . . . . . . . . Sediment-Filled Fissures: Neptunian Dikes and Fissure Fills . . . . . . . . . . . Origin, Development and Filling of Sedimentary Fissures . . . . . . . . . . . . . Microfacies Analysis of Neptunian Dikes . . . . . . . . . . . . . . . . . . . . . . . . . . Case Studies of Neptunian Dikes in Carbonates . . . . . . . . . . . . . . . . . . . . . Significance of Sediment-Filled Fissures . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfractures and Veins (Calcite Veins) . . . . . . . . . . . . . . . . . . . . . . . . . . . Origin and Classification of Calcite Veins . . . . . . . . . . . . . . . . . . . . . . . . . . Descriptive Criteria of Calcite-Filled Microfractures . . . . . . . . . . . . . . . . . Significance of Microfractures in Carbonate Rocks . . . . . . . . . . . . . . . . . . Carbonate Breccias and Conglomerates . . . . . . . . . . . . . . . . . . . . . . . . . . . . Terminology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . How to Describe Carbonate Breccias? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Breccia Types: Origin, Classification, Criteria . . . . . . . . . . . . . . Carbonate Conglomerates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Carbonate Breccias and Conglomerates . . . . . . . . . . . . . . .
206 206 211 215 216 217 217 221 221 223 223 224 225 225 228 228 229 233 239 242
6 6.1 6.1.1 6.1.1.1 6.1.1.2 6.1.1.3 6.1.2 6.1.2.1 6.1.2.2 6.1.2.3 6.2 6.2.1 6.2.1.1 6.2.1.2 6.2.1.3
Quantitative Microfacies Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Grain-Size Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Grain-Size Analysis: Methods and Aims . . . . . . . . . . . . . . . . . . . . . . . . . . . Measuring Grain Sizes and Describing Size Distributions . . . . . . . . . . . . . Approaches to the Environmental Interpretation of Grain-Size Data . . . . . Grain-Size Analysis in Thin Sections . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Grain-Size Studies of Modern and Ancient Carbonates . . . . . . . . . . . . . . . . Grain-Size Studies of Modern Carbonate Sediments . . . . . . . . . . . . . . . . . . Application of Grain-Size Analyses to Carbonate Rocks . . . . . . . . . . . . . . Significance of Grain-Size Studies of Carbonate Rocks . . . . . . . . . . . . . . . Frequency Analysis of Microfacies Data . . . . . . . . . . . . . . . . . . . . . . . . . . . Methods of Frequency Analyses . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Counting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Estimating . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Image Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
243 243 244 244 246 247 248 248 251 253 254 254 254 257 259
5 5.1 5.1.1 5.1.2 5.1.3 5.1.4 5.1.5 5.1.5.1 5.1.5.2 5.1.5.3 5.1.5.4 5.1.6 5.2 5.2.1 5.2.2 5.2.3 5.2.4
177 177 177 181 184 185 190 192 192 193 197 198 203 203 203 205
XIV
6.2.1.4 6.2.1.5 6.2.2 6.3 6.3.1 6.3.2
7 7.1 7.1.1 7.1.2 7.1.3 7.1.4 7.1.5 7.1.5.1 7.1.5.2 7.2 7.2.1 7.2.1.1 7.2.1.2 7.2.1.3 7.2.1.4 7.2.2 7.3 7.3.1 7.3.1.1 7.3.1.2 7.3.1.3 7.3.2 7.3.3 7.4 7.4.1 7.4.2 7.4.2.1 7.4.2.2 7.4.3 7.4.4 7.4.4.1 7.4.4.2 7.4.4.3 7.4.5 7.5 7.5.1 7.5.2 7.5.3 7.6 7.6 1 7.6.2 7.7 7.8 7.8.1
Contents
Constituent Ranking, Diversity and Maturity . . . . . . . . . . . . . . . . . . . . . . . Integrated Frequency Studies of Reef Carbonates . . . . . . . . . . . . . . . . . . . . Practical Advice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Multivariate Microfacies Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Methods: Variations between Constituents and between Samples . . . . . . . . Significance of Multivariate Studies: Constituent Analysis as a Clue to Environmental Conditions and Depositional Settings . . . . . . . . . . .
260 260 261 262 262
Diagenesis, Porosity, and Dolomitization . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Mineralogy and Diagenetic Processes . . . . . . . . . . . . . . . . . . . . Modern Carbonate Sediments and Ancient Carbonate Rocks . . . . . . . . . . . Common Carbonate Minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Processes and Controls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . From Soft Sediments to Hard Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Oscillating Trends in Phanerozoic Carbonate Mineralogy . . . . . . . . . . . . . Secular Variations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . How to Recognize Former Aragonite and Mg-Calcite Mineralogy in Ancient Low-Calcite Limestones? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Major Diagenetic Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Meteoric, Marine and Burial Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . Meteoric (Freshwater) Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mixing Zone and Marine Vadose Environment . . . . . . . . . . . . . . . . . . . . . . Marine Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Burial Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Early and Late Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Porosity of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Porosity Categories, Pore Geometry and Permeability . . . . . . . . . . . . . . . . Basic Definitions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pore Geometry and Permeability . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Porosity Measurements and Pore Types in Thin Sections . . . . . . . . . . . . . . Porosity Terminology and Classification . . . . . . . . . . . . . . . . . . . . . . . . . . . Porosity in Limestones and Dolomites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pore-Filling Processes: Cementation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Controls on Carbonate Cementation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Morphology and Fabrics of Cement Types . . . . . . . . . . . . . . . . . . . . . . . . . Cement Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cement Fabrics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cement Types and Diagenetic Environments . . . . . . . . . . . . . . . . . . . . . . . Facies-Controlled Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Platforms and Ramps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reefs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cold-Water vs. Warm-Water Diagenesis . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Pathways and Patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Textures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mechanical Processes: Compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Chemical Processes: Pressure Solution and Stylolitization . . . . . . . . . . . . . Significance of Compaction and Pressure Solution . . . . . . . . . . . . . . . . . . . Neomorphic Processes: Alteration and Recrystallization . . . . . . . . . . . . . . Recrystallized Carbonate Rocks: What to do? . . . . . . . . . . . . . . . . . . . . . . . How to DescribeRecrystallized Carbonate Rocks? . . . . . . . . . . . . . . . . . . . Sparite: Recrystallization Product or Carbonate Cement? . . . . . . . . . . . . . . Dolomitization and Dedolomitization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Descriptive Criteria and Terminology of Dolomite Fabrics . . . . . . . . . . . . .
267 267 267 270 271 272 272 272
263
272 274 276 276 276 277 277 278 278 278 278 282 282 283 288 289 292 295 295 299 299 301 301 302 308 308 310 311 314 321 321 321 324 324 325 328
XV
Contents
7.8.1.1 7.8.1.2 7.8.1.3 7.8.2 7.8.2.1 7.8.2.2 7.8.2.3 7.8.3 7.8.3.1 7.8.3.2 7.8.3.3 7.9
Thin-Section Description and Terminology of Dolomite Rocks . . . . . . . . . Dolomite Cement . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Value of Dolomite Textures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Some Dolomitization Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dolomites Associated with Evaporites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mixing-Water and Seawater Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Subsurface Burial Dolomites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dedolomitization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Textural Criteria for Recognizing Dedolomitization . . . . . . . . . . . . . . . . . . Origin of Dedolomite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Dedolomitization . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Metamorphic Carbonate and Marbles . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
328 329 330 330 330 331 332 332 333 333 334 334
8 8.1 8.2 8.3 8.3.1 8.3.2 8.3.2.1 8.3.2.2 8.3.2 8.3.2.1 8.3.2.2 8.4 8.4.1 8.4.2 8.5 8.6
Classification – A Name for Your Sample . . . . . . . . . . . . . . . . . . . . . . . . Basic Concepts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reef Limestones and Microbial Carbonates (Autochthonous Carbonates) . Classifications Based on Depositional Texture . . . . . . . . . . . . . . . . . . . . . . Prerequisites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Original and Expanded Dunham Classification . . . . . . . . . . . . . . . . . . . . . . Concepts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Original and Expanded Folk Classification . . . . . . . . . . . . . . . . . . . . . . . . . Concepts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Specific Classifications . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Changes in Depositional Textures . . . . . . . . . . . . . . . . . . . . . . . Some Nonmarine Carbonates Need very Specific Names . . . . . . . . . . . . . . Classification of Mixed Siliciclastic-Carbonate Rocks . . . . . . . . . . . . . . . . A Name for Your Samples: Some Practical Advice . . . . . . . . . . . . . . . . . . .
339 339 340 348 348 348 349 350 356 356 361 361 361 361 361 364
9 9.1 9.1.1 9.1.2 9.1.3 9.1.3.1 9.1.3.2 9.1.4 9.1.5 9.1.5.1 9.1.5.2 9.1.5.3 9.2 9.2.1 9.2.2 9.2.3
Limestones are Biological Sediments . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microbial Carbonates and Stromatolites . . . . . . . . . . . . . . . . . . . . . . . . . . . Bacterial Contribution to Carbonate Precipitation . . . . . . . . . . . . . . . . . . . . How to Recognize Microbial Carbonates? . . . . . . . . . . . . . . . . . . . . . . . . . . Describing and Classifying Benthic Microbial Carbonates . . . . . . . . . . . . . Terminology and Descriptive Criteria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Classification of Benthic Microbial Carbonates . . . . . . . . . . . . . . . . . . . . . Stromatolites are Laminated Microbialites . . . . . . . . . . . . . . . . . . . . . . . . . Occurrence and Significance of Microbialites and Stromatolites . . . . . . . . Development through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Paleoenvironmental Significance of Microbial Carbonates . . . . . . . . . . . . . Economic Importance of Stromatolites . . . . . . . . . . . . . . . . . . . . . . . . . . . . Biogenic Encrustations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Criteria and Constraints of Encrusters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Phanerozoic Encrusters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Encrustation Patterns in Recognizing Depositional Settings and Environmental Controls . . . . . . . . . . . . . . . . . . . Bioerosion, Boring and Grazing Organisms . . . . . . . . . . . . . . . . . . . . . . . . Recent and Fossil Microborers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Recent and Fossil Macroborers . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Micro- and Macroboring through Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . Qualitative Changes in Micro- and Macroborer Groups . . . . . . . . . . . . . . . Quantitative Changes in the Intensity of Macroboring in Coral Reefs . . . .
369 369 369 370 370 370 371 374 377 377 379 379 379 381 384
9.3 9.3.1 9.3.2 9.3.3 9.3.3.1 9.3.3.2
384 386 387 388 392 392 394
XVI
9.3.4 9.4
10 10.1 10.1.1 10.1.2 10.1.3 10.2 10.2.1 10.2.1.1 10.2.1.2 10.2.1.3 10.2.1.4 10.2.1.5 10.2.1.6 10.2.1.7 10.2.1.8 10.2.1.9 10.2.2 10.2.2.1 10.2.2.2 10.2.2.3 10.2.3 10.2.3.1 10.2.3.2 10.2.3.3 10.2.3.4 10.2.4 10.2.4.1 10.2.4.2 10.2.4.3 10.2.4.4 10.2.4.5 10.2.4.6 10.2.4.7 10.2.4.8 10.2.4.9 10.2.5 10.2.6 10.2.6.1 10.2.6.2 10.3 10.4
11 11.1 11.2 11.3
Contents
Microborer Associations are Proxies for Paleo-Water Depths . . . . . . . . . . . Practical Advice: How to Describe Microbialites and Stromatolites, Biogenic Encrustations and Borings? . . . . . . . . . . . . . . . . . .
396
Fossils in Thin Section: It is Not That Difficult . . . . . . . . . . . . . . . . . . . . Specifics of Thin-Section Fossils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . How to Determine Fossils in Thin Sections? . . . . . . . . . . . . . . . . . . . . . . . . Which Fossils in which Time Interval? . . . . . . . . . . . . . . . . . . . . . . . . . . . . Practical Advice . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagnostic Criteria of Fossils in Thin Sections . . . . . . . . . . . . . . . . . . . . . . Cyanobacteria and Calcareous Algae . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cyanobacteria and Calcimicrobes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Corallinacean and Peyssoneliacean Red Algae . . . . . . . . . . . . . . . . . . . . . . Solenoporacean Red Algae . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ancestral Red Algae and Problematic Red Algae . . . . . . . . . . . . . . . . . . . . Udoteacean Green Algae and Gymnocodiacean Algae . . . . . . . . . . . . . . . . Phylloid Algae . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dasyclad Green Algae . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Charophyta: Fresh-Water and Brackish-Water Algae . . . . . . . . . . . . . . . . . ‘Calcispheres’ and Algal Cysts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Foraminifera and other Protozoa . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Foraminifera . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Radiolaria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Calpionellids. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Benthic Sessile Organisms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sponges . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hydrozoans . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Corals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bryozoans . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Shells . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Brachiopods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Bivalves . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Gastropods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cephalopods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Tentaculitids and Other Conical Shells (Cricoconarida; Hyolithida) . . . . . . Serpulids and Other Worm Tubes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Crustacean Arthropods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Trilobites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Echinoderms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Rare Thin-Section Fossils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microproblematica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Inventory of Microproblematica . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion of Selected Microproblematica . . . . . . . . . . . . . . . . . . . . . . . . . Biozonation of Platforms with Thin-Section Fossils . . . . . . . . . . . . . . . . . . Where to Look for Thin-Section Pictures of Fossils and Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
399 399 399 400 401 402 404 408 412 418 422 422 427 430 447 452 453 453 482 487 491 491 508 509 514 518 519 522 529 533 537 540 542 547 548 558 559 559 560 570
Microfacies Interpretation Summarizing Microfacies Criteria: Microfacies Types . . . . . . . . . . . . . MFT Concepts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . How to Differentiate Meaningful Microfacies Types . . . . . . . . . . . . . . . . . Making Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
396
571
575 575 576 584
Contents
XVII
12 12.1 12.1.1 12.1.1.1 12.1.1.2 12.1.1.3 12.1.2 12.1. 2.1 12.1.2.2 12.1.2.3 12.1.3 12.1.3.1 12.1.3.2 12.1.4 12.1.4.1 12.1.4.2 12.1.5 12.1.5.1 12.1.5.2 12.1.5.3 12.1.6 12.1.6.1 12.1.6.2 12.1.7 12.1.7.1 12.1.7.2 12.1.7.3 12.1.8 12.1.8.1 12.1.8.2 12.1.8.3 12.2 12.2.1 12.2.2 12.2.2.1 12.2.2.2 12.2.2.3 12.3 12.3.1 12.3.2 12.4
Recognizing Paleoenvironmental Conditions . . . . . . . . . . . . . . . . . . . . . Reconstructing Environmental Constraints . . . . . . . . . . . . . . . . . . . . . . . . . Hydrodynamic Controls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hydrodynamic Energy Levels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Classifying Low-Energy and High-Energy Environments . . . . . . . . . . . . . Paleocurrent Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Storms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Storm Deposits (Tempestites) on Shelves, Ramps and Platforms . . . . . . . . Impact of Tropical Storms on Reefs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Carbonate Storm Deposits . . . . . . . . . . . . . . . . . . . . . . . . . Marine Carbonate Substrates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Substrate Types and Organism-Sediment Interactions . . . . . . . . Recognizing Substrate Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Light . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Zonation and Light Conditions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Recognition of Photic and Aphotic Conditions . . . . . . . . . . . . . . . . . . . . . . Oxygen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Terminology and Classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Recognizing Paleo-Oxygenation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Black Shale Development on a Carbonate Platform . . . . . . . . Seawater Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Seawater Temperature: Biotic Proxies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemical Proxies of Seawater Temperatures . . . . . . . . . . . . . . . . . . . . . Salinity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Biotic and Microfacies Proxies of Paleosalinity . . . . . . . . . . . . . . . . . . . . . Geochemical Proxies of Paleosalinity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Proxies for Hypersaline and Evaporitic Conditions . . . . . . . . . Productivity and Nutrients . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Nutrients . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Estimating Paleonutrient Levels . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Effects of Nutrient Excess on Reef and Platform Carbonates . . . . . . . . . . . Estimating Paleoclimatic Conditions: Grain Association Analysis . . . . . . . Concepts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Practical Advice, Examples and State of Current Information . . . . . . . . . . Distinguishing Grain Association Types . . . . . . . . . . . . . . . . . . . . . . . . . . . Examples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . State of the Art . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Assessing Water Depths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Hints to Paleowater Depths . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Assessing the Water Depth of a Carbonate Ramp . . . . . . . . . . Looking for Seismic Events . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
587 587 588 589 591 593 593 594 600 604 605 606 607 610 611 611 612 613 613 614 615 616 616 617 618 619 622 623 623 624 624 625 625 627 627 628 628 634 635 637 640
13 13.1 13.1.1 13.1.2 13.1.2.1 13.1.2.2 13.1.2.3 13.1.2.4 13.2 13.2.1 13.2.2
Integrated Facies Analysis . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Non-Carbonate Constituents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Insoluble Residues (IR): Clay Minerals and Detrital Quartz . . . . . . . . . . . . Authigenic Minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Silicification of Carbonates, Authigenic Feldspar and Glauconite . . . . . . . Sulfides: Pyrite . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sulfates: Evaporite Minerals . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Phosphates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Geochemical Proxies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Trace Elements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Strontium and Manganese - Favorite Tools for Facies Studies . . . . . . . . . .
641 641 641 643 643 646 647 648 652 652 652
XVIII
Contents
13.2.3 13.2.4 13.3
Significance of Trace Elements in Facies Studies of Carbonate Rocks . . . . Stable Isotopes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Organic Matter in Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
653 654 656
14 14.1 14.1.1 14.1.1.1 14.1.1.2 14.1.1.3 14.1.2 14.1.2.1 14.1.2.2 14.1.2.3 14.1.3 14.1.3.1 14.1.3.2 14.1.4 14.1.5 14.1.6 14.1.7 14.1.8 14.2 14.2.1 14.2.2 14.3 14.3.1 14.3.2 14.3.3 14.3.4 14.3.4.1 14.3.4.2 14.3.5 14.3.6 14.4
Depositional Models, Facies Zones and Standard Microfacies . . . . . . . Depositional Facies Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Conceptual, Dynamic and Computer Models . . . . . . . . . . . . . . . . . . . . . . . Conceptual Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Dynamic Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Numerical Models . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Basic Elements of Carbonate Facies Models . . . . . . . . . . . . . . . . . . . . . . . . Common Facies Belts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Common Depositional Patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Different Depositional Settings Require Different Facies Models . . . . . . . Facies Zones of Rimmed Carbonate Platforms: The Wilson Model . . . . . . Standard Facies Zones and the Modified Wilson Model . . . . . . . . . . . . . . . Discussion and Use of Standard Facies Zones . . . . . . . . . . . . . . . . . . . . . . . Carbonate Ramp Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Non-rimmed Shelves and Platforms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Isolated Platforms and Atolls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Epeiric Platform Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Epeiric Ramp Model . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Biotic Zonation Patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Concepts and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Foraminiferal Distribution in Late Triassic Reefs and Platforms Standard Microfacies Types (SMF) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Revised Standard Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Discussion of Standard Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . . . Stratigraphic Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Common Microfacies Types of Carbonate Ramps . . . . . . . . . . . . . . . . . . . Microfacies Criteria of Carbonate Ramps . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Types of Carbonate Ramps . . . . . . . . . . . . . . . . . . . . . . . . . . . Tracing Facies Zones with Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . Determining Standard Microfacies Types: A Practical Guide . . . . . . . . . . . Dynamic Microfacies Types and Environmental Changes . . . . . . . . . . . . .
657 657 657 657 658 658 659 659 659 660 660 660 661 664 667 667 669 669 671 671 676 680 681 712 716 716 716 718 718 720 723
15 15.1 15.1.1 15.1.2 15.2 15.2.1 15.2.2 15.3 15.4 15.4.1 15.4.1.1 15.4.1.2 15.4.1.3 15.4.2 15.5 15.5.1 15.5.1.1
Basin Analysis: Recognizing Depositional Settings . . . . . . . . . . . . . . . . . Pedogenic Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Criteria of Paleocaliche . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Significance of Paleocaliche and Paleosols . . . . . . . . . . . . . . . . . . . . . . . . . Paleokarst and Ancient Speleothems . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagnostic Criteria of Paleokarst and Paleospeleothem Structures . . . . . . . Significance of Paleokarst and Cave Carbonates . . . . . . . . . . . . . . . . . . . . . Travertine, Calcareous Tufa and Calcareous Sinter . . . . . . . . . . . . . . . . . . . Lacustrine and Palustrine Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies of Lacustrine Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Criteria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Types of Lacustrine Limestones . . . . . . . . . . . . . . . . . . . . . . . . Distribution of Lacustrine Microfacies Types . . . . . . . . . . . . . . . . . . . . . . . Palustrine Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Peritidal Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Criteria of Peritidal Limestones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Definitions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
725 725 726 728 730 731 734 734 737 738 738 740 740 742 745 745 745
Contents
15.5.1.2 15.5.1.3 15.5.2 15.6 15.6.1 15.6.2 15.6.2.1 15.6.2.2 15.6.3 15.6.3.1 15.6.3.2 15.6.3.3 15.7 15.7.1 15.7.2 15.7.2.1 15.7.2.2 15.7.2.3 15.7.2.4 15.7.2.5 15.7.2.6 15.7.3 15.7.3.1 15.7.3.2 15.7.3.3 15.7.3.4 15.7.4 15.7.5 15.7.5.1 15.7.5.2 15.8 15.8.1 15.8.2 15.8.2.1 15.8.2.2 15.8.3 16 16.1 16.1.1 16.1.1.1 16.1.1.2 16.1.1.3 16.1.2 16.1.2.1 16.1.2.2 16.1.2.3 16.2
XIX
Major Facies Criteria of Peritidal Carbonates . . . . . . . . . . . . . . . . . . . . . . . Synopsis of Diagnostic Criteria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Middle Devonian Peritidal Carbonates from Poland . . . . . . . . Carbonate Platforms and Ramps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ecological Controls on Platforms and Ramps . . . . . . . . . . . . . . . . . . . . . . . The Response of Carbonate Platforms to Drowning . . . . . . . . . . . . . . . . . . Microfacies Signals of Drowning History . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Platform Drowning Reflected by Microfacies . . . . . . . . . . . . . Case Studies: Platform and Ramp Carbonates . . . . . . . . . . . . . . . . . . . . . . . A Late Jurassic Bahamian-Type Carbonate Platform from the Northern Calcareous Alps . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . A Middle Devonian Ramp from Graz, Southern Austria . . . . . . . . . . . . . . . A Late Jurassic/Early Cretaceous Ramp from the Subsurface of Southern Bavaria, Germany . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Platform-Slope-Basin Transects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Types and Composition of Carbonate Slopes . . . . . . . . . . . . . . . . . . . . . . . Allochthonous Slope and Basin Deposits: Diagnostic Criteria . . . . . . . . . . Submarine Rockfalls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Breccias and Megabreccias . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Debris-Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Grain-Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Turbidites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sliding and Slumping . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies of Slope Carbonates: Case Studies . . . . . . . . . . . . . . . . . . . . . . Permian of Sicily: Megablocks and Base-of-Slope Carbonates . . . . . . . . . . Triassic of the Southern Alps: Allochthonous Slope Sediments . . . . . . . . . Jurassic of Morocco: Platform-Slope-Basin Transect . . . . . . . . . . . . . . . . . Jurassic of the Northern Calcareous Alps: Detailed Information from Limestone Turbidites on Source and Deposition Patterns . . . . . . . . . . . . . . Slope Stability Reflected by Texture and Microfacies . . . . . . . . . . . . . . . . . Tracing Platform-Basin Transitions Using Grain Composition Logs . . . . . Concept and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Late Triassic of the Gosaukamm Region, Austria . . . . . . . . . . Pelagic Deep-Marine Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Setting, Controls and Biota of Pelagic Limestones . . . . . . . . . . . . . . . . . . . Examples and Case Studies of Pelagic Carbonates . . . . . . . . . . . . . . . . . . . Microfacies of Paleozoic Basinal Carbonates . . . . . . . . . . . . . . . . . . . . . . . Microfacies of Mesozoic Basinal Carbonates . . . . . . . . . . . . . . . . . . . . . . . Contourites . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Realizing Depositional Constraints and Processes . . . . . . . . . . . . . . . . . Cyclic Carbonates, Microfacies and Sequence Stratigraphy . . . . . . . . . . . . Cyclic Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cyclic Carbonates: Some Basics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies and Cyclic Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Studies: The Lofer Cycle and the Latemar Cycle (Triassic of the Alps) . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Sequence Stratigraphy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Sequence Analysis: Some Basics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Data Applied to Sequence Stratigraphy . . . . . . . . . . . . . . . . . . Case Studies: Sea-Level Fluctuations and Systems Tracts Documented by Microfacies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Understanding Reef Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
746 749 752 753 753 756 756 757 759 760 768 768 769 769 771 771 772 772 773 773 779 780 780 781 784 786 786 788 788 789 793 793 795 795 796 801 803 803 803 804 805 808 815 816 818 822 830
XX
16.2.1 16.2.2 16.2.3 16.2.3.1 16.2.3.2 16.2.4 16.2.5 16.2.5.1 16.2.5.2 16.2.6 16.2.6.1
Contents
16.7.1 16.7.1.1 16.7.1.2 16.7.1.3 16.7.2 16.7.3 16.7.4
What is a Reef? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reef Types . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reef Fossils . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reef Biota: Compositional Changes during Time . . . . . . . . . . . . . . . . . . . . Reef Guilds: Ecologic Units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . How to Classify Reef Carbonates? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies Approach to Reef Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Basic Constituents of Reef and Mound Carbonates . . . . . . . . . . . . . . . . . . . Describing Reef Carbonates: A Practical Guide . . . . . . . . . . . . . . . . . . . . . Case Studies of Some Ancient Reefs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mud Mounds: The Early Devonian ‘Kess Kess’ Mounds in the Anti-Atlas, Southern Morocco . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Waulsortian Mud Mounds: Early Carboniferous (Lower Mississippian) Muleshoe Mound, Sacramento Mountains, New Mexico, U.S.A. . . . . . . . . Reefs: The Capitan Reef, Permian Reef Complex, Guadalupe Mountains, Texas and New Mexico, U.S.A. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Fingerprinting Lost Platforms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Platform Facies Patterns Derived from the Microfacies of Early Carboniferous Conglomerates (Southern Spain) . . . . . . . . . . . . . . . . Case Study: Reconstruction of Paleo-Escarpments from Microfacies Data Recognizing Ancient Cool-Water Carbonates . . . . . . . . . . . . . . . . . . . . . . . Microfacies Criteria of Non-Tropical Cold-Water Shelf and Reef Limestones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Case Study: Early Tertiary Cool-Water Coral Reef . . . . . . . . . . . . . . . . . . . Testing for Ancient Vent and Seep Carbonates . . . . . . . . . . . . . . . . . . . . . . Diagnostic Criteria of Ancient Seep and Vent Carbonates . . . . . . . . . . . . . . Case Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Late Eocene ‘Whiskey Creek’ Seep Carbonates of Washington State, U.S.A. Early Cretaceous Seep Carbonates in the Canadian Arctic . . . . . . . . . . . . . Mixed Carbonate-Siliciclastic Settings and Limestone/Marl Sequences . . . Carbonate-Siliciclastic Environments . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Modern Carbonate-Siliciclastic Environments . . . . . . . . . . . . . . . . . . . . . . Ancient Mixed Carbonate-Siliciclastic Environments . . . . . . . . . . . . . . . . . Describing Carbonate-Siliciclastic Sediments: Practical Advice . . . . . . . . . Limestone-Marl Sequences: Primary or/and Diagenetic Origin? . . . . . . . . Secular Variations in Carbonate Depositional Patterns and Temporal Changes in Microfacies Criteria . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Changes of Major Carbonate Depositional Environments During Time . . . Phanerozoic Carbonate Platforms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Phanerozoic Reef Patterns . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pelagic Carbonates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Differences in Phanerozoic Benthic Carbonate Factories . . . . . . . . . . . . . . Temporal Changes in Non-Skeletal and Skeletal Mineralogy . . . . . . . . . . . Temporal Changes in the Abundance and Significance of Microfacies Criteria
17 17.1 17.1.1 17.1.2 17.1.2.1
Practical Use of Microfacies Reservoir Rocks and Host Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate Hydrocarbon Reservoirs . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Distribution of Carbonate Reservoirs during Time . . . . . . . . . . . . . . . . . . . Depositional Setting and Environmental Controls of Carbonate Reservoirs Depositional Setting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
16.2.6.2 16.2.6.3 16.3 16.3.1 16.3.2 16.3.2.1 16.3.2.2 16.4 16.4.1 16.4.2 16.5 16.5.1 16.5.2 16.5.2.1 16.5.2.2 16.6 16.6.1 16.6.1.1 16.6.1.2 16.6.1.3 16.6.2 16.7
830 830 831 831 831 833 833 834 836 836 838 842 843 847 847 848 848 849 852 852 856 857 857 858 859 860 862 862 862 864 864 866 868 868 869 870 872 873 874 874 877 877 878 878 878
Contents
XXI
17.1.2.2 17.1.3 17.1.3.1 17.1.3.2 17.1.4 17.1.4.1 17.1.4.2 17.1.4.3 17.1.4.4 17.1.5 17.1.5.1 17.1.5.2 15.1.5.3 17.2 17.2.1 17.2.2
Environmental Controls . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Controls on Carbonate Reservoirs . . . . . . . . . . . . . . . . . . . . . . . Reservoir Properties . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Diagenetic Controls on Reservoir Properties . . . . . . . . . . . . . . . . . . . . . . . . Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Seismic Interpretation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Log Response . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Cores and Cuttings . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reservoir-Related Outcrop Analog Studies . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies, Lithofacies and Reservoir Rock Types . . . . . . . . . . . . . . . . . . Reservoir Heterogeneity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Relevant Microfacies Data . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Reservoir Rock Types and Facies Criteria . . . . . . . . . . . . . . . . . . . . . . . . . . Carbonate-Hosted Mineral Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ore Deposits and Carbonate Settings . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Microfacies and Ore Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
881 881 881 881 884 884 884 885 885 888 888 888 890 892 892 893
18 18.1 18.2 18.3 18.4 18.4.1 18.4.2
Carbonate Rock Resources, Facies, Weathering, Preservation . . . . . . . Industrial Use of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Exploration and Exploitation of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . Facies and Physical-Chemical Properties of Carbonate Rocks . . . . . . . . . . Weathering, Decay and Preservation of Carbonate Rocks . . . . . . . . . . . . . . Weathering of Carbonate Rocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Preservation . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
895 895 895 896 897 897 899
19 19.1 19.2 19.2.1 19.2.2 19.3 19.3.1 19.3.2 19.4 19.4.1 19.4.2 19.5 19.5.1 19.5.2 19.6 19.7
Microfacies and Archaeology . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Questions and Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Building Stones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Building Stones: Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Building Stones: Examples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mosaic Material . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mosaic Material: Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Mosaic Material: Examples . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Works of Art . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Works of Art: Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Works of Art: Example . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ceramics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ceramics: Methods . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Ceramics: Example . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Marble Studies . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Antony and Cleopatra: Tracing a Famous Love Affair . . . . . . . . . . . . . . . .
903 903 903 903 904 905 905 906 907 907 907 910 910 912 913 913
20
Adding Some Samples by Axel Munnecke . . . . . . . . . . . . . . . . . . . . . . . .
916
1 2 2.1 2.2 3
Appendix Answers to Exercises . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Looking at the Attached CD . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Comparison Charts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Using the List of References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Permissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .
925 927 927 927 927
...........................................................
929
Index
XXII
Plate going with the preface: Looking onto a vertically cut and polished slab of a Late Triassic reef limestone from Adnet, near Salzburg, Austria. The spectacular colors and patterns in polished limestone surfaces provide information about the origin of limestone – both its mode of accumulation and the subsequent diagenetic processes that have strongly influenced its final appearance. The limestone tells about the animals and plants which formed it and may still be preserved in it, about the depositional environment when the sediment forming the limestone was accumulating, and also about periods of non-deposition and loss of sediment. Many questions about a limestone can be answered in the field by studying it on the outcrop or on a cut or even polished quarry wall. For example, are the bush-like corals in the lowermost part of the picture in growth position or do the small rims of reddish infill indicate that the corals have been transported and overturned before final deposition? Is the smooth horizontal upper surface of the coral knobs due to strong erosion? Many of the thin-branched corals in the center of the picture seem to be broken and most are overturned; were they transported or did they grow here or near-by? What do we know about the influence of storms or of an unstable sea floor that may be part of the answer to this question? What can we tell from this slab about percolating fluids that destroyed all or most internal structures of the corals? What is the significance of reddish clay that fills cavities, broken shells and, in the upper part of the picture, complete shells? Is the blackish-gray, irregular line that crosses the picture an indication of a time of non-deposition? The light gray sediment at the top of the picture indicates a completely different environment with rare corals and less common debris. Many questions can be posed and answered from the study of the quarry wall as here illustrated by the polished slab in the Geozentrum Nordbayern (Paleobiology), Erlangen, but many questions remain unanswered by analysis at this scale. This slab illustrates the bridge between outcrop and thin-section scale, contributes essential and unique data about the depositional environment, sedimentation, and the processes that transformed the sediment into limestone. Late Triassic (Norian) limestone form the Tropf-Steinbruch in Adnet near Hallein (Salzburg, Austria). Width of picture is 1 m. Photo courtesy of Ch. Schulbert
Plate Introduction
XXIII
1
1 New Perspectives in Microfacies
A rapid evolution in our understanding of carbonate rocks was triggered in the mid- to late 50’s by the discovery of carbonate reservoirs in various parts of the world, followed by intensive research of recent carbonates. Modern and ancient carbonate environments, diagenetic processes and facies models were studied between about 1955 and 1965. During the late 60’s and the 70’s microfacies became an essential part of facies analysis and paleoenvironmental interpretation of limestones. The increasing importance of limestones and dolomites as reservoir rocks and the use of thin-section fossils in subdividing carbonate platforms gave substantial impetus to the progress of microfacies research. The 1980’s were characterized by an increased application of geochemical techniques with the object of developing predictive models for carbonate diagenesis and porosity. The subsequent period of time has seen the rapid development of sequence stratigraphy which now plays a major role in the characterization of the geometry and reservoir potential of carbonate rocks.
as 1927 by the use of microscopic features of limestones for oil exploration in Texas. Hovelaque and Kilian (1900) published the first illustrated volume of thin-section photographs of carbonates. This approach was continued by the ‘International Sedimentary Petrographical Series’ initiated by Cuvillier at the Third International World Petroleum Conference in Paris in 1951. The volumes published within this series as well as other comprehensive monographs containing numerous plates with thin-section photographs (see Sect. 10.4) contributed substantially to the rapid adoption of the microfacies approach. During the past decades microfacies has become an established part of the study of carbonate rocks. However, many authors use microfacies criteria in describing and classifying limestones but fail to explore the significance of these criteria for the interpretation of depositional and diagenetic histories of carbonate rocks. Part 2 and 3 of this book demonstrate the great potential of microfacies studies in basin analysis and applied carbonate geology.
1.1 The Microfacies Concept 1.2 New Perspectives As originally defined by Brown (1943) and again independently by Cuvillier (1952) the term ‘microfacies’ referred only to petrographic and paleontological criteria studied in thin-sections. Today, however, microfacies is regarded as the total of all sedimentological and paleontological data which can be described and classified from thin sections, peels, polished slabs or rock samples. Field geology, including mapping and profiling, is a prerequisite for successful microfacies analysis. The importance of combined field work and thin-section studies was already emphasized in the earliest thin-section based investigations of carbonate rocks directed towards genetic interpretations of limestones (Peters 1863; Gümbel 1873; Sorby 1879; Hantken 1884). The practical application of thin-section criteria of limestones was demonstrated by Udden and Waite as early
New perspectives in the study and application of microfacies have resulted from the application of new techniques, from the availability of new comparative data on modern and ancient carbonates, from new data concerning the origin and distribution patterns of carbonate grains and the dominant biological control on carbonate sedimentation, from new concepts on defining facies models, and from new ideas concerning the course of carbonate sedimentation through geological time. New techniques. Whilst microscopy is essential in the study of carbonate rocks there are many additional techniques which should also be utilized to maximize the information offered by microfacies analysis (Tucker 1988). Cathodoluminescence microscopy, fluid inclusion microscopy, scanning electron microscopy as well
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_1, © Springer-Verlag Berlin Heidelberg 2010
2
New Perspectives in Microfacies
Fig. 1.1. Limestones as paleoclimatic proxies. Shelf limestones formed in tropical warm-water and in non-tropical coolwater environments differ in the association of specific grain types (see Sect. 12.2). The sample is a skeletal grainstone consisting predominantly of bryozoans (B) and benthic foraminifera (F). Gastropods (G) are rare. Black grains are mainly red algae. The skeletons of many bryozoan fragments have lost their structure due to the attacks of microborers and have been transformed into microcrystalline grains. The sediment is only poorly lithified as shown by the high open porosity. Marine carbonate cement occurs as thin rims of fibrous calcite crystals (arrows). The grain association as well as the weak cementation are characteristic of cool-water carbonates. The sample comes from Oligocene deposits of Leg 133 of the Ocean Drilling Program which dealt with the geological evolution of the Great Barrier Reef Province off northeastern Australia. Courtesy of T. Brachert (Mainz).
as stable isotope geochemistry have become common tools in sedimentological studies (Chap. 3). Chemical analysis and X-ray techniques are well-established routine methods used to determine major, minor and trace element concentrations and the composition of carbonate rocks. The combination of these techniques with microfacies studies offers new possibilities for the interpretation of diagenetic pathways of carbonates (Chap. 7). Quantification of thin-section criteria, based on computer-assisted frequency analysis and using multivariate methods, has become an essential prerequisite for the categorization of microfacies types, the evaluation of the sedimentary budget of basin fills and for computer simulations having the potential to constrain quantitative sedimentary models (see Sect. 6.3.2). Lacustrine and terrestrial carbonates. Despite the increasing economic importance of non-marine sediments, many students interested in microfacies are somewhat hesitant in applying microfacies analysis to
non-marine carbonate rocks. This attitude may reflect classification problems related to the poverty in characteristic non-marine microfacies types as compared with the wealth of types known from marine environments. This hesitancy is, however, only partly justified: there exist succinct and informative reviews of modern and ancient fresh and saltwater lake deposits, travertines and pedogenic carbonates (see Sect. 2.4.1 and Sect. 15.1 to Sect. 15.4). Ancient calcretes, characterized by a set of distinctive microfacies criteria, are very important indicators of subaerial exposure and are potentially very useful for defining the nature of sequence boundaries and recognizing changes in paleoclimate. The same is true for paleokarst which is easily recognized from microfacies features. New comparative data of modern and ancient sediments. Modern carbonate depositional systems only partly reflect the wide range of settings and environments established during the earth’s history because biological, geological and chemical controls have
New Perspectives in Microfacies
3
Fig. 1.2. Limestones as potential reservoir rocks. The sample has been impregnated with blue-dye-resin that fills open pores between and within skeletal grains (miliolid foraminifera, MF; bryozoans, B). Primary inter- and intra-particle porosity has been somewhat reduced by syntaxial overgrowth cements (O) on echinoderm fragments, dogtooth and granular cements within foraminiferal shells. Highly porous limestones may be good reservoir rocks depending on their permeability, and the geometry and distribution of pores. Tertiary: Paris Basin, France.
changed with time. Modern carbonates, however, demonstrate how organisms are involved in carbonate production and which physical, biological and chemical processes may be recorded in ancient carbonate rocks. Current microfacies interpretations are strongly influenced by new data both on shallow- and deep-marine sedimentation (see Chap. 2). Carbonate platforms and ramps. Facies interpretations for ancient shallow-marine carbonates have been heavily dependent on comparisons with sedimentation patterns of the Great Bahama Bank, the Florida shelf and parts of the Persian Gulf. Whilst ‘platform carbonates’ have previously been regarded as the normal type of shallow-marine sedimentation, case studies of ancient shallow-marine carbonates together with the investigation of modern examples now indicate that gently sloping carbonate ramps were also of major importance during the Phanerozoic. Ramps and platforms differ in their geometry, depositional depths and the distribution patterns of facies zones: microfacies reflect these differences (Sect. 14.3). Platform and ramp carbonates are controlled by variations in biogenic production as well as by fluctuations in both sea level and in accommodation and sedimentation rates. Microfacies
may reflect short-term environmental changes and highfrequency sea-level fluctuations as well as long-term patterns in the formation of carbonate buildups (Sect. 15.6). Platform-basin relations are recorded in allochthonous carbonates formed in shallow-water environments and deposited on the slope and within the basins (Sect. 15.7.5). The reconstruction of ‘vanished’ shallow-marine depositional environments is a very promising tool for microfacies studies (Sect. 16.3). Warm-water, temperate-water and cold-water carbonates. The paradigm of a predominantly warm-water origin for ‘sun-born’ shallow-marine carbonates has become obsolete as a result of current investigations of temperate, boreal, subarctic and even polar shelf carbonates (Sect. 2.4.4.3; Sect. 16.4). High-latitude marine carbonate production, cold-water organic reefs, deep-marine seep and vent communities (Sect. 16.5) as well as high bioclastic sedimentation rates in cold ocean waters are factors which need to be considered in the environmental interpretation of ancient carbonates. Increasing numbers of ancient examples of these types of carbonates are being recognized. The ability to distinguish ancient ‘tropical’ warmwater, and ‘non-tropical’ temperate and cold-water car-
4
A First Glance at a Thin Section
bonates and to recognize ‘warm-mode’ and ‘cold-mode’ paleolatitudes are improving rapidly. Carbonate grains. Increasingly new more information is available concerning the biological and non-biological controls on the origin of carbonate grains, the site of deposition of carbonate grains, and on the distribution and dominance of grain types during the Phanerozoic. Characteristic grain type associations and the specific composition of grain types (e.g. ooids, oncoids, pisoids) provide hints on paleolatitudes and environmental controls for both marine and non-marine carbonate deposition (Sect. 4.2; Sect. 12.2).
Bioerosion. The biological destruction of carbonate substrates by borers, raspers and crushers is currently being intensively studied. Important objectives of these studies, which are also significant in microfacies analysis, are the application of boring morphology and boring patterns for paleobathymetric reconstructions and the estimation of the quantitative contribution of bioeroders for the marine carbonate budget (Sect. 9.3). Microbial carbonates: Carbonate deposits produced or localized by microbial communities are known from marine, marginal-marine, freshwater and terrestrial environments. The role of bacteria and cyanobacteria
Plate 1 A First Glance at a Thin Section The plate exhibits a thin section of a Mesozoic limestone from the Austrian Alps. Although the first appearance may be puzzling, a thorough investigation of the microfacies criteria reveals an amazing history for this rock. 1 The thin section shows fossils and dark-colored areas. The fossils correspond to organisms once living in a reef and contributing to the formation of a reef limestone. The dark-colored areas represent various types of fine-grained sediment. The reef limestone contains calcareous sponges, serpulid worms (SW), tube-like microproblematica (Microtubus, M), and bivalves. Sponges and encrusting Microtubus and serpulids formed an organic framework. The sediment within the framework (S1) consists of tiny, densely packed microcrystalline particles (peloids) and some ostracods and foraminifera. Both elements can only be seen under greater magnification. The sponges are represented by non-chambered inozoans (IS), chambered sphinctozoans (SS) and chaetetid sponges (CS). Microtubus occurs encrusted on and within sponges, and is concentrated near the periphery of the framework, but the outermost part is a microbial crust (MC) formed by bacterially induced carbonate precipitation. The surface of the microbial crust outlines a distinct relief. The eye-catching large bivalve shell (BI) was heavily attacked by boring organisms including boring bivalves (BB) and microborers drilling tiny holes (MB). The bivalves are infilled with dark calcareous sediment (S2) different in composition from the sediment within the organic reef framework. The conspicuous circular section might represent a pelagic fossil (H, Heterastridium). The reddish fine-grained sediment (S3) yields densely packed skeletal grains including sponge spicules and pelagic shell debris as well as some radiolarians and pelagic ostracods. This sediment and the reef limestones are covered by a thin Fe/Mn crust (arrows) forming small microbial microstromatolites. The overlying sediment S4 contains (white) skeletal elements of crinoids, mostly parts of arms associated with pelagic shell debris. Radiolarians, small ammonites and pelagic algal cysts are too small to be visible on the plate. Interpretation: Microtubus indicates a Late Triassic (Norian to Rhaetian) age for the reef limestone. The composition of the reef-building assemblage characterizes reefs formed in protected low-energy settings. Environmental change is indicated by the peripheral microbial crust covering the reef fabric. The reef limestone and the microbial crust were transected by a fissure that remained open for some time as indicated by the calcite tapestry. Later on the fissure was infilled by the open-marine pelagic sediment S3. The Fe/Mn crust covering both the surface of the microbial crust and the fissure fill marks a discontinuity surface corresponding to a hardground and indicating a gap in the sedimentary record. The microfacies of sediment S4 reflects the renewed onset of pelagic sedimentation. The sediments S3 and S4 are Liassic in age. The time hidden in the pelagic sedimentary record, however, can not be determined from the thin-section data. Solution took place within the reef limestone as shown by solution voids (V) within the originally aragonitic sponges. In contrast primary calcitic fossils, e.g. Microtubus, are well preserved. These voids and the cement type of the calcite tapestry within the fissure point to freshwater diagenesis, an interpretation that also is supported by stable isotope data. The small-scale microfacies data of the thin section reflect the large-scale developments in sedimentation at the TriassicJurassic boundary in the Northern Calcareous Alps: Platform and reef development – platform destruction related to subaerial exposure and drowning – transition to open-marine pelagic sedimentation. The sample comes from the Tropfbruch quarry in Adnet near Salzburg, Austria, where Rhaetian reef limestones (Fig. 8.2) are overlain by Early Jurassic carbonates. See also the polished palte of this outcrop on page XXIII.
The story behind this thin section is rather difficult, but the following chapters will answer to all of your questions and give the information needed.
Plate 1: A First Glance at a Thin Section
5
6
in the precipitation of fine-grained carbonate and in the formation of microbialites is becoming increasingly well understood. Laboratory experiments simulating bacterially-controlled carbonate precipitation, the recognition of bacterially constructed carbonate crystals and grains and the study of modern biofilms combine to indicate the importance of microbes in the formation of many items included within the list of microfacies criteria (e.g. micrite, peloids, ooids, oncoids, micrite envelopes, stromatactis, carbonate crusts; Sect. 9.1). Combined studies of microfacies and geobiochemical data are necessary in order to understand constructive microbial controls (e.g. biogenic crusts and mud mounds), destructive processes (e.g. boring microbes) and carbonate cementation. Carbonate sequences and cycles. The recognition and interpretation of cyclic patterns is a major objective of modern carbonate sedimentology. Because cyclicity is reflected by systematic changes in biota, grain composition, texture and early diagenetic criteria, microfacies studies are able to contribute to a better understanding of short-term depositional variations (Sect. 16.1.1). The three major prerequisites for sequence stratigraphy – a sound paleoenvironmental interpretation, an estimate of synchronicity, and the differentiation between regional and local effects – may be vigorously supported by microfacies studies which also improve the interpretation of sequence boundaries with regard to origin, timing and location (Sect. 16.1.2). New concepts in defining facies models. The value of ‘models’ in facies analysis depends largely on the significance attributed to a sedimentary model by the author. The majority of models suggested for the different modes of carbonate sedimentation are based on comparative studies between modern and ancient carbonates. These ‘conceptual models’ remain important tools because they facilitate the attribution of characteristic sedimentological or paleontological data to specific facies belts (Chap. 14). However, a too rigorous use of conceptual models for one’s own facies studies may result in significant errors if the more descriptive, static character of the model is not critically considered. ‘Dynamic models’, concentrating on processes and controls and computer simulations dealing with the major controlling factors of carbonate buildups, provide a promising platform for the discussion of potential steps in the development of sedimentary bodies (e.g. carbonate platforms), organic structures (e.g. reefs) or individual sequences (e.g. cycles). The evaluation of computer-generated facies models should be strongly de-
A First Glance at a Thin Section
pendent on microfacies and paleontological criteria because only these criteria deliver the paleoenvironmental information that can verify, refine or disprove theoretical concepts. Secular changes during the Phanerozoic. There is considerable evidence for temporal and secular fluctuations in major facies types, the dominant types of carbonate cements, the mineralogy of skeletal grains and probably also in diagenesis and carbonate geochemistry. Some of these fluctuations can be explained by global changes in biological factors (e.g. role of carbonate plankton, nutrients), others by global changes in climate and oceanography. Microfacies criteria for carbonate rocks provide evidence for changes through time as shown by distinct differences in the composition of Phanerozoic bioclastic sands, types of calcareous ooids, composition and size of oncoids, open-space structures etc. (Sect. 16.7). The recognition of these changes has a major impact on depositional and diagenetic models for ancient carbonates. Applied microfacies. Carbonate rocks contain more than 50% of the world’s oil and gas reserves, are important hosts to ore deposits and form very important raw materials. Depositional facies and diagenetic patterns determine the physical and chemical properties of these limestones and dolomites which control their reservoir potential (Chap. 17) and their technological attributes (Chap. 18). Microfacies studies assist in understanding the origin and history of these characteristics. Basics: New perspectives in microfacies Brown, J.S. (1943): Suggested use of the word microfacies. – Economic Geology, 38, p. 325 Carozzi, A.V. (1989): Carbonate rock depositional models. A microfacies approach. – 604 pp., Englewood Cliffs (Prentice Hall) Cuvillier, J. (1952): Le notion de ‘microfacies’ et ses applications. – VIII Congreso Nazionale di Metano e Petroleo, sect. I, 1-7 Cuvillier, J. (1962): Étude et utilisation rationelle de microfacies. – Revue de Micropaléontologie, 4, 3-6 Flügel, E. (1982): Microfacies analysis of limestones. – 633 pp., Berlin (Springer) Spence, G.H., Tucker, M. (1999): Modelling carbonate microfacies in the context of high-frequency dynamic relative sea-level and environmental changes. – Journal of Sedimentary Research, 69, 947-961 Udden, J.A., Waite, V.V. (1927): Some microscopic characteristics of the Bend and Ellenburger limestones. – Texas University Bulletin, 27, 8 pp. Wilson, J.L. (1975): Carbonate facies in geologic history. – 471 pp., Berlin (Springer) Further reading: K002
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2 Carbonate Depositional Environments
Knowing where modern carbonates occur, what they are composed of, and what their controls are is essential for evaluating microfacies data. The objectives of this chapter are to summarize the settings and environments in which carbonate sediments are formed and to document which classifications are used for differentiating these environments. Emphasis is placed on the definitions of terms. The potential of microfacies analysis in recognizing and interpreting the environments discussed in this chapter will be demonstrated later. If you want to know how to describe paleoenvironmental criteria, and which microfacies criteria can be used in deciphering the biological and depositional constraints on ancient limestones, have a look at Chap. 12 and 15.
2.1 Introduction Microfacies analysis of carbonate rocks requires a knowledge of modern carbonates and an understanding of biological and geological changes during earth history. R.N. Ginsburg’s eloquent statement ‘carbonate buildups are like Shakespeare; the plays go on only the actors change’ emphasizes the remarkable similarities and analogues in the criteria of modern and ancient carbonates. The aphorism, however, ignores the guiding role of the stage-managers who often decide the basic mode of the play, similar to the changing geological and biological earth system parameters which shape the appearance of carbonate buildups.
2.1.1 Carbonates are Born not Made This simple phrase by Noel James (1979) highlights the main theme of carbonate sedimentation and the differences between carbonate and siliciclastic sediments. Carbonates are ‘born’. They originate as skeletal grains or precipitates within the depositional environment. By contrast, terrigenous clastic sediments are formed primarily by the disintegration of parent rocks and are transported to the depositional environment.
Shallow-marine carbonate depositional systems differ fundamentally from siliciclastic depositional systems in their intrinsic controls, i.e. the growth potential including the ‘aggregation potential’ (the ability to grow vertically and to track sea level) and the ‘production potential’ (the ability to produce and export sediment). The aggradation potential is critical for the history of carbonate platforms and reefs, the production potential for the progradation and retreat of platforms. These differences are of central importance in the application of sequence stratigraphy to carbonates (Schlager 1992). 2.1.2 The ‘Sorby Principle’: Limestones are Predominantly Biogenic Sediments Well over 90% or more of the carbonates found in modern marine environments are biological in origin, i.e. the sediments are biotically induced (by an organic trigger, e.g. microbial micrites) or biotically controlled (skeletal autotrophic and heterotrophic organisms determine the composition, location and timing of carbonate production). Some of the ‘abiotic’ carbonate precipitation (represented by marine cements) is also triggered by organics or the activity of organisms. The dominant role of organisms in the formation of limestones was recognized as early as 1879 by Henry Clifton Sorby. Studying Paleozoic, Mesozoic and Tertiary carbonate rocks in thin sections, he recognized the overwhelming abundance of fossils and their importance in the composition of carbonate sands and muds. The distribution and frequency of carbonate-producing organisms depend strongly on environmental factors, such as light, water temperature and sedimentary influx. These controls as well as the paleoenvironmental settings are reflected by microfacies criteria and paleontological data. 2.1.3 Modern Carbonates: Obligatory Reading There is a wealth of well-written reviews and concise updated syntheses of modern carbonate sediments (see
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_2, © Springer-Verlag Berlin Heidelberg 2010
Marine Environments
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Fig. 2.1. Settings of carbonate depositional environments.
list at the end of this section). The strength of these papers is that they outline the main criteria of recent marine and non-marine carbonates and their major controls in the context of interpreting ancient carbonate rocks. The information included in these fundamental texts is crucial for microfacies analyses. The ‘Microfacies of Carbonate Rocks’ does not intend to duplicate this information, but focuses on those data which are specifically relevant for microfacies studies.
2.2 Carbonate Sediments Originate on Land and in the Sea Carbonate sediments originate on land and in the sea. They are formed in three major settings: On the continents, within the transitional area between land and sea, and in the shallow and deep sea. Today only around 10 % of marine carbonate production takes place in shallow seas. 90% of the modern carbonate production is related to the deposition of calcitic plankton in the deep sea. These proportions were very different during most parts of the Phanerozoic. Major depositional environments are shown in Fig. 2.1. Terrestrial and aquatic non-marine carbonates are receiving increasingly more attention from geologists because of their economic importance and their significance in paleoenvironmental analyses. These carbonates are discussed in Sect. 15.1 to 15.4. Owing to their very shallow-water setting, marginal-marine carbonates have been studied in great detail and offer a wealth of information on coastal sedimentation processes and diagenesis. About 70% of microfacies studies dealing with marine carbonates concerns shallow-marine carbonates formed on the shelf and near the shelf break. Future research should be focused on the huge deep-marine carbonate depositional regions on the continental slopes.
2.3 Classification of Marine Environments Marine environments are classified into the benthic, for the sea bottom, and the pelagic, pertaining to the water mass. There is no universally accepted scheme of subdivision of marine environments which is equally acknowledged by biologists, oceanographers and geologists. Fig. 2.2 summarizes categories that are frequently used. The categories are predominantly based on schemes proposed by Hedgpeth (1957) and Edwards (1979), who discuss marine terminology. 2.3.1 Boundary Levels Several levels at the sea bottom and within the water column are commonly used in a vertical subdivision of marginal-marine and marine environments (Box 2.1). Microfacies analysis provides the basic data for recognizing these levels in ancient carbonate rocks (see Chap. 12). Essential critical interfaces that control sedimentary patterns and the distribution of organisms are: (1) The lower and the upper boundaries of the tides (control the distribution of organisms), (2) the base of the photic zone (controls the distribution of light-dependent phototrophic organisms), (3) the base of the zone of wave abrasion (above which bottom currents and wave action may lead to erosion and cementation), (4) the base of the action of storms on the sea bottom, (5) the O2 minimum zone (strongly limiting life on and in the sea bottom), (6) the thermocline (the layer of water that is too cold for most carbonate-producing organisms), and (7) the pycnocline (the layer of water where salinity is too high for most organisms). High and low tide, wave base and storm wave base are used as basic boundaries in the classification of the major shallow-marine environments.
Marine Zonations
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Fig. 2.2. Marine depositional environments. Modified from Kennett (1982).
2.3.2 Vertical and Horizontal Zonations Subdivisions of marine environments are both vertical and horizontal:
2.3.2.1 Vertical Zonations Benthic depth zones: The depth of the sea bottom and critical levels controlling the sedimentation subdivide the benthic environments into six zones: (1) coastal sublittoral zone (above high tide, corresponding to the supratidal zone), (2) littoral (between high and low tide, identical with the intertidal zone), (3) sublittoral (below low tide, corresponding to the major part of the
continental shelf), (4) bathyal (approximately equal to the continental slope), (5) abyssal (corresponding to the abyssal plains), and (6) the hadal zone (deep-sea trenches). Modern carbonate sedimentation takes place within the range of zone 1 to parts of zone 5. –> Geologists are in favor of the terms supratidal, intertidal and subtidal instead of the ecologically defined terms supralittoral, littoral and sublittoral. Pelagic depth zones: Five zones are defined by the vertical distribution of floating and swimming life. These are, in descending order, the epipelagic zone (the upper region of the oceans, extending to a depth of about 200 m), the mesopelagic, bathypelagic, abyssopelagic
Box 2.1. Glossary of terms used in the subdivision of marine environments. Fair-weather wave base (FWWB): Intersection of the wave base with submarine topography. The water depth below which surface wave action no longer stirs and moves the sediment. The depth of the wave base varies widely, depending on wave amplitude and fetch, bottom topography, activity of storms, orientation of shelves and ramps, latitudes etc. High tide: High water. The maximum height reached by each rising tide. Low tide: Low water. The minimum height reached by each falling tide. Mean sea level: The average height of the surface of the sea for all stages of the tide for a defined period. Storm wave base: See storm wave weather base. Storm wave weather base (SWWB): The water depth down to which storm-generated waves and resulting bottom currents rework bottom sediments and produce specific texture types (tempestites). The depth of the storm wave base varies strongly and depends on shelf and shelf edge topography, storm intensity and climate zone. Surf: a) The wave activity in the area between the shoreline and the outermost limit of breakers. b) A collective term for breakers. Tidal range: The difference in height between consecutive high and low waters. Tide: The periodic rise and fall of the sea resulting from the gravitational attraction of the moon and the sun acting upon the rotating earth and causing flood and ebb currents. Wave base: See fair-weather wave base.
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(corresponding to oceanic environments below about 4 000 m), and hadopelagic zone. –> Note that the boundaries of the benthic and pelagic depth zones are not fixed accurately. These boundaries reflect the situation in modern oceans that are not necessarily equivalent to depth zones visualized for ancient oceans.
2.3.2.2 Horizontal Zonations The lateral distribution of pelagic organisms with respect to their distance from the coast characterizes two major zones of the ocean: The neritic zone is the water that overlies the continental shelf, today generally with water depth less than 200 m and covering about 8% of the ocean floor. The term oceanic zone refers to the water column beyond the shelf break, overlying the slope and the deep-sea bottoms, generally with water depths greater than 200 m and down to more than 10 000 m.
Depositional Environments
Rao, C.P. (1996): Modern carbonates. Tropical, temperate, polar. – 206 pp., Howrah (Carbonates) Reading, H.E. (1996): Sedimentary environments, processes, stratigraphy. Third edition. – 688 pp., Oxford (Blackwell) Scholle, P.A., Bebout, D.G., Moore, C.H. (eds., 1983): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 708 pp. Scoffin, T.P. (1987): An introduction to carbonate sediments and rocks. – 274 pp., Glasgow (Blackie) Tucker, M.E., Wright, V.P. (1990): Carbonate sedimentology. – 482 pp., Oxford (Blackwell) Varney, M. (1996): The marine carbonate system. – In: Summerhayes, C.P., Thorpe, S.A. (eds.): Oceanography. An illustrated guide. – 182-194, New York (Wiley) Walker, R.G., James N.P. (eds., 1992): Facies models. Response to sea level change. – 409 pp., Ottawa (Geol. Ass. Canada) Wilson, J.L. (1975): Carbonate facies in geologic history. – 471 pp., Berlin (Springer)
2.4 Review of Modern Carbonate Depositional Environments
–> Again, keep in mind that these water depths are not compatible with the situation in many ancient oceans. Note that the term ‘neritic’ is often used to describe sea bottom environments below the neritic water column, or shallow-marine environments characterized by significant terrigenous influx.
The following review is focused on the definitions, main criteria and subdivisions of carbonate settings. The criteria that help recognize these settings and the evaluation of paleoenvironments in ancient carbonate rocks are discussed in Chap. 12 and 15.
Basics: Classification of carbonate environments and obligatory reading
2.4.1 Non-Marine Carbonate Environments
Classification Edwards, A.R. (1979): Classification of marine paleoenvironments. – Geol. Soc. New Zealand Newsletter, 48, p. 48 Hedgpeth, J. (1957): Classification of marine environments. – In: Hedgpeth, J.E. (ed.): The treatise on marine ecology and paleoecology. Vol.1, Ecology. – Geol. Soc. America Mem., 67, 17-28 Kennett, J. (1982): Marine geology. – 813 pp., Englewood Cliffs (Prentice Hall) Nybakken, J.W. (1993): Marine biology. An ecological approach. Third edition. – 462 pp., New York (Harper Collins) Summerhayes, C.P., Thorpe, S.A. (eds., 1996): Oceanography. An illustrated guide. – 352 pp., New York (Wiley) Obligatory reading Bathurst, R.G.C. (1994): Carbonate sediments and their diagenesis. 2nd enlarged edition – Developments in Sedimentology, 12, 660 pp., Amsterdam (Elsevier) Bosellini, A. (1991): Introduzione allo studio delle rocce carbonatiche. – 317 pp., Ferrara (Italo Bovolenta) James, N.P., Kendall, A.C. (1992): Introduction to carbonate and evaporite facies models. – In: Walker, R.G., James N.P. (eds.): Facies models. Response to sea level change. – 265-275, Ottawa (Geol. Ass. Canada) Milliman J.D. (1974): Marine carbonates. – 375 pp., Berlin (Springer)
Non-marine carbonates originate in terrestrial and aquatic environments without marine influence. These carbonates, formed by abiotic and/or biotic processes differ in many aspects from marine carbonates, which are much better known to geologists. These carbonates are formed in terrestrial subaerially exposed settings and in submerged aquatic settings (Fig. 2.3). The major non-marine carbonate environments are listed in Box 2.2. Non-marine carbonates formed in terrestrial subaerially exposed settings comprise pedogenic carbonates, palustrine carbonates, cave carbonates, eolian carbonates and glacial carbonates.
2.4.1.1 Pedogenic Carbonates, Paleosols, and Caliche/Calcretes Terminology: The term pedogenic refers to soil genesis. Most pedogenic carbonates are formed within soils rich in calcium carbonate. The term paleosol describes ‘a buried soil horizon of the geologic past’ but is often
Non-Marine Environments
11
Fig. 2.3. Non-marine carbonate depositional environments.
used in a broader sense, embracing various descriptive terms associated with horizons of subaerial exposure (e.g. caliche). Important processes are chemical and physical weathering of host carbonate rock, soil development and accumulation of calcium carbonate within the soil by evaporation of saturated porewaters driven to the surface by capillary forces. These processes may result in a physicochemically and microbially controlled formation of hard crusts on the surface or in the upper horizon of soils (duricrusts). Depending on climatic conditions and the host substrate, the duricrusts consist of carbonate (forming calcretes or dolocretes) or other mineral phases (e.g. silica forming silcretes, gypsum –> gypscretes, hematite and goethite –> ferricretes). The following discussion is limited to near-surface accumulations of calcium carbonate (Low-Mg calcite) in terrestrial, alternating wet settings, generally called calcrete or caliche. These distinctive carbonates result
from multiple processes induced by life, climate and exchange between mineral phases: ‘Calcrete’ was proposed for calcareous-cemented deposits in Africa, the roots of the word caliche are Latin, meaning lime and limestones (Reeves 1976). Calcrete and caliche are used synonymously or separately (calcrete in order to describe highly indurated parts of caliche profiles). Other names are calcareous duricrust, cornstone, kalkar (Agarwal et al. 1992) or ‘Krustenkalk’. These calcium–rich soil accumulations cover an estimated 20 million km2 or about 13% of the total land surface (Yaalon 1988). They are widespread in temperate (e.g. Mediterranean), semiarid and arid climates, particularly in tropical and subtropical latitudes (e.g. Caribbean) with sparse rainfall and a mean annual temperature of about 18 °C. ‘Calcicretes (syn. calcretes)’ and ‘dolocretes’ can be differentiated.
Box 2.2. Carbonates originating in non-marine environments. Non-marine carbonates formed in terrestrial subaerially exposed settings Pedogenic carbonates, paleosols, caliche/calcrete: Formed by the accumulation of calcium carbonate within unconsolidated carbonate-rich soils. Palustrine carbonates: Carbonates formed in lacustrine and transitional non-marine/marine environments originating by short-term oscillations of the water levels and characterized by a mixture of subaerial and freshwater facies criteria. Cave carbonates, karst: Formed by the precipitation of calcium carbonate within caves resulting in ‘speleothems’ (e.g. flowstones, dripstones). Eolian carbonates: Sedimentation of fine-grained, wind-born sand-sized carbonate material (–> eolianites). Glacial carbonates: Dissolution and recrystallization of glacially transported carbonate debris, occurs in glacial-marine deposits. Non-marine carbonates formed in terrestrial aquatic settings Freshwater carbonates: Travertine, calcareous tufa, calcareous sinter: Formation of carbonate deposits at warm, hot ore cold spring orifices in subaquatious settings (e.g. pools), by combined biotic and abiotic processes. Lacustrine carbonates: Deposition and precipitation of calcium carbonate in lakes of different salinity (freshwater lakes, salt lakes, playa lakes) and different settings. Fluvial carbonates: Carbonate deposits formed in rivers, creeks and waterfalls, originating from the combined activity of biotic and abiotic processes.
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Important studies of pedogenic carbonates deal with late Pleistocene and Holocene calcretes on the Bahamas, Florida Keys, Yucatan, Barbados, Persian Gulf, Western Australia, and the Mediterranean (see Basics and references with Code K030). Calcareous paleosols typically contain a higher proportion of mineral insoluble residue than the enclosing limestone units. Two models have been developed for explaining the mechanisms that lead to the concentration of insoluble residue: (1) Selective dissolution and removal of calcium carbonate from carbonate host rock, and (2) accretion of wind-driven dust at the limestone exposure surface without major dissolution of the limestones (Multer and Hoffmeister 1968). Most authors favor the first explanation. Calcretes are defined as ‘secondary accumulations of calcium carbonate (Low- and High-Mg calcite) in nearsurface settings, which result from the cementation and/or replacement of host material by the precipitation of calcium carbonate from soil water or ground water’ (Wright 1990). In addition to the calcium carbonate derived from the host rock, eolian input is still an important factor (Goudie 1973). Currently, two basically different models are being discussed with respect to the formation of calcretes. The commonly held model, that calcretes are pedogene structures resulting from leaching of carbonate from the surface horizon, is questioned by the second model, claiming that calcretes are complex sedimentary formations that have undergone phases of deposition, pedogenesis, and transformation of sediments during early diagenesis heavily related to microbially influenced processes. According to the first model, calcrete is formed in the interior of a soil; according to the second model calcretes are accumulations that build up by aggradation from bottom to top (Verrecchia 1996). Classification: Calcretes are classified according (a) to their micromorphology and hydrological setting or (b) dolomite content (Netterberg 1978). Using micromorphological criteria, massive ‘alpha calcretes’ (exhibiting a dense pseudonodular microfabric with complex microcracks, microsparite and coarse-crystalline calcite, floating sediment grains (Pl. 128/6), and calcretes exhibiting microfabrics of biogenic (particularly fungal) origin, called ‘beta calcretes’, can be distinguished (Wright 1990). The latter group is characterized by microbial coatings, calcified tubules, Microcodium, alveolar septal fabrics, and needle fibre calcite (Pl. 128). Two calcrete types are differentiated according to their hydrologic setting: (1) Pedogenic calcretes derived from vertically percolating soil water and typically occurring in the form
Caliche
of well-horizonted soil profiles. Characteristic features are distinct profiles made up of a succession of different horizons and specific microfacies attesting to the presence of a vegetated soil cover. ‘Rhizogenic calcretes’ produced by calcification associated with plant roots occur in the fluvial sequences of floodplains (Wright et al. 1995), palustrine and lake margin areas with relatively high water-tables, on subaerially exposed marine carbonates (Wright 1994), in carbonate dunes (Hay and Reeder 1978) or in beachrock deposits. (2) Non-pedogenic calcretes, derived from laterally migrating groundwater and composed of well-indurated, variably extensive sheets of mottled carbonate (Goudie 1983). In contrast to the rather thin pedogenic calcretes (a few meters), these calcretes may have a thickness of more than 10 meters as shown by Australian examples. Compared with pedogenic calcretes, non-pedogenic calcretes are characterized by the absence of rhizoconcretions, black pebbles and peloids, and by more silica than carbonate. Critical factors in the formation of caliche are (1) climate (a significant net annual moisture deficit in the soil; control on dust accretion caused by dust storms), (2) topography, (3) vegetation, (4) presence of carbonate and oxalic phases, (5) carbonate content, texture, porosity and permeability of the hosting substrate (Thériault and Desrochers 1993), (6) action of microorganisms (Loisy et al. 1999) as well as (7) exposure time (which may range between a few tens of thousands, as shown by Holocene calcretes, and millions of years as suggested by Carboniferous calcretes). Vegetation is the main factor in rhizogenic calcretes (Klappa 1979; Wright et al. 1995) exhibiting textures which can be interpreted as calcification in, on or around roots: Intracellular calcification creates Microcodiumtype microstructures. Extracellular calcification around roots produces laminar rhizolites (Wright et al. 1988) consisting of vertical rhizoconcretions or as calcareous cementation around root mats (Pl. 20/6). Incomplete calcification of root mats may be responsible for the formation of rhizogenic calcretes characterized by peloidal and coated grain fabrics (Calvet and Julia 1983). Exposure time, substrate composition and vegetation control the maturity levels of calcrete profiles. Relatively mature levels can be recognized from the sequence of horizons forming the calcrete profiles and by the number of calcrete microfabrics (which increase with increasing maturity). A number of models for pedogenic calcrete profiles has been developed, focused on the progressive increase in the volume of secondary calcium carbonate within
Cave Carbonates
the host material (soil, sediment or rock) via a combination of cementation, replacement and displacement (Gile et al. 1966; Arakel 1982; Machette 1985; West et al. 1988; Rabenhorst et al. 1990). Diagnostic criteria of caliche and paleocaliche: see Chap. 13 and Pl. 14/6, Pl. 128, and Pl 141/9-10.
2.4.1.2 Palustrine Carbonates Palustrine limestones are deposits of shallow freshwater environments exhibiting extensive evidence of pedogenic modification (Pl. 48). These carbonates are common in many ancient continental and marginalmarine sequences. The term (derived from ‘paludal’, meaning marshy or swampy) was first proposed while describing late Cretaceous-Early Tertiary limestones in southern France (Freytet 1973, 1984; Freytet and Plaziat 1982). Palustrine carbonates were originally considered as near-shore deposits of extremely shallow lakes with oscillating lake levels and densely vegetated shorelines. Because of a very low relief on the lake bottom, fluctuation in the water table should result in an interbedding of lacustrine facies (characterized by the occurrence of specific fossils), and pedogenic facies (characterized by textures which indicate a pedogenic overprint of dried-out parts of shores). This interpretation presents problems since many palustrine limestones are not associated with lake deposits. A suitable modern analogue for a palustrine environment may be the freshwater carbonate marshes of the interior Everglades in Florida (seasonal wetland model: Platt and Wright 1992). Seasonal variations in water depth and minor topographic variations across the area could explain the association of soil and calcrete criteria with dissolution features (pseudomicrokarst), and the presence of sublittoral carbonates which were formed in perennial shallow lakes within the wetland.
2.4.1.3 Cave Carbonates, Speleothems and Karst Speleothems (carbonates formed in caves) are important paleoclimatic recorders documenting short- and long-term climatic fluctuations. Terminology: The term karst refers to physical structures formed by dissolution from meteoric waters. It includes ‘all of the diagenetic features – macroscopic and microscopic, surface and subterranean – that are produced during the chemical dissolution and associated modification of a carbonate sequence’ (Choquette
13
and James 1988). Karst develops (a) at the air-limestone or soil-limestone interface (‘surface karst’) and (b) within carbonate bodies (‘subsurface karst’). Diagnostic features: Surface karst or exokarst is characterized by surface solution or corrosion features comprising small-scale and large-scale criteria (microkarst, residual paleosols, karren, solution pans or kamenitzas, and dolines or sinkholes). Large-scale geomorphological features (e.g. karst valley and karst towers) are responsible for characteristic landforms. Subsurface karst (see Pl. 129) is represented by caves and vugs formed above, at and below the water table. Caves contain (a) secondary mineral deposits (porous calcareous tufa: moonmilk, and hard crystalline sinter or speleothems) precipitated by vadose and shallow phreatic waters and (b) internal cave sediments and karst breccia caused by the collapse of caves. Karst breccias exhibit a wide variety of internal composition depending on the site and the mode of origin (Sect. 5.3.3.3 and Sect. 15.2). Speleothems are precipitated from thin water films flowing over the rock and forming flowstones and dripstones, or in pools of water, forming ‘cave pearls’ (globoids, pisoids; vadoids: Chafetz and Butler 1980; Peryt 1983. See Pl. 14/8). Dripstones include stalagmites and stalactites as well as finely curved twiglike structures (helictites). Carbonate sinters are made up of calcite with various amounts of Mg, and/or aragonite, dolomite and rarer minerals (Gonzalez and Lohmann 1988; Tietz 1988). The mineralogical composition is often seasonally controlled; it depends on the Mg/Ca ratio of the percolating pore-fluids, pCO2, and evaporative conditions within the cave. Speleothems exhibit diagnostic cave cements (Folk and Assereto 1976; Kendall and Broughton 1978; Assereto and Folk 1980) which can be used in order to recognize ancient cave carbonates (Sect. 15.2.1). Common cement types in speleothems are length-slow and length-fast calcites and dendritic calcite crystals. Changes in mineralogical composition and cement types of vadose-zone speleothems may reflect short and long-ranging cyclic oscillations of climatic factors (Railsback et al. 1994; Shopov et al. 1994). There is increasing evidence for microbial contributions to the growth and fabric of speleothems (Danielli and Edington 1983; Jones and MacDonald 1989). Controls: Prerequisites for the formation of extant karst and cave systems are fluids that are undersaturated with respect to the country rock, and fluid flow transporting dissolution products (Lohmann 1988; Dreybrodt 1990). The most important intrinsic factors influencing karst development are the general lithol-
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ogy (including mineralogy, bulk purity, fabric and texture), stratal permeability, bedding planes, fractures as well as the existence of confined or unconfined aquifers and structural conduits. Limestones are several orders of magnitude more soluble than dolomites in meteoric waters. Major extrinsic factors are climate (rainfall and evaporation, temperature), base level (elevation and relief, sea level or local water bodies), vegetation and the duration of time involved in karstification. Different climatic conditions produce different karst types (James and Choquette 1988). Karst formed in areas of warm temperatures and high rainfall is characterized by the development of soils and terra rossa, abundant sinkholes, and dissolution-collapse breccia. In semiarid and temperate climates both karst and calcrete may occur together, depending on alternating wet and dry conditions (Esteban and Klappa 1983). The potential of microfacies studies for the recognition of ancient speleothems and karst and the significance of paleokarst in basin analyses is discussed in Sect. 15.2.
2.4.1.4 Eolian Carbonates Terminology: The term eolianite has been coined for bioclastic wind-borne deposits that make up more than 90% of the Pleistocene deposits of Bermuda (Sayles 1931). The term is now used for eolian sands that have been cemented by calcium carbonate in a subaerial environment. The coalescence of dunes results in linear ridges with large cross-bed sets which may stand a few tens of meters above the source beach. Glacial, interglacial and Holocene eolian carbonate dune deposits are known from Bermuda, the Bahamas, northeast Yucatan, California, the Trucial Coast, the Mediterranean (Mallorca, Tunisia, Israel) and western Australia. Generally modern eolianites are associated with coastal marine deposits and pedogenic carbonates. Wind can transport carbonate particles (ooids, microfossils) over wide distances (e.g. from Africa to the Caribbean). Recognizing carbonate eolianites: Summaries of features common among carbonate eolianites were given by McKee and Ward (1983) and Abegg et al. (2001). Important criteria are: (1) Grain texture - good overall sorting, fine- to medium-grained, fenestrae, sorting differences between laminae or beds. (2) Carbonate cements - vadose and phreatic Low-Mg calcite cements (Pl. 2/2 and Pl. 31/5, 6, meniscus, pendant, needle-fiber, blocky, drusy). (3) Biota - plant structures (rhizoconcretions, calcified root systems of dune plants),
Travertine
Microcodium, ichnofossils: crustacean and insect burrows, bioturbation; terrestrial gastropods. (4) Stratification - high- to low-angle beds, see Fig. 15.1. (5) Associated facies - upward-shallowing succession, shelf/ reef - beach grainstone - eolian grainstone. Grains are ooids or bioclasts, sometimes accumulations of particular foraminiferal tests (forming the ‘miliolite’ of the Trucial Coast, a grainstone composed predominantly of miliolid foraminifera). Significance: The differentiation of wind-borne calcarenites from grainstones formed in marine environments is important because calcareous eolianites indicate carbonate shorelines, sea-level fluctuations (aragonite cements in eolianites indicate a later marine transgression; White 1995, San Salvador Island, Bahamas), paleoclimatic conditions and paleowind directions. 2.4.1.5 Glacial Carbonates The dissolution and recrystallization of glacially transported carbonate debris may result in the formation of micritic and sparitic carbonates, leading to a chemically distinctive carbonate matrix in meltout tills or subglacial crusts within glacial-marine deposits (Hallet 1976; Aharon 1988; Fairchild and Spiro 1990; Fairchild et al. 1993), which are also recorded in Precambrian and Phanerozoic glacial deposits. Enigmatic calcite pseudomorphs, known as glendonites and interpreted as pseudomorphs after ikaite, a metastable hexahydrate of calcium carbonate, were reported from some modern as well as ancient glaciomarine environments. Ikaite occurs in morgenstern-like and stellate aggregates, several centimeters in size, and indicates sub-zero temperatures at the seafloor. Glendonites are typically found in cold-water deposits dating from the Precambrian to the Pleistocene (Suess et al. 1982; Shearman and Smith 1985; Kemper 1987). Non-marine carbonates formed in aquatic settings comprise freshwater carbonates, lacustrine carbonates and fluvial carbonates.
2.4.1.6 Travertine, Calcareous Tufa and Calcareous Sinter Origin: Travertine and tufa are freshwater carbonates formed as a result of physical and/or biochemical CO2 degassing around carbonate- and CO2-rich springs, along streams and in pools, and often precipitated together with and on cyanobacteria, bacteria, algae (Pl.
Travertine
2/6), mosses, and higher plants (Pl. 2/7) as well as on various organic and inorganic substrate. Cyanobacteria, bacteria and algae are able to trigger the precipitation of calcite and aragonite forming microbial travertines (Chafetz and Folk 1984; Folk et al. 1985). Millimeter- to centimeter-sized laminae (Pl. 2/1, 2), and arborescent shrub-like calcite precipitates (Pl. 2/1), consisting of associations of micrite aggregates and rhombic spar crystals aggregates appear to be microbially controlled (Guo and Riding 1994). Stromatolitic travertines exhibit alternating sparry and micritic laminae which are interpreted in terms of seasonality (Chafetz et al. 1991), climatic changes as well as changes in microbial and algal growth and sedimentation (Monty 1976). Travertines and tufa are often associated with higher plants (Pentecost 1990), but the ability of macrophytes to induce the precipitation of carbonate during photosynthesis is regarded as controversial. Consumption of CO2 leads to carbonate precipitation in the immediate vicinity of photosynthesizing vegetation. Both field and laboratory studies show that the intensity of photosynthesis varies from least effective in springs and waterfalls to highly effective in lakes and rivers (Merz 1992). Morphological differentiation: Travertine deposits exhibit distinctive morphologies (Chafetz and Folk 1984): •
•
•
• •
Waterfall deposits, formed at sites of break-in slope of streams by the growth of plants and contemporaneous calcification, resulting in a massive, irregular contorted tangle of cement crusts and plant molds. Lake-fill accumulations, characterized by thick, horizontally stratified and laterally extensive travertines. These are composed of thin summer layers of (bacterial) shrubs intercalated with finely laminated winter mud layers and thicker, crudely laminated muds. Contorted zones composed of calcite rays, intraclasts and pisoids, represent spring orifices on the lake bottom. Sloping mounds, fans and cones, characterized by structures paralleling a slope at the time of deposition. Terraced mounds, exhibiting a step-like morphology. Warm spring waters, flowing down both sides of fissure ridges.
Which names are appropriate? The terminology of freshwater-carbonates is controversial because of the mixed use of descriptive and genetic terms. This is particularly true for carbonates called travertine (Pl. 2/1,
15
Fig. 2.4. Practical classification of travertines, tufa and sinter (after Koban and Schweigert 1993).
2, 3, 5), tufa (Pl. 2/6, 7) or sinter (Pl. 2/4) (Schweigert 1996). Travertine often is used in a rather loose manner to designate hard and compact carbonate deposits; tufa for porous deposits. Many authors use travertine in a rather broad sense, covering all kinds of non-marine carbonates of springs, creeks, streams, pools and lakes, as well as caves (e.g. Julia 1983; Pentecost and Whitton 2000). More precise definitions were offered by Riding (1991): The term travertine is restricted to non-marine, layered autochthonous carbonates deposited at thermal springs and commonly exhibiting a bushy fabric (Pl. 2/ 1). Tufa refers to often conspicuously porous carbonates, whose formation in non-marine cold-waters is strongly controlled by aquatic plants. Calcareous sinters are characterized by well-developed lamination and lack of visible porosity. Currently there is a tendency to restrict the term travertine to thermal deposits (Ford and Pedley 1996; Fouke et al. 2000) formed at hot and warm, often mineralized springs, and calcareous tufa to settings where calcium-rich waters bubble up from springs to the earth, cool down and precipitate calcium carbonate, supported by photosynthesizing plants. How to classify freshwater carbonates? Koban and Schweigert (1993) proposed a classification which reflects the genetic factors responsible for the fabric of non-marine freshwater carbonates (Fig. 2.4): The triangle underlines the inferred origin, setting and visible porosity (see Pl. 2 for examples). Thermal travertines: Comparative studies of modern thermal spring systems (Cady and Farmer 1996; Fouke et al. 2000) provide a basis for constructing facies frameworks which can assist in the interpretation of ancient travertines. The spectacular carbonate-precipitating thermal springs of the Mammoth Hot Springs in Yellowstone National Park, Wyoming, demonstrate
16
Travertine,Tufa, Sinter
Plate 2 Terrestrial Freshwater Carbonates: Travertine, Calcareous Tufa and Sinter Travertines, tufa and calcareous sinter are excellent paleoclimatic archives (Andrews et al. 2000). These carbonates and their corresponding water bodies can serve as analogues of Precambrian and Phanerozoic microbialites and their oceans (Arp et al. 2001). Travertines of hot springs are of interest for recognizing potential microbial signatures on Mars (Kempe and Kazmierczak 1997). Travertines are very important building stones that have been used since antique times (Winkler 1994). Many travertines and tufa originated during interglacial phases of the Pleistocene, but in places are still growing today. Terrestrial carbonates produced by precipitation in freshwater occur as travertines (–>1-3, 5), calcareous tufa (–> 6, 7), and calcareous sinter (–> 4). Travertine is a layered deposit of calcium carbonate with moderate to high primary porosity, and often with a dendritic fabric. Calcareous tufa is an originally porous, laminated or massive deposit which commonly originates at cold subaerial springs, but also forms at subaquatic springs in soda lakes and highly alkaline salt lakes (‘tufa pinnacles’: Kempe et al. 1991; Benson 1994; Arp et al. 1998). Calcareous sinter is characterized by well-developed lamination and the lack of visible porosity. These carbonates orignate from degassing of CO2 and HCO3- caused by changes in temperature, pressure and photosynthesis. Travertine is often found on warm or mineralized hydrothermal waters, tufa on temperate waters. Tufa precipitation takes place on organic substrates (higher plants, algae, mosses) in freshwater. In both types carbonate chemistry and the temperature of the spring waters, the overall morphology of the accumulation and biologically induced processes are major controls. Characteristic textural criteria of freshwater limestones are (1) macroscopically visible voids (often corresponding to moulds and leaves of higher plants); (2) void-filling radiating calcite (–> 3); (3) small spar-filled cellular structures; (4) wavy and laminar layers (–> 1, 2); (3) concentric structures consisting of smooth and wavy laminae of micrite and sparite (–> 1); (5) alternating dark and light laminae with vertical thin ‘filaments’ (–> 4); (6) layers with bubble structures (–> 1); (7) shrub-like structures (–> 1); (8) ovoid to subangular peloids with diameters from 50 to 300 µm, occurring in clumps or aggregates (–> 5); (9) travertine fragments occurring as intraclasts, and detrital grains derived from underlying or adjacent rocks; (10) pisoids (see Pl. 14/1). 1 Travertine. The name travertine is derived from the Latin lapis tiburtinus, meaning ‘the stone of Tiburnium’ (now Tivoli), a locality near Rome where these carbonate rocks have been quarried for more than two thousand years. At the ‘type locality’, the precipitation of carbonate occurs from spring waters. The importance of inorganic or organic (bacteriallyinduced) modes of precipitation is primarily controlled by water temperature and water chemistry (Chafetz and Folk 1984). In the Bagni di Tivoli locality, travertine occurs as stratified lake-fill deposits formed in shallow ponds.The travertine is composed of laterally continuous and vertical repetitions of layers with micritic shrubs (S; interpreted as result of bacterially induced precipitation) intercalated with laminated muds (–> 3). The shrubs nucleate on a substrate and grow and branch upward, indicating a phototrophic control. The bubbles (B) at the bottom may represent lithified gas-filled vesicles surrounded by bacterially precipitated walls or relics of chironomid insect larvae. The samples (–> 1 and 2) are classified as travertine according to the vadose setting, microbial-induced precipitation, and medium porosity (Fig. 2.4). 2 Travertine. ‘Bacterial stromatolite’. Thin laminae aggregate into packets with large lenticular intermat cavities bulged up by large gas bubbles (B). Note the abundant micritic clumps within the mats (arrows) interpreted as bacterial aggregates embedded or surrounded by calcite crystals. Bagni di Tivoli near Rome. 3 Travertine. Laminated micrite layers bordering voids which are partly filled with bladed prismatic (‘dogtooth’, DC) calcite cement. Same locality as –> 2. 4 Calcareous sinter. The layers consist of fibrous cements arranged in bundles. The low-calcite sinter originated from carbonic acid-rich and calcium-rich mineral waters. The travertine was formed during warm interglacial periods in the Pleistocene. The sample is classified as calcareous sinter owing to inferred inorganic precipitation, low porosity and the vadose setting. Stuttgart-Bad Cannstatt, Germany. 5 Travertine. Peloids and bacterial mud reworked into intraclasts. Same locality as –> 2. 6 Recent freshwater tufa. The fabric is characterized by incrustations of oscillatoriacean cyanobacteria. The threads have a diameter of about 10 µm. Fossil Oscillatoriaceae are indicative of freshwater and brackish water environments. Schleierfälle, Ammer, southern Germany. 7 Modern freshwater tufa. Tufa deposition is controlled by a number of intervening factors including carbonate equilibrium, crystal nucleation, crystal morphology, and diagenesis (Ford and Pedley 1996; Merz-Preiß and Riding 1999: Jannsen et al. 1999). Note the porous texture and the plant stem (P) surrounded by euhedral calcite crystals some of which exhibit dark ‘centers’ (arrow) commonly interpreted as an indication of bacterial contributions to carbonate precipitation. The sample still fits into the triangle classification (Fig. 2.4) because of its high porosity, considerable organic contributions (–> 6) and origin in a phreatic setting. Same locality as –> 6.
Plate 2: Travertine, Tufa, Sinter
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Lakes
18
the distinctive control of travertine microfacies types and dominating organisms (eucaryonts, photosynthetic and chemosynthetic bacteria) by water temperatures and pH values (Farmer 2000). Pamukale near Denizli (southeastern Turkey) is another famous locality where thermal spring waters at a hill built, and still build, large rimstone pools of travertine while flowing down in cascades about 100 m of height difference. Because of the large surface of the very thin layer of water, the water looses deep-source carbon dioxide, resulting in carbonate precipitation. The resulting white travertines have been used since Roman times as building stones. The Turkish name Pammukale means ‘cotton castle’ and refers to the strikingly white color of the large steps. Phytoherms: Plant-dominated freshwater calcareous tufa, formed in river systems, are called ‘phytoherms’ (Cipriani et al. 1977; Pedley 1987, 1992) underlining the dominant role of macrophytes and procaryotemicrophyte biofilms in the construction of carbonate buildups (Collinson 1988). Phytoherms of river systems, capable of damming the water courses to produce chains of lakes (e.g. Plitvica region, Croatia: Emeis et al. 1987), are characterized by rapid growth in cool, agitated waters. Continuous water circulation is necessary for bringing in nutrients and providing replenishment of CaCO3 for the extensive precipitation of LowMg-calcite cements within organic frameworks. The primary frameworks are provided by the in-situ growth of erect and trailing hydrophytes and semi-aquatic macrophytes. Secondary frameworks are formed by oscillatoriacean cyanobacteria, diatoms and insect larvae colonizing exposed surfaces. The shelly fauna associated with the phytoherms contains ostracods, gastropods, rare bivalves and crustaceans. The stability of these bioconstructions and also their preservation in the fossil record is provided by a concomitant development of thick Low-Mg calcite calcite cements (micrite fringes and bladed calcite spar fringes) and by an intra-frame sediment infill. Both micritic fringe cements and the abundant small-scaled peloids are interpreted as the product of biological mediation associated with a procaryote-microphyte biofilm and the presence of organic detritus. Significance: Travertines are important building and decoration stones, used since antique times. High quality travertine possesses a great mechanical strength and remarkable resistance to weathering as compared with other building stones. Technological qualities of travertines are strongly controlled by depositional facies and early diagenesis. Differences in technological properties of travertines and tufas are primarily caused by
fabric-selective porosity patterns. Changes in porosity, permeability, specific surface, and water sorption due to the dissolution and precipitation of calcite within pores lead to a decrease in porosity and increase in rock hardness with time (Koch et al. 1999). The rapid precipitation of carbonate favors the preservation of fossil faunas and floras, and of archeological artifacts. Chronological dating of Pleistocene travertines aids greatly in the age determination of prehistoric archeological material (Kahlke 1984). 2.4.1.7 Lacustrine Carbonates: Lakes Lacustrine environments differ from marine environments in the properties, kinetics, chemistry and productivity of lake waters as well as in the sedimentation patterns originating in hydrologically open and hydrologically closed lakes. Lakes are dynamic systems. Short-term and long-term changes in climate, water chemistry and bathymetry of lakes result in significant variations of litho- and biofacies, and their lateral and vertical distribution patterns. Our review focuses on the carbonates formed in lakes. Carbonates are only a part of lacustrine sediments, which in addition, often include claystones, marls and also sandstones. The review neglects the siliciclastic sediments as well as some chemical deposits (cherts, phosphorites). Lacustrine carbonates are a specific type of terrestrial carbonate deposited within a lake or at the shorelines of a lake. Fluctuating shorelines may result in a pedogenic overprint of lacustrine sediments leading to the formation of ‘palustrine carbonates’ (Sect. 2.4.1.2), which were defined as ‘near-shore deposits of extremely shallow lakes with oscillating lake levels and densely vegetated shorelines’ (Freytet and Plaziat 1982). Frequently, lake deposits are associated with paleosols as well as with fluvial sediments. Because wind-driven currents and waves often do not affect the lake bottom, lake waters tend to become permanently or seasonally stratified: Warmer and lighter waters (epilimnion) overlie colder and denser waters of the bottom zone (hypolimnion). These zones are separated by a zone characterized by a rapid rate of change of temperature with depth (mesolimnion). Permanent stratification is characteristic of tropical oligomictic lakes with temperature stratification and of meromictic lakes in which the bottom water has a higher salinity than the surface water. In both cases, the bottom water can become completely stagnant and be depleted of nutrients and oxygen, resulting in the accumulation of laminated muds rich in organic matter.
Lakes
Seasonal stratification is common in temperate climates. During the summer, surface waters are warmer and less dense than deeper waters. Cooling of the surface waters during the winter or an underflow by cool waters caused by rivers crossing the lake will produce surface waters of the same or higher density than that of the bottom waters. Consequently, the lake water body is turned over once or twice yearly and mixed (monomictic or dimictic lakes; tropical and polar lakes). Both near-surface waters and bottom waters are well-oxygenated, enabling a seasonal biogenic production (e.g. algae), which may be recorded by annual varves consisting of fine-grained siliciclastics, biogenic silica and carbonate formed in the deeper parts of the lakes. Fine-scale lamination, caused by the seasonal precipitation of carbonates or periodic influxes of sediment into lakes, is a common feature of both freshwater and saltwater lakes. Biologically induced precipitation of light carbonate laminae alternating with dark organic-rich laminae and/or detrital laminae form varved sediments in the deeper parts of many lakes. Open and closed lakes: Freshwater and salt lakes Lakes are classified according to their salinity. Freshwater lakes usually contain <1 g/l dissolved constituents, brackish water lakes 1–5 g/l, and salt lakes >5 g/l. The setting of lakes in hydrologically open or hydrologically closed basins as well as the chemistry of lake waters is used in separating major types of lacustrine environments (Allen and Collinson 1986): (1) Hydrologically open lakes (freshwater lakes) have an outlet, relatively stable shorelines and are filled with meteoric lake waters. They are characterized by the input of siliciclastic sediments in nearshore and offshore zones as well as by inorganic/biochemical and biogenic deposition of calcareous sediments in the shallower littoral zone and the deeper profundal zone. Most freshwater lakes are fed by rainfall; the salinity, therefore, is low. Sedimentation patterns vary in response to changes in depositional regimes and source areas (Calvo et al. 1989). Hydrologically open freshwater lakes are normally perennial and occur mainly in temperate humid climates. The subdivision of lake environments is based on biological criteria and embraces the following zones (Wright 1990): • Littoral zone: Below the wave base, characterized by rooted macrophytes. The eulittoral zone is the area between the highest and lowest sea levels, the supralittoral zone corresponds to the landward zone, which is rarely submerged. • Sublittoral zone: Still within the photic zone, colonized by algae and cyanobacteria.
19
• Profundal zone: Aphotic zone. Often, but not always corresponds to the zone below the thermocline. • Pelagial zone: Open-water, planktic zone. Carbonates in freshwater lakes result from: • Removal of CO2 by photosynthesis of plants (microbes, algae, floating phytoplankton, macrophytes), less commonly by changes in temperature, evaporation or mixing of water masses, resulting in the precipitation of Low-Mg calcite. Basin-wide carbonate blankets are produced in larger and deeper stratified lakes resulting from a seasonal change of algal blooms and sedimentation of organic matter, diatoms and detrital components, producing couplets of light carbonate-rich laminae and dark organic-rich laminae. • Deposition of ‘algal sediments’, algally-coated grains (e.g. oncoids; Pl. 12/5, Pl. 131/5) and of characean oogonia (Pl. 65) in nearshore shallow parts of the littoral zone and accumulation of carbonate shells (e.g. gastropods, bivalves, or ostracods; Pl. 130/2, 3) in deeper parts of the littoral zone, resulting in biogenic sedimentation. Differences in the composition of the algal carbonates may reflect the generally well-defined zonal distribution of vegetation. • Growth of stromatolitic bioherms (Pl. 131/1) or ‘algal reefs’ (Pl. 130/1). • Input of allochthonous carbonate particles derived from the drainage basin and deposited by debris flows and turbidity currents in deeper part of the lake. • Early diagenetic precipitation of carbonates (nonmarine carbonate cements; Chafetz et al. 1985). Case studies: Europe: Lake Constance, Germany: (Müller 1971; Schäfer and Stapf 1978; Dominik 1981); Eifel, Maar Lakes, Germany (Negendank et al. 1990); Lake Zürich and Lake Biel, Switzerland (Kelts and Hsü 1978; Wright et al. 1980); Austria: Behbehani et al. (1986); North America: Green Lake, New York State (Eggleston and Dean 1976); Minnesota Lakes (Megard 1967; Murphy and Wilkinson 1980). The Green Lake exhibits wave-resistant stromatolite bioherms, formed by cyanobacteria and mosses trapping carbonate sediment, and subsequent cementation by precipitated calcium carbonate. Lake Michigan is an impressive example of facies differentiation in a typical northern temperate hard-water lake. (2) Hydrologically closed lakes (salt lakes) lack an outlet. They are common in regions with semi-arid warm and hot climates. If evaporation of the lake water is about equal to inflow, a permanent water body is
20
maintained (perennial lake). If the evaporation exceeds inflow for a long time, salinity and the Mg/Ca ratio of the water will increase, as Ca is depleted by early-stage precipitates. Finally salts will be precipitated in these shallow ‘salt lakes’ and ‘soda lakes’ which normally fall dry, with the exception of a central pond containing highly concentrated brines (ephemeral lakes, playa lakes: Last 1993). High and extreme salinities can originate from a separation of former parts of the sea (e.g. separation of the Caspian Sea from the Mediterranean Sea), prolonged evaporation of lake waters, or increased input of dissolved salts. Case studies: Mono Lake, Death Valley, California; Great Salt Lake, Utah (Eardly 1966; Halley 1976); seamarginal ponds of the eastern Mediterranean e.g. Solar Lake, Israel (Krumbein and Cohen 1974; Friedman 1978); sea-marginal flats in the Red Sea, Gulf of Elat and Aqaba (Sneh and Friedman 1985); Australia: (De Deckker 1988); Lake Clifton (Burne and Moore 1983); Lake Thetis (Grey et al. 1990); Coorong Lagoon (Rosen et al. 1988). One of the most intersting saline lakes is the Dead Sea whose bottom sediments consists of precipitated aragonite, gypsum and halite, and detrital grains (Garber et al. 1987). Many hydrologically closed lakes are characterized by fringes of supralittoral mudflats (commonly termed ‘playa’ or ‘inland sabkha’), sandflats in the nearshore zone, and depositions of mixed siliciclastic-evaporitic sediments as well as organic-rich sediments (‘oil shales’) within the lakes. Saline minerals are precipitated in perennial brines, as efflorescent crusts and salt pans and as cements within the sediments of the mudflats. Climatic changes may turn freshwater lakes into saline water bodies, as exemplified by the history of the modern Great Salt Lake, Utah, or the great lakes in the East African rift system. Carbonates in saltwater lakes are characterized by: • Precipitation of High-Mg calcite and aragonite (e.g. Dead Sea) or Low-Mg calcite and High-Mg calcite (e.g. Lake Balaton, Hungary) because of the increase in the Mg/Ca ratio, • Formation of carbonate laminites, consisting of CaMg carbonates, detrital quartz, silicates and organic layers, • Ooid sands are formed in shallow, near-shore areas of current- or wave-swept lakes, sometimes (e.g. Great Salt Lake, Utah) associated with • Algal bioherms, composed of mm-scale laminated micritic aragonite or of porous carbonate, • Loss of CO2 by degassing or photosynthesis of spring
Lakes
waters on the playa surfaces cause supersaturation with respect to calcite, resulting in the formation of pisoids and carbonate encrustations (Risacher and Eugster 1979), • Hydrothermal fluids emerging from the lake bottom may produce sinter crusts consisting of aragonite and magnesium calcite (Lake Tanganyika, East-Central Africa: Cohen and Thouin 1987, Stoffers and Botz 1994).
Compositional characteristics of lake carbonates Similar to marine carbonates, lacustrine carbonates are composed of allochthonous and autochthonous constituents, the former represented by skeletal and nonskeletal grains and fine-grained calcite mud caused by seasonal phytoplankton blooms (‘Seekreide’: Schäfer 1973), the latter by predominantly biogenic structures (stromatolites, bioherms and ‘reefs’) formed or induced by cyanobacteria and algae. All major grain categories known from marine carbonates also occur in lacustrine settings. Biota: Differences exist in the inventory of the biota. Common organisms are (a) calcareous algae (charophycean algae, see Sect. 10.2.1.8 and Pl. 65; planktonic and benthic green algae, see Pl.130/1) and cyanobacteria, siliceous algae (diatoms), (b) gastropods and bivalves, cf. Pl. 130/2, 3, (c) ostracods (see Pl. 130/2), and (d) arthropods that are responsible for the production of abundant peloids. Otherwise marine organisms, e.g. foraminifera, also may occur in salt lakes. These marine biota are transported by wind and may thrive in lakes if environmental conditions are favorable (Lee and Anderson 1991). Bioturbation and bioerosion are very important processes modifying the depositional texture of lacustrine sediments (Schneider 1977, Schröder et al. 1983). Peloids and reworked and redeposited sedimentary fragments (intraclasts, lithoclasts) are common constituents. Small ooids and large oncoids (Pl. 131/5) are of specific interest in microfacies studies of ancient lacustrine limestones because these grains may aid in the differentiation of non-marine and marine settings. Ooids occur in marine as well in non-marine settings (Pl. 13/1), they are known from low- and highsaline environments. Ooids are not common in recent freshwater lakes. Ooids originating in salt lakes have been described from the Great Salt Lake, Utah (Eardley 1966; Kahle 1974; Sandberg 1975; Halley 1977; Pl. 13/2) and the Pyramid Lake, Nevada (Popp and Wilkinson 1983). Ooids formed in a temperate ‘marl’ lake were
Fluvial Carbonates
reported from Michigan (Wilkinson et al. 1980). Saltwater and freshwater ooids exhibit differences in the morphology, microfabrics and mineralogy of the cortical layers, and in size ranges. The same is true for oolitic carbonate particles formed in situ in hypersaline algal mats (Friedman et al. 1973). These differences offer the possibility of tracing the development of ancient freshwater- and saltwater lakes (Sect. 4.2.5). Lacustrine oncoids have been described by various names (algal nodules, water biscuits, lake balls, lacustrine pisoliths) for many years (Pia 1926, 1933, Schäfer and Stapf 1978, Jones and Wilkinson 1978; Pl. 12/5). The origin of these grains is related to the activity of algae and microbes (Sect. 4.2.4.1). Cyanobacteria are known or assumed to be essential in the formation of these oncoids. Modern oncoids are not restricted to freshwater environments but also occur in brackish and hypersaline waters. Today, oncoids are rare in tidal and shallow subtidal marine environments. Recent marine oncoids are non-lithified, as opposed to modern hard freshwater oncoids that are closely similar to Paleozoic and early Mesozoic marine oncoids. This may be related to the replacement of porostromate oncoids by spongiostromate oncoids in marine settings during the Jurassic. Stromatolites occur both in freshwater and salt lakes. The stromatolites are usually domes, passively paving the relief of the substrate and reacting to environmental controls like currents. Stromatolites of less saline lakes are composed of calcium carbonate whose precipitation is induced by algal/bacterial activity. Trapping of clastic particles, widely known from modern stromatolites, appears to be only of secondary importance. Stromatolites of saline lakes are well-known from hypersaline lakes in Utah, Nevada and California as well as from the San Salvador Island, Bahamas (Mann and Nelson 1989). The stromatolites have complicated microstructures consisting of micrite, pellets, filament molds and casts as well as sparry calcite. Lacustrine stromatolites contribute to the growth of ‘algal mounds’ as demonstrated by the ‘reefs’ described from the Great Salt Lake in Utah, which exhibit a nonskeletal framework of algally induced aragonite precipitates, internal sediments and carbonate cements, and cover about 100 square kilometers of the nearshore areas. The mounds extend from the shoreline down to a depth of 4 m; they are surrounded by rippled ooid sand. Coccoid bluegreen algae on the present-day living surface are different from the filamentous algae recognizable in the indurated stromatolites. The change in the algal flora may be due to an increase in salinity. A similar pattern is known from the famous intertidal stromato-
21
lites of Shark Bay in western Australia. Spectacular meter-sized stromatolites consisting of hydromagnesite were described from the highly alkaline Salda Golu Lake in Southern Turkey (Braithwaite and Zedef 1996). The composite stromatolites originate from microbially induced precipitation of hydromagnesite, which is possible because of the dissolution of magnesium and calcium during the passage of water through ultramafic rocks.
2.4.1.8 Fluvial Carbonates Carbonate deposits originate in fluvial waters in various settings including small creeks, large rivers or waterfalls where tufa is formed. Stromatolites, oncoids and ooids in present-day creeks and streams are usually formed by the complex activity of algae, mosses and microbes (Nickel 1985; Galli and Sarti 1989). Fluvial oncoids and ooids differ in morphological and compositional criteria as well as in transport patterns (Verrecchia et al. 1997) from marine grains and are, therefore, valuable environmental indicators. Basics: Non-marine carbonates Esteban, M., Klappa, C.F. (1983): Subaerial exposure environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (ed.): Carbonate depositional environments. – Mem. Am. Ass. Petrol. Geol., 33, 1-54 Etheridge, F.G., Flores, R.M. (eds., 1981): Recent and ancient nonmarine depositional environments: models for exploration. – Soc. Econ. Paleont. Min. Spec. Publ., 31, 349 pp. James, N.P., Choquette, P.W. (1990): Limestone - The meteoric diagenetic environment. – In: McIlreath, I.A., Morrow, D.W. (eds.): Diagenesis. – Geoscience Canada, Reprint Series, 4, 35-73 Pia, J. (1926): Pflanzen als Gesteinsbildner. – 355 pp., Berlin (Borntraeger) Pia, J. (1933): Die rezenten Kalksteine. – MineralogischPetrographische Mitteilungen, Neue Folge, Ergänzungsband, 420 pp., Leipzig Further reading: K028 Pedogenic carbonates, caliche, paleosols Esteban, M.C. (1975): Vadose pisolites and caliche. – Amer. Ass. Petrol. Geol. Bull., 60, 1048-1057 Klappa, C.F. (1983): A process-response model for the formation of pedogenic calcretes. – In: Wilson, R.C.C. (ed.): Residual deposits. – Geol. Soc. London Spec. Publ., 11, 211220 Reeves, C.C. (1976): Caliche. Origin, classification, morphology and uses. – 233 pp., Lubbock, Texas (Estacado Books) Reinhardt, J., Sigleo, W.R. (1988): Paleosols and weathering through geologic time: principles and application. – Geol. Soc. America Spec. Paper, 216, 181 pp. Retallak, G.J. (2001): Soils of the past. Second edition. – 416 pp., Oxford (Blackwell)
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Wright, V.P. (ed., 1986): Paleosols. Their recognition and interpretation. – 315 pp., Oxford (Blackwell) Wright, V.P., Tucker, M.E. (1991): Calcretes. – Reprint Series Int. Ass. Sed., 2, 352 p., Oxford (Blackwell) Further reading: K030 Cave carbonates, speleothems and karst Borsato, A., Frisia S., Jones B., Van der Borg, K. (2000): Calcite moonmilk: crystal morphology and environment of formation in caves in the Italian Alps. – J. Sed. Research, A70, 1171-1182 Chafetz, H.S., Butler, J.C. (1980): Petrology of recent caliche, pisolites, spherulites and speleothem deposits from central Texas. – Sedimentology, 27, 497-518 Drybrodt, W. (1988): Processes in karst systems - physics, chemistry, geology. – Springer Series in Physical Environments, 4, 288 pp., New York (Springer) Folk, R.L., Assereto, R. (1976): Comparative fabric of lengthslow and length-fast calcite and calcitized aragonite in a Holocene speleothem, Carlsbad Cavern, New Mexico. – J. Sed. Petrol., 46, 486-496 Frisia, S., Borsato, A., Fairchild, I.J., McDermott, F. (2000): Calcite fabrics, growth mechanisms, and environments of formations in speleothems from the Italian Alps and southwestern Ireland. – J. Sed. Research, A70, 1183-1196 Choquette, P.W., James, N.P. (eds., 1988): Paleokarst. – 416 pp., New York (Springer) Fritz, R.D., Wilson, J.L., Yurewicz, D.A. (eds., 1993): Paleokarst related hydrocarbon reservoirs. – Soc. Econ. Paleont. Min. Core Workshop, 275 pp. Klinchouk, A.B., Ford, D.C., Palmer, A.N., Dreybrodt, W. (eds., 2000): Speleogenesis: evolution of karst aquifers. – 527 pp., Huntsville (National Speleological Society) Mylroie, J.E., Carew, J.L., Vacher, H.L. (1995): Karst development in the Bahamas and Bermuda. – In: Curran, H.A., White, B. (eds.): Terrestrial and marine geology of Bahamas and Bermuda. – Geol. Soc. America Spec. Paper, 300, 251-267 Purdy, E.G. Waltham, D. (1999): Reservoir implication of modern karst topography. – Amer. Ass. Petrol. Geol. Bull., 83, 1774-1794 Thrailkill, J. (1976): Speleothems. – In: Walter, M.E. (ed.): Stromatolites. – 73-86, Amsterdam (Elsevier) Tietz, G.F. (1988): Zur Genese rezenter Karbonatbildungen in Dolomithöhlen Frankens. – Karst und Höhle, 1988, 7-79 Further reading: K031 Eolian carbonates Abegg, F.E.R., Harris, P.M., Loope, D.B. (eds., 2001): Modern and ancient carbonate eolianites: sedimentology, sequence stratigraphy and diagenesis. – SEPM Spec. Publ., 71, 214 pp. Bigarella, J.J. (1972): Eolian environments: their characteristics, recognition, and importance. – Soc. Econ. Paleont. Min. Spec. Publ., 16, 12-61 Brookfield, M.E. (1992): Eolian systems. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 143-156, Ottawa (Geol. Ass. Canada) Caputo, M.V. (1995): Sedimentary architecture of Pleistocene eolian calcarenites, San Salvador Island, Bahamas. – Geol. Soc. America Spec. Paper, 300, 63-76 McKee, E., Ward, C.W. (1983): Eolian environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbon-
Basics: Non-Marine Carbonates
ate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 131-170 Pye, K. (1994): Aeolian sediments: ancient and modern. – Spec. Publ. Int. Ass. Sedimentologists., 16, 192 pp. Sayles, R.W. (1931): Bermuda during the Ice Age. – American Academy of Arts and Sciences, 66, 381-468 Further reading: K 032 Glacial carbonates Edwards, M. (1986): Glacial environments. – In: Reading, H.G. (ed.): Sedimentary environment and facies. – 445470, Oxford (Blackwell) Eyles, N., Eyles, C.H. (1992): Glacial depositional systems. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 73-100, Ottawa (Geol. Ass. Canada) Fairchild, I.J., Bradby, L., Spiro, B. (1993): Carbonate diagenesis in ice. – Geology, 21, 901-904 Fairchild, I.J., Spiro, B. (1990): Carbonate minerals in glacial sediments: Geochemical clues to paleoenvironment. – In: Scource, J.D., Dowdeswell, J.A. (eds.): Glacimarine environments and processes. – Geol. Soc. London Spec. Publ., 53, 241-256 Suess, E., Balzer, W., Hesse, K-F., Müller, P.J., Ungerer, C.A., Wefer, G. (1982): Calcium carbonate hexahydrate from organic rich sediments of the Antarctic shelf: precursors of glendolites. – Science, 216, 1128-1131, Washington Shearman, D.J., Smith, A.J. (1985): Ikaite, the parent mineral of jarrowite-type pseudomorphs. – Proc. Geol. Ass. London, 96, 305-314, London Further reading: K210 Calcareous tufa and travertine Andrews, J.E., Pedley, M., Dennis, P.F. (2000): Paleoenvironmental records in Holocene Spanish tufas: a stable isotope approach in search of reliable climatic archives. – Sedimentology, 47, 961-971 Arp, G., Wedemeyer, N., Reitner, J. (2001): Fluvial tufa formation in a hard-water creek (Deinschwanger Bach, Franconian Alb, Germany). – Facies, 44, 1-22 Chafetz, H.S., Folk, R.L. (1984): Travertines: depositional morphology and the bacterially constructed constituents. – J. Sed. Petrol., 54, 289-316 Folk,, R.L., Chafetz, H.S. (1983): Pisoliths (pisoids) in Quaternary travertines of Tivoli, Italy. – In: Peryt, T.M. (ed.): Coated grains. – 471-487, Berlin (Springer) Ford, T.D., Pedley, H.M. (1996): A review of tufa and travertine deposits of the world. – Earth-Science Reviews, 41, 117-175 Irion, G., Müller, G. (1968): Mineralogy, petrology and chemical composition of calcareous tufa from the Schwäbische Alb, Germany. – In: Müller, G., Friedman, G.M. (eds.): Recent developments in carbonate sedimentology in central Europe. – 157-171, Berlin (Springer) Julia, R. (1983): Travertines. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists Mem., 33, 64-72 Koban, C.G. (1993): Faziesanalyse und Genese der quartären Sauerwasserkalke von Stuttgart, Baden-Württemberg. – Profil, 5, 47-118 Koban, C.G., Schweigert, G. (1993): Microbial origin of travertine fabrics. Two examples from southern Germany
Basics: Non-Marine Carbonates
(Pleistocene Stuttgart travertines and Miocene Riedöschingen travertine). – Facies, 29, 251-264 Pedley, H.M. (1990): Classification and environmental models of cool freshwater tufas. – Sed. Geol., 68, 143-154 Pedley, M. (1992): Freshwater (phytoherm) reefs: the role of biofilms and their bearing on marine reef cementation. – Sed. Geol., 79, 255-274 Pedley, M. (2000): Ambient temperature freshwater microbial tufas. – In: Riding, R.E., Awramik, S.M. (eds.): Microbial sediments. – 177-186, Berlin (Springer) Pentecost, A., Whitton, B.A. (2000): Limestones. – In: Whitton, B.A., Potts, M. (eds.): The ecology of cyanobacteria. – 257-279, Dordrecht (Kluver) Renault, R.W., Jones, B. (2000): Microbial precipitates around continental hot springs and geysers. – In: Riding, R.E., Awramik, S.E. (eds.): Microbial sediments. – 187-195, Berlin (Springer) Riding, R. (1991): Classification of microbial carbonates. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 21-51, Berlin (Springer) Schweigert, G. (1996): Vergleichende Faziesanalyse, Paläoökologie und paläogeographisches Umfeld tertiärer Süßwasserkarbonate auf der westlichen Schwäbischen Alb und im Hegau (Baden-Württemberg). – Profil, 9, 1-100 Further reading: K029 Lacustrine carbonates Allen, P.A., Collinson, J.D. (1986): Lakes. – In: Reading, H.G. (ed.): Sedimentary environments and facies. Second edition. – 63-94, Oxford (Blackwell) Anadon, P., Carrera, L., Kelts, K. (eds., 1991): Lacustrine facies analysis. – Spec. Publ. Intern. Ass. Sed., 13, 318 pp. Anderson, R.Y., Dean, W.E. (1988): Lacustrine varve formation through time. – Palaeogeogr., Palaeoclimat., Palaeoecol., 62, 215-235, Amsterdam Dean, W.E., Fouch, T.D. (1983): Lacustrine environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 97-130 Forester, R.M. (ed., 1987): Paleo-lacustrine theme issue. – Palaios, 2, 412-522 Eugster, H.P., Hardie, L.A. (1978): Saline lakes. – In: Lerman, A. (ed.): Lakes. Chemistry, geology, physics. – 237-293, Berlin (Springer) Friedman, G.M., Krumbein, W. (1984): Hypersaline exosystems. – Ecological Studies, Analysis and Synthesis, 35, 500 pp., Berlin (Springer) Hakanson, L., Jansson, M. (1983): Principles of lake sedimentation. – 316 pp., Berlin (Springer) Katz, B.J. (ed., 1990): Lacustrine basin exploration. Case studies and modern analogs. – Mem. Amer. Ass. Petrol. Geol., 50, 340 pp. Kelts, K., Talbot, M. (1990): Lacustrine carbonates as geochemical archives of environmental change and biotic/abiotic interactions. – In: Tilzer, M.M., Serruya, C. (eds.): Large lakes: ecological structure and function. – 288-314, Berlin (Springer) Matter, A., Tucker, M.E. (eds., 1978): Modern and ancient lake sediments. – Int. Ass. Sedimentologists Spec. Publ., 2, 290 pp., Oxford Picard, M.D., High, L.R. (1972): Criteria for recognizing lacustrine rocks. – In: Rigby, J.K., Hamblin, W.K. (eds.): Recognition of ancient sedimentary environments. – Soc. Econ. Paleont. Min. Spec. Publ., 16, 108-145
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Picard, M.D., High, L.R. (1981): Physical stratigraphy of ancient lacustrine deposits. – In: Ethridge, F.G., Flores, R.M. (eds.): Recent and ancient nonmarine depositional environments: models for exploration. – Soc. Econ. Paleont. Min. Spec. Publ., 31, 233-259 Platt, N.H., Wright, V.P. (1991): Lacustrine carbonates: facies models, facies distribution and hydrocarbon aspects. – In: Anadon, P., Carrera, L., Kelts, K. (eds.,): Lacustrine facies analysis. – Spec. Publ. Intern. Ass. Sedimentologists, 13, 55-74, Oxford Renault, R.W., Last, W.M. (eds., 1994): Sedimentology and geochemistry of modern and ancient saline lakes. – Soc. Econ. Paleont. Min. Spec. Publ., 50, 334 pp. Serruya, C., Pollingher, U. (1983): Lakes of the warm belt. – 569 pp., New York (Cambridge Univ. Press) Talbot, M.R., Kelts, K. (eds., 1989): The Phanerozoic record of lacustrine basins and their environmental signals. – Palaeogeogr., Palaeoclimat., Palaeoecol., Special Issue, 70, 304 pp., Amsterdam Torgersen, T. De Deckker, P., Chivas, A.R., Bowler, J.M. (1986): Salt lakes: a discussion of processes influencing paleoenvironmental interpretations and recommendation for future studies. – Palaeogeogr., Palaeoclimat., Palaeoecol., 54, 7-19 Wright, V. P. (1990): Lacustrine carbonates. – In: Tucker, M., Wright, V.P. (eds.): Carbonate sedimentology. – 164-189, Oxford (Blackwell) Further reading: K025, K026, K027 Fluvial carbonates Galli, G., Sarti, C. (1989): Morphology and microstructures of Holocene freshwater-stream cyanobacterial stromatolites (Villa Ghigi, Bologna, Italy). – Rev. Paleobiol., 8, 3949 McGannon, D.E. (1975): Primary fluvial oolites. – J. Sed. Petrol., 45, 719-727 Nickel, E. (1985): Carbonates in alluvial fan systems. An approach to physiography, sedimentology and diagenesis. – Sed. Geol., 42, 83-104 Ordonez, S., Garcia Del Cura, M.A. (1983): Recent and Tertiary fluvial carbonates in Central Spain. – In: Collinson, J.D., Lewing, J. (eds.): Modern and ancient fluvial systems. – Int. Ass. Sedimentologists Spec. Publ., 6, 485-497 Smith, N., Rogers, J. (1999): Fluvial sedimentology VI. – Int. Ass. Sedimentologists Spec.Publ., 28, 488 pp., Oxford Verrecchia, E.P., Freytet, P., Julien, J., Baltzer, F. (1997): The unusual hydrodynamical behaviour of freshwater oncolites. – Sed. Geol., 113, 225-243 Palustrine carbonates Freytet, P., Plaziat, J.-C. (1982): Continental carbonate sedimentation and pedogenesis - Late Cretaceous and Early Tertiary of southern France. – Contributions to Sedimentology, 12, 213 pp., Stuttgart Monty, C.V., Hardie, L.A. (1976): The geological significance of freshwater bluegreen algal marsh. – In: Walter, M.R. (ed.): Stromatolites. – Developments in Sedimentology, 20, 447-477, Amsterdam (Elsevier) Platt, N.H. (1992): Fresh-water carbonates from the Lower Freshwater Molasse (Oligocene, western Switzerland). sedimentology and stable isotopes. – Sed. Geol., 78, 81-99 Platt, N.H., Wright, V.P. (1992): Palustrine carbonates and the Florida Everglades: towards an exposure index for the
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freshwater environment. – J. Sed. Petrol., 52, 1058-1071 Wright, V.P., Platt, N.H. (1995): Seasonal wetland carbonate sequences and dynamic catenas: a re-appraisal of palustrine limestones. – Sed. Geol., 99, 65-71 Further reading: K 030
2.4.2 Transitional Marginal-Marine Environments: Shorelines and Peritidal Sediments Marginal-marine environments (deltas, beaches and barrier-islands, estuaries, coastal lagoons and tidal flats) occur in a narrow nearshore zone along the boundary between continental and marine depositional realms. The shoreline is marked by great environmental instability resulting from intensive interaction between highenergy forces related to waves, tides, wind, and currents and constantly changing sea levels. Water energy, sediment transport and resulting sediment structures change across the shoreline from offshore to shoreface and foreshore environments. Carbonate deposition takes place in the coastal shoreline zone at the beach, in coastal lagoons behind barriers, and within the peritidal zone. Sediments formed in arid and humid shoreline environments differ in depositional patterns, composition and biota (Sect. 15.5). Key areas for the understanding of modern carbonate shoreline sedimentation are the Yucatan Peninsula, Florida Bay, the Bahamas, Belize shelf, the Arabian Gulf, and western Australia. Excellent reviews of the environments, depositional patterns and facies sequences in these areas are given by Inden and Moore (1983) and Tucker (1990). Particularly handy overviews of modern peritidal carbonates were presented by Shinn (1983), Hardie and Shinn (1986), and Pratt et al. (1992).
2.4.2.1 Beach (Foreshore), Barriers and Coastal Lagoons The beach is strongly wave-dominated and constantly changing. Sediments are produced by the influx from rivers, erosion of cliffs and headlands, erosion of the sea floor, accumulation of shells and the deposition of unconsolidated sediment by tides. A common feature of low-ltitude warm-water beaches is the presence of beachrock (Scoffin and Stoddard 1983). Beachrock originates from the rapid cementation of sand grains and gravels by aragonite and Mg-calcite crystals growing in intergranular pores. Beachrocks are formed both in tropical (see Pl. 3/7, 8; Pl. 33/3, Pl. 35/1, 2) and temperate environments (Pl. 31/3, 4, Pl. 92/2).
Transitional Marginal-Marine Sediments
Modern carbonate coastal and offshore environments can be described by two models, the beach-barrier island-lagoon model and the beach-strandplain model (Reinson 1984, Tucker 1990). The first model develops inregions with moderate to high wave energy, relatively high tidal range and high carbonate sand production resulting in barriers and sand shoals and ridges. The beach-strandplain model is also characterized by high wave energy and high sand production rate, but low tidal range. Most modern carbonate shoreface-foreshore-backshore environments are dominated by skeletal grainstones composed of skeletons of organisms living in these environments and in small coastal patch reefs; ooid grainstones are common in some regions. Coastal lagoons differ in sediment production and biota. Protected lagoons with varying salinities behind beach barriers (islands or off-shor sand bars) are sites for the deposition of fine-grained sediments forming packstone through to mudstones; the biota are low-diversity. Open lagoons connected to the open sea via tidal inlets and with a normal salinity contain a diverse fauna and a variety of sediment types. Freshwater marshes, tidal flats, brackish water and salt ponds or sabkhas may occur adjacent to the lagoons.
2.4.2.2 Peritidal Environments Peritidal carbonates are sediments formed ‘around the tides’ (Folk 1973, Wright 1984) and include deposits formed in supratidal, intertidal and shallow subtidal areas (Box 2.3). The term describes sedimentation on lowenergy tidal zones, especially on tidal flats accreting from the shorelines of land areas or around islands, and on shelves, platforms and ramps. Hardie and Shinn (1986) defined ten basic features characterizing tidalflat deposition: (1) Carbonate tidal flats occur in settings which are protected from open ocean waves. Protection from the open ocean can be caused by wide shelf lagoons, dampening the incoming waves (e.g. Andros tidal flats on the Great Bahama Banks), a position behind barrier islands that separates back reef lagoons from the open ocean (e.g. Abu Dhabi, Persian Gulf), or within restricted embayments (e.g. Shark Bay, Australia). (2) The basic physiography of the tidal-flat system is determined by the scale of the tidal ranges. (3) Tidal flats are subdivided into three depositionalecological zones by the daily oscillations of the tides (Box 2.3). (4) Tidal flats are partly sea and partly land. They are affected by steep gradients of environmental con-
Peritidal Encironments
25
Box 2.3. Criteria of peritidal and subtidal zones. Supratidal zone (supralittoral): Shore zone above normal high tide, flooded only during storms and semi-monthly high spring tides. The zone includes marshes, mangrove forests and diverse coastal ponds of different salinity. May become evaporitic in arid or semiarid climates. Evaporitic supratidal flats are called sabkhas. The width of this zone depends on the relief and dip of the coasts. Intertidal zone (littoral, eulittoral): Near-coast zone between normal high- and low-tide levels which is alternatively flooded and exposed on a diurnal or semi-diurnal basis. Transitional area from marine to terrestrial conditions, dotted with brackish and saline ponds and dissected by subtidal creeks. This zone exhibits a variety of biological, sedimentary and diagenetic criteria. Because many of these criteria are preserved in the geological record, ancient intertidal carbonates are of major importance as key markers in facies studies. Subtidal zone (sublittoral, circalittoral): Marine zone below low-tide level, seaward of the tidal flats, extending down to the edge of the continental shelf and underlying the neritic zone. Water depths between a few decades of meters and 100 to 200 m. Includes low- and high-energy environments. The shallowest parts can be influenced by tidal currents and may be briefly exposed by semi-monthly neap tides. The zone is illuminated and includes shallow and deep subtidal subenvironments. Subdivisions are predominantly based on the distribution and composition of benthic biota.
ditions resulting in an extreme rich variety of sedimentary, early diagenetic and paleontological features recorded in carbonate rocks. (5) Owing to the large number of combinations of circulation, salinity, tides and climate, care must be taken, to use just a single ‘model’ in describing ancient tidal deposits. (6) Tidal flats are wave-protected and, therefore preferential sites for the accumulation of mud. Offshore mud is transported shoreward onto tidal flat currents and storms, and mud can also be produced on supratidal flats. (7) Carbonate tidal flats nucleate on a continental mainland (Persian Gulf), isolated islands with a lagoon (Florida Bay), and platform-margin islands (Great Bahama Banks). (8) Warm and clear waters of shallow tropical environments are conducive to the profilific growth of algal and microbial mats which influence carbonate deposition by binding, trapping and precipitation in supraand intertidal zones. (9) Early cementation is crucial in the preservation of primary structures and textures. (10) Carbonate tidal-flat facies typically occur as stacked cycles, each ideally composed internally of an upward succession of subtidal to intertidal to supratidal units. Peritidal environments exhibit a high degree of variability in relation to the tidal range, mean sea level, topographic position and meteorological factors acting within each zone. To overcome these problems, Ginsburg et al. (1977) have proposed the use of an ‘exposure index’ defined as the percentage of the time that a specific area is exposed. This approach also can be applied to ancient tidal carbonates (Sect. 15.5).
Basics: Modern transitional environments Ginsburg, R.N. (ed., 1975): Tidal deposits, a case book of Recent examples and fossil counterparts. – 428 pp., Berlin (Springer) Hardie, L.A. (ed. 1977): Sedimentation on the modern tidal flats of Northwest Andros Island. – John Hopkins University Studies in Geology, 22, 202 pp. Hardie, L.A., Shinn, E.A. (1986): Carbonate depositional environments, modern and ancient. Part 3. Tidal flats. – Colorado School of Mines Quarterly, 81, 1-74 Inden, R.F., Moore, C.H. (1983): Beach environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Asociation of Petroleum Geolists Memoirs, 33, 211-265 Pratt, B.R., James, N.P., Cowan, C.A. (1992): Peritidal carbonates. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 303-322, Ottawa (Geological Association Canada) Shinn, E.A. (1983): Tidal flat environments. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists Memoir, 33, 171-210 Shinn, E.A. (1986): Modern carbonate tidal flats: their diagnostic features. – Colorado School of Mines Quart. 81, 735 Neumeier, U. (1999): Experimental modelling of beachrock cementation under microbial influence. – Sedimentary Geology, 126, 35-46 Wright, V.P. (1984): Peritidal carbonate facies models: a review. – Geological Journal, 19, 309-325 Wright, V.P.: (1990): Peritidal carbonates. – In: Tucker, M.E., Wright, V.P.: Carbonate sedimentology. – 137-164, Oxford (Blackwell) Further reading: K019, K020
2.4.3 Shallow-Marine Sedimentary Environments: ‘Shallow’ and ‘Deep’ Geologists reconstructing ancient depositional environments like to distinguish between marine ‘shallow-water’ and ‘deep-water’ carbonates, but have a lot of
26
Shelfs, Ramps, Platforms
‘Deep sea’ as used in marine biology to mean that part of the marine environment that lies below the level of effective light penetration for phytoplankton synthesis in the open ocean (photic zone), and deeper than the depth of continental shelves (corresponding approximately to > 200 m). The upper part of the deep sea retains some light (disphotic zone). The permanently dark region below sunlit waters is the aphotic zone which starts in temperate waters at rather shallow depths (about 100 m), and in tropical waters at deeper depths (about 600 m).
2.4.3.1 Pericontinental vs Epicontinental Shallow Seas
Fig. 2.5. Rimmed and non-rimmed carbonate shelves (platforms) and ramps (slightly modified after James and Kendall 1992). Note the different position of the ‘carbonate factories’ at the seafloor. Carbonate factories (James 1979) are subtidal areas characterized by high carbonate production by predominantly benthic organisms. The optimum autochthonous carbonate production on rimmed platforms occurs near the platform margin and behind the margin. On ramps optimum areas of carbonate production are distributed over the entire extension of the ramp. Rimmed and unrimmed platforms are common in tropical and subtropical sunlit waters, today approximately 30° N and S of the equator, where the carbonate factory is primarily controlled by high water temperatures favoring phototroph carbonate-secreting organisms. Ramps are common in cool-water zones, extending polward from the limit of the tropical factory to polar latitudes, and characterized by the dominance of heterotroph organisms (Sect. 16.4). Schlager (2000) differentiated a third carbonate factory (mud mounds), characterized by the in-situ production of biotically induced and abiotic carbonate mud.
trouble doing so. Microfacies analysis assists in overcoming these difficulties (see Sect. 12.3). Common criteria used by marine geologists and biologists in distinguishing ‘shallow-marine’ from ‘deepmarine’ environments in modern oceans are the shelf break and the lower boundary of the well-illuminated zone. Another important boundary level is the storm wave weather base (see Box 2.1). Shallow seas occur above the continental shelf and include marginal-marine and marine environments, both in pericontinental and epicontinental settings. Deep-marine sea bottom environments are located on the slope, at the toe-ofslope and within basins.
Most modern shallow seas cover shelves around the margins of the continents. These pericontinental seas differ strongly from the epicontinental (epeiric) setting of many ancient shallow seas. Epeiric seas occupied extensive areas of the continental interior or the continental shelf. They differ from pericontinental seas with regard to wave and current regimes, sediment input and sea level controls. Modern examples of epeiric shelf seas are rare in contrast to the geological record, where epeiric carbonate platforms and ramps were common. 2.4.3.2 Carbonate Shelves, Ramps and Platforms Carbonate depositional models developed in the 1960s and early 1970s were strongly based on modern analogues from the Bahamas, Florida, the Yucatan and the Persian Gulf (Wilson 1975). These models describe sedimentation patterns on shelves characterize by the deposition of shallow-water sediments on flat platform tops, and by significant rims marked by reefs or shoals near the shelf break. Ahr (1973) noted that the ‘rimmed shelf’ model does not fit many ancient carbonate successions, and that a different model was necessary, especially for the epeiric shallow-water carbonates (Irwin 1965). The term carbonate ramp was adopted to describe a gently sloping depositional surface which passes gradually without slope break from a shallow, high-energy environment to a deeper, low-energy environment. A ramp is attached to a shoreline at one end and to correlative basinal beds at the other end. Fig. 2.5 exhibits the basic differences between rimmed and unrimmed shelves (platforms) and ramps with respect to the position of the optimum carbonate production (carbonate factory) and sediment transport.
Shelves, Ramps, Platforms
27
Carbonate Platforms: Broad Categories Rimmed carbonate shelves (Ginsburg and James 1974): Shallow pericontinental flat-topped platforms. Outer wave-agitated edge characterized by morphological rims (barrier reefs, shoals, islands) that absorb wave action and a pronounced break of slope into deeper waters. Rimmed shelf margins comprise accretionary, bypass and erosional margins. High-energy facies occurs predominantly at the outer shelf margin. Widths typically a few to about 100 km. Modern examples: Queensland Shelf off eastern Australia with the Great Barrier Reef, South Florida Shelf, Belize Shelf, Gulf of Suez. Non-rimmed carbonate shelves: Shallow platforms without a pronounced marginal barrier. Characterized by open water circulation. The term includes ‘open carbonate platforms’ as well as carbonate ramps. Widths a few to more than 100 km. Modern examples: western Florida, Yucatan, Brazil. Atlantic, common in cool-water settings. Homoclinal ramps (Read 1982): Shallow epeiric and pericontinental platforms characterized by gentle uniform depositional slopes, passing downwards from a shallow nearshore high-energy facies offshore into a more muddy deeper-water facies without a marked break in slope. The angle of slope is commonly less than 1° (of the order of a few meters per kilometer), but steeper dips may occur. Width between 10 and > 100 km. Modern examples: Trucial Coast of the Arabian Gulf, Shark Bay of Western Australia. Distally steepened ramps (Read 1982): Similar to homoclinal ramps, but with a distinct increase in gradient in the outer, deep ramp region. Width between 10 and 100 km. Modern examples: Northeastern Yucatan, western Florida.
Epeiric platforms (Shaw 1964): Very extensive, flat, cratonic areas covered by shallow seas. Dominated by shallow subtidal-intertidal low-energy facies and tidal flat cycles. The oceanward margin can be gentle (ramp-like) or steep (shelf-like), and can be rimmed. Width between 100 and 10 000 km. No good modern examples of epeiric platforms exist. Isolated carbonate platforms: Isolated or detached shallow-water platforms, offshore from continental shelves surrounded by deep water. Platform margins partly with reefs and sand shoals, platform interior with low-energy facies and tidal flats. Most isolated platforms have steep margins and slopes into deep and very deep water. Widths between 10 and 100 km. Modern examples: Bahama Bank, Glovers reef/Belize. Oceanic atolls: Formed on extinct, subsiding volcanoes that rise several hundreds to thousands of meters from the deep ocean floor. Characterized by raised reefal rims, steep outer slopes and low coral islets encircling shallow and deep lagoons with pinnacle reefs. The circular or elliptical structures are surrounded by deep water of the open sea. Atolls range in diameter between less than 1 km to more than 130 km. Modern examples: Maldive Islands, Indian Ocean. Common in the western and central Pacific Ocean. Drowned carbonate platforms (Schlager 1981): Rapid sea-level rises, strong subsidence, or drastic reduction in carbonate productivity as result of environmental stress can cause complete or incipient drowning of rimmed shelves, ramps and isolated platform. Pelagic platforms are drowned fragments of an ancestor shallow-marine platform. Modern examples: Atolls in the Pacific and Indian Ocean, Blake Plateau north of the Bahamas.
Fig. 2.6. Carbonate platforms: Broad categories.
Shelves and Ramps
28
Box 2.4. Lateral classification of carbonate shelves (Wilson and Jordan 1983) and carbonate ramps (Burchette and Wright 1992). Carbonate shelves Inner shelf: Near-coast tide-dominated zone including peritidal and shallow subtidal environments varying and restricted salinity; sluggish circulation; biota low-diverse. Mid-shelf: Extended shallow subtidal zone between the near-shore area and the shelf break; below fair-weather wavebase, but above storm-wave base; mud-dominated but with grainy storm sediments; water depths between a few tens of meters and 100 to 200 m; normal marine, but different conditions in local restricted areas; biota high-diverse. Outer shelf: Rimmed shelves: A narrow zone near the shelf break, with shoals and reefs. Non-rimmed shelves: A wide zone below normal storm-wave base which may be affected by intruding ocean currents. Carbonate ramps Inner ramp: Between upper shoreface (beach or lagoonal shoreline) and fair-weather wave base; seafloor more or less constantly affected by wave agitation; includes shoreline deposits, sand-shoals, and back-barrier peritidal sediments. Mid-ramp: Between fair-weather wave base and storm-wave base. The bottom is frequently reworked by storm waves and swells. Sediment composition and textures reflect proximal-distal trends. Outer ramp: Below normal storm-wave base, down to the basin plain. Mud-dominated, but with few storm beds. In deeper zones, restricted bottom conditions may develop.
Today the terms shelf, ramp and platform are used in describing the setting and depositional geometry of shallow-marine carbonates, but some confusion still exists with regard to the definition of these terms: Originally, the term platform described shallow-marine carbonate bodies with flat tops and steep flanks, formed by the accumulation of sediment on the shelf or in the ocean. Today the term often is used rather loosely for thick, often flat-topped sequences of shallow-marine carbonates. Steep-sided flanks are not necessarily definition criteria. Many authors regard carbonate platform as a very general term that includes ramps, shelves and various types of flat-topped platforms or they use the term if they cannot yet make an assignment to any of these categories. ‘Carbonate shelf’ characterizes a carbonate depositional system which develops a constructional relief above the sea floor and is bordered on the shoreward side by marginal-marine or continental sediments, and offshore by slope and basinal sediments. The transition from shallow water to basin occurs over a relatively short distance and is marked by a distinct break in the slope. In this definition the term means the same as ‘attached platform’ a term, sometimes used for platform settings starting at the coast and continuing offshore. One of the major tools of microfacies studies is the understanding of ancient facies belts and depositional areas related to particular marine settings (Chap. 14). This aim requires the differentiation of well-defined categories. Fig. 2.6 exhibits sketches and definitions of broad categories of shallow-marine depositional settings. The figure illustrates definitions used by Ahr (1973, 1998), Ginsburg and James (1974), Schlager
(1981), Enos 1983), Wilson and Jordan (1983), Read (1982, 1985), Harris et al. (1985), Wright and Burchette (1998), and relies on the excellent reviews by Tucker (1985), Tucker and Wright (1990), Jones and Desrochers (1992), and Burchette and Wright (1992). Lateral subdivision of shelves and ramps Carbonate shelves and homoclinal carbonate ramps are laterally subdivided into at least three parts (Box 2.4). Recognizing the various parts of ramps (Fig. 2.7) is important for using the ‘standard microfacies’ approach (Sect. 14.3). The differentiation of inner and mid ramp environments can be strongly improved if the degree of storm, wave, or tidal influence is considered. The identification of sediments, deposited in a ‘basin’ adjacent to the outer ramp of homoclinal or distally steepened ramps is difficult. The lack of coarse tempestites and the absence of turbidites can be used as diagnostic criteria. Sediments in shallow basins, but also in restricted basins and outer ramp settings may be strongly bioturbated (and, therefore, be mistaken for lagoonal facies). In distally steepened ramps the slope break is commonly located around the mid- or outer ramp. The corresponding ‘outer slope’ would form an appropriate additional subdivision of ramps (Burchette and Wright 1992).
2.4.3.3 Shelf Margins The boundary between the shelf and the continental slope is marked by the ‘shelf break’, a point at the shelf edge where a significant change in gradient occurs in the outermost edge of the continental shelf (Stanley and
29
Reefs
Fig. 2.7. Generalized subdivision of carbonate ramps. Compare with Box 2.4.
Moore 1983). Carbonate shelf-slope breaks are stationary, offlap or onlap. Many carbonate shelf and ramp margins are characterized by reefs or by sand banks (Wilson 1974; Halley et al. 1983; James and Mountjoy 1983) which are the line source provenance of most deep-water resedimented carbonates (Schlager and Chermak 1979). Platform margins are subdivided into depositional margins, by-pass margins and rimmed margins. Depending on the geometry of the shelf margin and the declivity of the slopes, depositional and by-pass margins and slopes can be differentiated. Depositional margins exhibit a low relief from platform to slope and basins; by-pass margins are characterized by a steeper relief triggering the deposition of most platform-derived sediment on lower slopes, toe-ofslopes and in adjacent basin plains. For the recognition of margin types with microfacies see Sect. 15.7.
2.4.3.4 Reefs Reefs are singled out from the broad spectrum of marine environments because of the dominating role of benthic organisms in the formation of these structures. Modern reefs originate predominantly in shallow-marine environments,but are also formed in deeper positions on the slope and even in basins. Contradictory long-held opinions that reefs are more or less confined to tropical and subtropical, low-latitude warm-water zones and the new knowledge that modern and ancient reefs can also be formed in mid- and high-latitude tem-
perate and cold waters has been increasing (see Sect. 16.4). Reefal carbonate production is generally considered as an essential factor in the formation of carbonate platforms (Bosscher and Schlager 1993; Kleypas 1997), but this opinion has recently been questioned (Kiessling et al. 2000). Reefs and their eminent role in the development of carbonate depositional systems are discussed in Chap. 14. Box 2.5 contains a glossary of some of the (too) many terms used in the context of reef studies. Basics: Modern shallow-marine carbonate environments Ahr, W.M. (1973): The carbonate ramp: an alternative to the shelf model. – Transact. Gulf Coast Ass. Geol. Soc., 23, 221-225 Ahr, W.M. (1998): Carbonate ramps, 1973-1996: a historical review. – In: Wright, V.P., Burchette, T.P. (eds.): Carbonate ramps. – Geol. Soc. London Spec. Publ., 149, 7-14 Burchette, T.P., Wright, V.P. (1992): Carbonate ramp depositional systems. – Sed. Geol., 79, 3-57 Crevello, P.D., Wilson, J.L., Sarg, J.F., Read, J.F. (eds., 1989): Controls on carbonate platform and basin development. – Soc. Econ. Paleont. Min. Spec. Publ., 44, 405 pp. Enos, P. (1983): Shelf environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 267296 Halley, R.B., Harris, P.M., Hines, A.C. (1983): Bank margin environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 463-506 Harris, P.M., Moore, C.H., Wilson, J.L. (1985): Carbonate depositional environments. Modern and ancient. Part 2: Carbonate platforms. – Colorado School of Mines Quart., 80, 1-60
Tropical and Non-Tropical Carbonates
30
Box 2.5. Glossary of terms used in distinguishing major reef types. Bioherm: Mound or lens-shaped reefal buildup. Biostrome: Tabular rock body, usually a single bed of similar composition. Laterally extended, dense growth of skeletal organisms. No depositional relief. A rigid framework may or may not be present. Buildup: A carbonate rock mass that is thicker than laterally equivalent strata, and probably stood above the sea floor during some or all of its depositional history. The term is often very loosely used for reefs, banks or thick massive limestone structures. Ecologic reef: An ancient reef interpreted as having been built by organisms into a rigid, wave resistant, topographic high on the sea floor (Dunham 1970). Framework reef: Built by organisms forming a rigid calcareous frame. Microbial mound: Biogenic mounds, formed by the action of microbes which initiate carbonate precipitation, and bind and trap sediment (James and Bourque 1992). Mound: A rounded hill-like structure. In the context of reef studies used for counterparts of framework reefs. See microbial mounds, skeletal mounds and mud mounds. Mud mound: Mud-dominated carbonate buildups (Wilson 1975). Organisms are minor constituents. Syndepositional relief. Reef: Laterally confined biogenic structures, developing due to the growth or activity of sessile benthic organisms and exhibiting topographic relief. This broad definition covers framework reefs, reef mounds, mud mounds as well as biostromes (Flügel and Kiessling 2002). Reef mound: Lenticular carbonate bodies consisting of bioclastic mud with minor accounts of organic binding (James 1980). Skeletal organisms are common, but there is no evidence for a prominent in situ skeletal framework. Lime mud/carbonate cement and skeletal organisms are about equally important. Syndepositional relief. Skeletal mound: Biogenic mounds made of small delicate skeletal or encrusting organisms that are thought to baffle, trap, bind and stabilize lime mud (James and Bourque 1992). Skeletal reef: Corresponds to framework reefs with organisms, forming a rigid calcareous framework. Stratigraphic reef: A thick, laterally restricted mass of carbonate rock, without genetic connotations (Dunham 1970). Hottinger, L. (1989): Conditions for generating carbonate platforms. – Mem. Soc. Geol. Ital., 40, 265-271 Irwin, M.L. (1965): General theory of epeiric clear water sedimentation. – Amer. Ass. Petrol. Geol. Bull., 49, 445-459 James, N.P., Bourque, P.-A. (1992): Reefs and mounds. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 323-348, Ottawa (Geol. Ass. Canada) James, N.P., Kendall, A.C. (1992): Introduction to carbonate and evaporite facies models. – In: Walker, R.G., James, N.P. (eds.): Facies Models. Response to sea level change. – 265-275, Ottawa (Geol. Ass. Canada) James, N.P., Mountjoy, E.W. (1983): Shelf slope break in fossil carbonate platforms: An overview. – Soc. Econ. Paleont. Min. Spec. Publ., 33, 189-206 Jones, B., Desrochers, A. (1992): Shallow platform carbonates. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 277-301, Ottawa (Geol. Ass. Canada) Read, J.F. (1982): Carbonate platforms of passive (extensional) continental margin-types, characteristics and evolution. – Tectonophysics, 81, 195-212 Read, J.F. (1985): Carbonate platform facies models. – Amer. Ass. Petrol. Geol. Bull., 69, 1-21 Schlager, W. (1981): The paradox of drowned reefs and carbonate platforms. – Geol. Soc. Amer. Bull., 92, 197-211 Schlager, W. (1989): Drowning unconformities on carbonate platforms. – Soc. Econ. Paleont. Min. Spec. Publ., 44, 1525 Schlager, W. (1992): Sedimentology and sequence stratigraphy of reefs and carbonate platforms. – Continuing Education Course Notes, 34, 71 pp. Schlager, W. (2000): Sedimentation rates and growth potential of tropical, cool-water and mud-mound carbonate systems. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems: components and interactions.
– Geol. Soc. London Spec. Publ., 178, 217-227 Scholle, P.A., Bebout, D.G., Moore, C.H. (eds., 1983): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 708 pp. Sellwood, B.W. (1996): Shallow-marine carbonate environments. – In: Reading, H.G. (ed.): Sedimentary environments and facies. – 283-356, Oxford (Blackwell) Tucker, M.E. (1985): Shallow-marine carbonate facies and facies models. – In: Brenchley, P., Williams, B.P.J. (eds.): Sedimentology, recent developments and applied aspects. – 147-169, Oxford (Blackwell) Tucker, M.E., Wilson, J.L., Crevello, P.D., Sarg, J.R., Read, J.F. (eds., 1990): Carbonate platforms. Facies, sequences and evolution. – Intern. Ass. Sedimentologists Spec. Publ., 9, 328 pp., Oxford Wilson, J.L. (1975): Carbonate facies in geologic history. – 411 pp., New York (Springer) Wright, V.P., Burchette, T.P. (1996): Shallow-water carbonate environment. – In: Reading, H.G. (ed.): Sedimentary environments: processes, facies, stratigraphy. – 325-394, Oxford (Blackwell) Wright, V.P., Burchette, T.P. (eds., 1999): Carbonate ramps: an introduction. – Geol. Soc. London Spec. Publ., 149, 1-5 Further reading: K018, K019, K020, K190
2.4.4 Tropical and Non-Tropical Carbonates: Different in Composition, Controls and Significance For a long time modern shelf carbonates have been considered to be restricted predominantly to warm-water tropical environments. However, extensive shelf car-
Tropical and Non-Tropical Carbonates
31
Fig. 2.8. Map exhibiting the occurrence of selected tropical and non-tropical shelf carbonates. There are, of course, much more records (see maps provided by Wilson 1975, Rao 1996, and James 1997). The sites shown in the figure were selected with regard to their importance in comparative facies analyses. For numbers see Tab. 2.1. Base map after Nybakken (1993).
bonates are forming in all latitudes and at warm, cool and cold water temperatures. Modern non-tropical carbonates occur in tidal flats, on shelves and platforms as well as in the deep sea, similar to tropical shelf carbonates. Tropical carbonates are relatively well understood and are, therefore, discussed rather briefly in this chapter. The following text concentrates chiefly on nontropical shelf and reef carbonates, because microfacies analysis will have a high potential in the future for recognizing and evaluating ancient temperate and coolwater limestones.
2.4.4.1 Latitudinal Zonation and Diagnostic Criteria of Tropical and Non-Tropical Carbonates On the basis of mean surface ocean temperatures and overall organism distribution, three major biogeographical zones can be established whose boundaries are roughly parallel to latitudes: Tropical, temperate and polar (Fig. 2.8). Surface waters in the tropical zone are warm throughout the year, but in the temperate zone warm only in the summer. The thermocline (a zone of rapid temperature change between warm surface waters and cooler deeper waters) is a permanent feature of the tropical zone, occurs seasonally in temperate waters and is absent in polar zones since the ocean surface is covered with ice in winter and solar radiation is low in summer.
Subzones are differentiated according to surface water temperatures (and bottom water temperatures in the polar zone) as well as the latitudinal range (Fig. 2.9). The boundaries between the zones and subzones are transitional rather than absolute. Some trouble exists in the precisely defining of subzones and the use of the terms warm, temperate, cool and cold, particularly in the context of facies studies. Warm-temperate and cool-temperate subzones are rather well-defined (Briggs 1974) in contrast to the subtropical zone. Some authors consider this zone to be an equivalent of the warm-temperate zone; others define the subtropical one as a zone which borders the tropics. Here the climate is intermediate in character between tropical and temperate, though more like the former than the latter. Modern shelf carbonates and reefs are not limited to low-latitude tropical and subtropical warm-water environments, but also occur in temperate cool-water, and even polar cold-water ‘non-tropical’ settings of midto high-latitudes. Non-tropical settings can be loosely defined as lying poleward of the present limits of hermatypic corals reefs, or beyond the mean annual surface-water isotherm of about 20 °C, today typically near the 30° latitude. Carbonate production also takes place along a transitional spectrum of environments between tropical and non-tropical settings depending on ocean circulation and upwelling patterns. The northern, central and southern parts of the Gulf of California, Mexico, offers a unique possibility for studying carbonate proText continued on p. 34
32
Occurrence of Modern Shelf Carbonates
Table 2.1. Modern shelf carbonates (except coral reefs) in non-tropical (ct - cold-temperate, p - polar, wt - warm-temperate) and tropical (tr) settings. The table lists sites that are particularly relevant for the understanding of ancient shelf carbonates and the interpretation of microfacies data. Numbers refer to the numbers in the map (Fig. 2.8).
ct
1
ct
2
p
3
p
4
Sites
Reference
Remarks
Alexander Archipelago (Alaska) Eastern Canadian shelf: Newfoundland, Nova Scotia Vesterisbanken, central Greenland Sea Spitsbergen Bank (Barents Sea) Southwestern Svalbard shelf
Hoskin 1969
Skeletal sands, inner shelf
Barnhardt and Kelly 1995
Carbonate-rich sediments, spongebryozoan mounds Submarine seamount; nearly year-round sea ice cover; sponge-bryozoan mounds Largest open-shelf cold-water carbonate platform in the Arctic region; balaniddominated carbonates; bioclastic carbonates under glaciomarine conditions Open-shelf carbonates; discussion of bioerosion and preservation potential of bioclastic cold-water deposits; siliceous sponge mounds, Maerl, intertidal and subtidal communities; open shelves; deep shelves
ct 5
Western Canadian shelf: Vancouver Island, Hecate Strait
ct 6
Northwestern Europe: Scottish shelves: Rockall Bank, Orkney Islands; Ireland, Porcupine Bank; English Channel; Southern England; Northern Brittany Peninsula, France Northern Norway: off Troms
ct
7
Henrich et al. 1992; Henrich et al. 1995 Bjorlykke et al. 1978; Freiwald 1995; Andruleit et al. 1996; Henrich et al. 1997 Nelson and Bornhold 1983; Young and Nelson 1988; Carey et al. 1995; Krautter et al. 2001 Bosence 1980; Scoffin et al. 1980; Farrow and Fyfe 1988; Scoffin 1988; Freiwald 1993; Henrich and Freiwald 1995; Schäfer et al. 1996; Light and Wilson 1998; Wehrmann 1998; Bader 2001 Freiwald 1995, 1998; Henrich and Freiwald 1995
wt 8
Mediterranean Sea: Southern French coast; Balearic Islands, Spain; Malta Island; Gulf of Naples, southern Italy, Tyrrhenian Sea, off Southern Sicily
Walther 1910; Bosence 1985; Laborel 1987; Bellan-Santini et al. 1994; Canals and Ballesteros 1997; Fornos and Ahr 1997; Basso 1998
wt 9
Adriatic Sea, Italy: Gulf of Trieste; Apulian shelf Off Algeria and Tunisia
Zuschin and Piller 1994; Toscano and Sorgente 2002 Caulet 1972; Burollet 1981
Southern Gulf of California (Mexico) Bermuda Island
Halfar et al. 2000
wt 10 wt 11 tr 12 wt 13 tr 14
Off northwest Africa Belize-Yucatan, central America
tr 15
Bahamas; Florida
Ginsburg and Schroeder 1973; Kuhn 1984; Curran and White 1995 Rad et al. 1982 Logan et al. 1969; James and Ginsburg 1979; Viniegra-Osario 1981; Gischler and Lomando 1999 Purdy 1963; Ginsburg 1956; Multer 1977; Schlager and Ginsburg 1981; Ginsburg 1991; Jones and Desrochers 1992; Curran and White 1995; Reid et al. 1995; Carew and Mylroie 1995
Near-shore coastal platforms and fjords; coralline algal reefs; deep shelf; bioclastic sands produced by carbonatesecreting organisms in kelp forests ‘Coralligene’, maerl; ramp Note: On this map the Mediterranean is indicated as cold-temperate. Many authors consider the Mediterranean a predominantly warm-temperate region (Briggs 1974) Mixed siliciclastic-carbonate facies; relict sediments and modern shelf carbonates Relict and modern calcareous sands and bioclastic carbonate mud Detailed microfacies analysis ! Semitropical conditions adjacent to the temperate realm due to the Gulf Stream Rimmed shelf Tropical ramp (Campeche Bank); protected shelf lagoon (Belize); open shelf (Yucatan) Isolated platform/bank (Great Bahama Bank); unrimmed and rimmed shelves (western and southern Florida), subtropical setting
Occurrence of Modern Shelf Carbonates
33
Table 2.1. continued: Modern shelf carbonates tr 16
ct 17 tr 18 tr 19
tr 20
tr 21
tr 22 te 23 tr 24 tr 25 tr 26
tr 27 tr 28
tr 30 wt 31
wt 32
Caribbean Sea: Caicos (British West Indies), Grand Cayman, Jamaica, Nicaragua Rise, St.Croix (U.S. Virgin Islands), Barbados Northeastern Brazilian shelf Gulf of Suez; Gulf of Aqaba/Elat Red Sea: Bay of Safaga (Egypt); Sudanian coast; Saudi Arabian coast Persian Gulf and Trucial Coast
Multer and Gerhard 1974; Hubbard et al. 1981; Liddell et al. 1984; Wanless and Dravis 1989; Tongpenyai and Jones1991; Glaser and Droxler 1991; Geister 1992; Triffleman et al. 1992; Vecsei 2001 Jindrich 1983; Carannante et al. 1988 Friedman 1985; Reiss and Hottinger 1984; Sellwood and Netherwood 1984 Montaggioni et al. 1986; Piller and Mansour 1990; Brachert and Dullo 1990; Piller 1994; Schuhmacher et al. 1995; Dullo et al. 1996 Purser 1973; Loreau 1982; Friedman 1995
Western Indian Ocean: Mayotte (Comoro Islands), La Réunion, Seychelles Maldives Islands (Indian Ocean) South China Sea
Masse et al. 1989; Lewis 1969; Dullo et al. 1996; Camoin and Davies 1998; Purdy and Winterer 2001 Purdy and Bertram 1993; Ciarapica and Paseri 1993; Aubert and Droxler 1996 Li Desheng 1984; Guozhong 1998 Boichard 1985 Arp et al. 1996; Kempe et al. 1996 Emery et al. 1954; Davies et al. 1991; Camoin and Davies 1998; Vecsei 2001 Logan 1974; Playford 1990 Maxwell 1968; Scoffin and Tudhope 1985; Davies et al. 1989; Orpin et al. 1999 Agegian and Mackenzie 1989
Kalimatan (Borneo) Satonda Crater Lake, N’ Sumbawa, Indonesia Western Pacific: Coral Sea, Queensland Plateau Shark Bay (W. Australia) Great Barrier Reef (Australia)
Eastern Pacific: Hawaian Archipelago Southeastern Brazil
ct 34
Off southwestern Australia: Spencer Gulf Off southern and southeastern Australian margin: Great Australian Bight: Lacepede shelf, Otway Tasmania
ct 35
New Zealand
p 36
Ross Sea (Antarctica): McMurdo
p 37
King George V coast (Antarctica)
ct 33
Summerhayes et al. 1976; Alexandersson and Milliman 1981 Collins 1988 Collins 1988, James et al. 1992; Boreen et al. 1993; Bone and James 1993; James et al. 1994 Rao and Adabi 1992; Rao and Nelson 1992; Rao 1994; Rao and Amini 1995 Carter 1975; Nelson et al.1988; Rao and Nelson 1992 Taviani et al. 2001
Domack 1983
Isolated platforms, shallow lagoons, reefs, carbonate banks
Cool-water shelf carbonates in tropical regions ! Near-coastal sediments, reefs Rimmed shelf, sedimentation adjacent to coral reefs
Arid subtropical, homoclinal ramp (Trucial Coast, Arabian Gulf), reefs, atoll lagoons; open shelf (Persian Gulf) Ramps, platforms, atoll lagoons, reefs
Oceanic atolls, banks Some cold-water carbonates because cold currents reduce sea-water temperatures Platform ‘Marine’ lake with increased alkalinity Banks, platforms, reefs
Shallow embayment, living stromatolites Largest modern reef province in the world, rimmed and unrimmed shelf, protected shelf lagoon, coral reefs, Halimeda banks, subtropical setting Mid-depth carbonate banks Extended carbonate ramps, effect of upwelling and low terrigenous input on carbonate production Rimmed shelf One of the world’s largest modern accumulations of ramp-type coastal and shelf carbonates, high-energy open shelf Shelf carbonates Platform and shelf carbonates Production and accumulation of skeletal grains on continental shelves and offshore banks at depths < 350 m Sponge-bryozoan mound
34
duction in cool-temperate, warm-temperate and tropical/subtropical environments (Halfar et al. 2001). The distribution of modern tropical warm-water belt and the northern and southern cold-water belts is controlled by the global circulation patterns of surface currents and the intensity and availability of nutrient recycling between surface and deep waters (Henrich and Freiwald 1995). The rocky coasts of open shelf coldwater belts are characterized by kelp-forests. Kelps are giant seaweeds (brown algae). The thalli provide habitats for numerous carbonate-secreting organisms (cirripeds, bryozoans, serpulids, crustose coralline algae, sponges). These epibionts contribute considerably to carbonate production not only on coastal platforms, but also in intrashelf troughs and coastal lagoons due to the drift of algal blades or the transport of pebbles attached to the algae (kelp-drafting). Modern tropical and non-tropical carbonate sediments exhibit many biological and sedimentological criteria (Table 2.2) which are used in discriminating warm- and cool-water carbonates in the context of facies analyses of ancient limestones (see Sect. 16.4). One of the most significant features concerns the biotic carbonate production in near-coastal shelves:
NON-TROPICAL Carbonates Heterozoan
Latitudinal Sea-water range temperature
TROPICAL Carbonates Photozoan
,
Occurrence of Polar, Temperate and Tropical Carbonates
POLAR CARBONATES
>50° N and S
Cold water <5 -10 °C (mean) -1.5 to16 °C (range)
Subtropical and tropical coasts are fringed by coral reefs. Non-tropical temperate rocky coasts of the Northern Seas are characterized by huge kelp forests which provide habitats for sediment-producing epibionts (predominantly barnacles, bivalves, echinoids, and bryozoans).
2.4.4.2 Tropical and Subtropical ShallowMarine Carbonates Shelf carbonates The major settings of subtropical carbonate-producing shelves fall into two categories: (1) Rimmed shelves (= protected shelf lagoons), and (2) open shelves (Fig. 2.10). These categories can be subdivided into those shelves which are influenced by terrigenous input, and those which are not (Sellwood 1986). This classification fits many ancient carbonate shelves. Nevertheless, some caution is necessary in using recent shelf carbonates as analogues for ancient shelf and platform carbonates: Modern shelves have been strongly influenced by the Quaternary sea-level fluctuations, particularly by the erosional unconformity formed during the last
Subdivision
Latitudinal range
Sea-water temperature
polar
>60° N and S (to >70° N)
>5 °C
subpolar
>50° to <60° N and S
5 - 10 °C
Beyond the Arctic Circle: Central Greenland Sea, Barents Sea, Ross Sea, Antarctica Arctic eastern Canada, Northern Norway, Western Canadian Shelf Northwestern Europe
TEMPERATE CARBONATES
30° - 50° (60°) N and S
30°N to 30°S TROPICAL CARBONATES
Cool water ~10 -18 °C (mean)
cooltemperate
30° to 50° N and S
>10 to 25 °C (range)
warm temperate
25° to > 30° N 10 - 18 °C 25° to 30° S
Warm water
subtropical
18 - 22 °C
Bahama, Florida, Bermuda, Persian Gulf, Shark Bay
tropical
>22 °C
Great Barrier Reef, Indian Ocean, Pacific
5 - 10 °C
18 to >22 °C (mean) 18 to 30 °C (range)
Southern Australia, Tasmania, New Zealand Mediterranean Sea, Off North Africa, Southwestern Australia
Fig. 2.9. Latitudinal distribution and critical sea-water temperatures of modern tropical carbonates and temperate and polar carbonate settings. Note that cool-water carbonates can also form in tropical regions, where cold currents reduce sea-water temperatures (e.g. off the east coast of SouthAmerica, off the west coast of Africa and off the coasts of southern Asia).
Comparing Tropical and Non-Tropical Carbonates
35
Table 2.2. Main features of modern (warm-water) and non-tropical (cool- and cold-water) carbonates (based on Leonard et al. 1981, Nelson 1988, Rao 1997). Bold-type criteria are of specific value in microfacies studies of ancient non-tropical limestones.
Latitude Environment Tectonics Terrigenous input Geometry of buildups Mean water temperature Salinity Carbonate saturation Water circulation Accumulation rates Reefs Algal mats Biodegradation Carbonate content Texture Carbonate mud Origin of allochthonous mud Ooids Aggregate Grains Peloids Intra/Lithoclasts Calcareous red algae Rhodoids Calcareous green algae Stromatolites Seaweed (brown algae) Seagrass Zooxanthellate corals Azooxanthellate corals Coral growth forms Benthic foraminifera Bryozoans Bivalves Vermetid gastropods Serpulid worms Barnacles Echinoderms Siliceous sponges Skeletal grain associations Biotic diversity Equatability Association with non-carbonates Glauconite Phosphorite Biogenic silica Low-Mg calcite High-Mg-calcite Aragonite Dolomite Early sea-bed lithification Diagenetic trends
Tropical carbonates
Non-tropical carbonates
30° N - 30° S Platform; shelf Stable Low Attached and isolated platforms with marginal rims > 23 °C Normal to hypersaline Supersaturated to saturated Restricted to open 50 to >1000 cm/1000 a Common to abundant Common Common to abundant >90% Muds, sands, gravels Common to abundant Disintegration of algae, microbial Common Common Common Common Common to abundant Common Abundant Common to abundant Rare Abundant Abundant Rare Diverse Common, high-diversity Rare to common Common and diverse Rare Rare to common Rare Common Rare to common Few High Low Rare Rare Rare Rare Rare to common Abundant Abundant Common Common to abundant Constructive
beyond 30° N and 30°S Open shelf; ramp Stable to unstable Low to high Homoclinal and distally steepened ramps <20 °C Normal (to reduced) Saturated to undersaturated Open >1 to >10 cm/1000 a Rare to absent Absent Abundant, strong 50-100% Sands, gravels, muds Rare to abundant Physical and chemical degradation of skeletons Absent Absent Rare or absent Rare to common Abundant Abundant Very rare Rare Abundant Less common Rare Common to abundant Solitary and fasciculate Abundant, low-diversity Abundant Abundant and diverse Common Common Common to abundant Abundant Common to abundant Diverse Moderate to high High Rare to abundant Common Common Rare to common Common to abundant Abundant to common Rare Rare Less common Destructive
Subtropical Shelf Carbonates
36
Clastic-free rimmed shelves Clastic-influenced rimmed shelves
Great Bahama Bank Anselmetti et al. 2000; Ball 1967; Curran and White 1995; Eberli and Ginsburg 1989; Ginsburg et al. 1991; Hardie 1977; Harris 1979; Hine 1977; Milliman et al. 1993; Newell et al. 1959; Neumann and Land 1975; Purdy 1963a; 1963b; Schlager and Ginsburg 1981; Shinn et al. 1969; Wanless and Dravis 1989; Ginsburg 2001 South Florida Shelf Bosence 1989; Enos and Perkins 1977; Ginsburg 1956; Halley et al. 1983; Messing et al. 1990; Multer 1969; Shinn 1963; Shinn 1983; Shinn et al. 1989; Stockman et al. 1967; Turmel and Swanson 1976 Cuba Bandy 1964; Daetwyler and Kidwell 1959; Bathurst 1975 Belize Gischler and Lomando 1999; Gischler et al. 2000; James and Ginsburg 1979; Purdy 1974; Purdy et al. 1975; Scholle and Kling 1972 Great Barrier Reef Drew 1983; Gagan et al. 1990; Larcombe and Woolfe 1999; Maiklem 1968; Maxwell 1968; Maxwell and Swinchatt 1970; Scoffin and Tudhope 1985; Tudhope 1989
Clastic-free open shelves
Sea floor gently inclined towards the shelf break; sedimentation controlled by wave action, and tidal and oceanic currents; calcareous sands in shallow settings; mud in deeper settings; wide irregular facies belts; lack of resedimentation by mass flow mechanism
Examples with important references
Northwestern Yucatan Folk 1967; Folk and Robles 1964; Gischler and Lomando 1999; Hoskin 1963; Logan et al. 1969; Viniegra-Osario 1981; Wilson 1970; Ward and Brady 1973
Clasticinfluenced open shelves
RIMMED SHELVES (Protected shelf lagoons)
Shallow sea floor (~100 m average depth); topographical fringing barriers (reefs, shoals, islands) at the seaward margin; low-energy environments and mud deposition within the shelf lagoon, high-energy environments predominantly at the outer margins
OPEN SHELVES
Main features
Eastern Gulf of Mexico Folk et al. 1962; Gardulski et al. 1986; Rezak et al. 1985; Siringan and Anderson 1994 Persian Gulf Houbolt 1957; Patterson and Kinsman 1981; Purser 1973; Scholle and Kinsman 1974
Fig. 2.10. Major settings of modern subtropical shelf carbonates (based on Sellwood 1986).
glacial stage. Facies distribution and setting of modern reefs and shoals, therefore, are strongly controlled by the Pleistocene topography (Purdy 1974). Excellent reviews and detailed descriptions of modern tropical and subtropical shelf carbonates from different settings can be found in Sellwood (1986), Tucker and Wright (1990), Jones and Desrochers (1992), and Wright and Burchette (1996). In the context of interpreting of microfacies data, these references must be thoroughly consulted. Composition of tropical and subtropical shallow-marine carbonates: Similar to ancient limestones, modern warm-water shelf carbonates consist of polygenetic carbonate mud and different categories of grains. Grains include biogenic material (detrital bioclasts or in situ calcareous organisms), sand-sized grains of mudgrade carbonate (peloids), various coated grains (ooids, oncoids, rhodoids), sand-sized aggregates of agglutinated particles (grapestones), and lithified and reworked sediment clasts (lithoclasts, intraclasts). The
definitions and significance of these constituents are discussed in Sect. 4.2. Plate 3 offers a first look at the diversity of the grain types. Reefs Modern tropical and subtropical reefs are situated at shelf and platform margins, and on shelves, platforms and ramps. The distributional pattern of recent coral reefs suggests that reefs are arranged nearly symmetrically in both hemispheres, do not occur at latitudes greater than 34°, are concentrated at the western margins of the oceans, are absent in the vicinity of large river mouths, and exhibit two reef realms (Caribbean and Indo-Pacific realm). Furthergoing explanation of the distributional pattern suggests that reefs prefer shallow waters and tropical to subtropical climates, and can grow in land-locked settings and in open oceanic settings, and that reefs avoid settings with high clastic input and areas affected by cool eastern boundary currents and coastal upwelling. In summary, sea-water temperature, sunlight
Modern and Pleistocene Carbonates
and water depths, and nutrients and siliciclastic input are the major controls on the distribution of modern zooxanthellate coral reefs. Cool-water reefs are discussed in Sect. 16.4.1.
2.4.4.3 Non-Tropical Shelf and Reef Carbonates Shallow-marine temperate carbonate sediments constitute about one third of modern global shelf sediments (Nelson 1988). Modern temperate cool-water carbonates differ from tropical carbonates in the type of skeletal grains, mineralogy and diagenesis, oxygen and carbon isotopes and in the range of trace elements. Nontropical and tropical shallow-marine carbonates differ significantly in their prevailing geometries: The scar-
37
city of rapidly growing calcareous frame-building organisms and the absence of aragonite (which would favor early lithification) hinder the development of platforms with wave-resistent margins and steep flanks. Non-tropical shelf sediments are primarily controlled by bottom currents which distribute the bioclastic sediment over extended areas. The results are homoclinal ramps, developing in low-water energy, and distally steepened ramps in high-energy environments. Non-tropical carbonates are found almost continuously from mid to high latitudes in shelf areas with low or missing clastic input. Similar settings which occur along a latitudinal transect from warm to cold water conditions, may be developed along downslope profiles at any latitude (including tropical carbonate platforms, Brazil shelf).
Plate 3 Modern and Pleistocene Carbonates: Shallow-Marine and Windblown Sands; Beachrocks The samples shown on this plate document the common composition of calcareous sands deposited in shallowmarine (–> 1-3) and subaerially exposed environments (–> 4, 5, 7, 8). Both the Bahamas and the Bermudas offer excellent possibilities for studying the controls on the deposition of calcareous sediments and the processes involved in their transformation into carbonate rocks. The Bahamas are a large isolated platform, 700 x 300 km in size. The subtidal sediments exhibit specific distributional patterns (see map on next page). Lithofacies characterized by skeletal grains (coralalgal), ooids and reefs (see Pl. 51/7) occur at the platform margins, whereas sediments rich in aggregate grains (grapestone) or carbonate mud (pellet mud) are platform-interior facies. Modern shallow-marine carbonate sands formed in tropical and subtropical warm-water environments consist predominantly of skeletal grains (foraminifera, –> 6; calcareous algae, –> 5; mollusk shells, –> 2, and other biota), associated with non-skeletal grains including peloids (–> 2), ooids (–> 1, 3) and composite grains represented by aggregate grains (–> 1; Pl. 15/1-3), and reworked and redeposited sediment (intraclasts, –> 1). Grains are usually aragonite or High-Mg calcite. The average size ranges between < 0.5 and about 2.0 mm. Sand accumulations are common on or near the seawards edge of banks, platforms or shelves (Halley et al. 1983). Sand production and the development of modern bank-margin sand bodies are strongly controlled by antecedent topography, physical processes (wind, waves, tides, and storms), and sea-level changes. Various encrusting organisms contribute to the stabilization of the sediment binding of grains and to the formation of beachrocks (–> 8). Submarine cementation within inter- and intragranular pores as well as early diagenetic cementation of subaerially exposed sediments lead to an incipient lithification (–> 4). The microstructure of many grains becomes highly altered due to the activity of microborers (–> 2; ‘micritization’) and diagenetic alterations. These grains are formed in specific high- and low-energy environments (see map). Eolianites, consisting of windblown marine carbonate grains, are known from modern and ancient tropical and subtropical coastal environments (see Fig. 15.1). Eolianites are commonly associated with coastal marine deposits and paleosols. Because the sediment consists of grains, originally formed in shallow-marine environments, eolian grainstones may be confused with marine calcarenites. Important diagnostic criteria of eolianites are lamination and stratification with high- to low-angle beds (see Fig. 15.1), fine- to medium-sized grains, good overall sorting (–> 4), sorting differences between laminae or beds, inverse grading, rounded tabular grains, and meteoric vadose meniscus and microstalactitic calcite cements (–> 4). Other cement types are fibrous cements consisting of long calcite crystals, and micrite rim cement that superficially resemble micrite envelopes. Calcified plant root systems in calcarenites associated with paleosols and carbonate crusts (see Pl. 128) suggest a pause in sedimentation under subaerial conditions.
38
Plate 3: Description of Figures
Plate 3 Text Continued: Modern and Pleistocene Carbonates: Shallow-Marine and Windblown Sands; Beachrocks 1 Constituents of recent marine calcareous sands: Ooids (O), aggregate grains (grapestones, G; see Pl. 15/1-3), intraclasts and peneroplid foraminifera (P). Ooid nuclei are large peloids and skeletal grains. Holocene: Exuma Kays, Great Bahama Bank. 2 Skeletal grains (forminifera, algal debris, Halimeda HA) and micritized grains. Note the crossed-lamellar microstructure of the large aragonite gastropod shell. Parts of the shell were destroyed by microborers (arrows; see Pl. 52/1-4). The sediment in the shell interior was cemented prior to deposition. Subtidal unstable sand of the coralgal facies. Holocene: Bimini Island, Great Bahama Bank. 3 Ooids. Tangentially structured ‘normal ooids’ (TO) exhibiting a cortex of several aragonite lamellae and ‘superficial ooids’ (SO), characterized by only a few laminae (see Pl. 13). Nuclei are large peloids. Note the small dark micrite-filled spots within some ooids (bottom left) representing microborings. The Bahamian ooid shoals occur along the platform margin. The Joulters ooid shoal north of Andros Island consists of ooid grainstone (formed on current-swept, rippled sea bottoms), ooid packstone (formed on stabilized burrowed bottoms), fine-peloid packstone (formed on stabilized and burrowed bottoms), pellet wackestone and lithoclast packstone (formed in active tidal channels) as well as skeletal grainstone (Harris 1983). Holocene: Joulters Cays, Great Bahama Bank. 4 Constituents of a poorly cemented wind-blown ooid sand (ooid grainstone) forming eolian dunes. Tangentially structured ooids. Cementation has taken place in the meteoric vadose zone. Note rims of Low-Mg calcite cements (arrows) connecting the grains. These ‘meniscus cements’ and phreatic granular calcite cements indicate freshwater diagenesis (see Pl. 32/56). Cementation affects densely packed grains more strongly than loosely packed grains. Most pores between the grains are open (but were filled with epoxy during the preparation of the thin section). The grains were formed within a marine environment, but an eolian deposition is indicated by the good overall sorting and vadose cements (Caputo 1995). Late Pleistocene (Grotto Beach Formation; age about 125 000 a): San Salvador Island, Bahamas. 5 Skeletal grains, predominantly fragments of the green alga Halimeda (see Pl. 57/1, 2), few foraminifers and gastropods. The disintegration of Halimeda and other green algae is a major process in the production of carbonate mud (see Pl. 7/ 1, 2). Late Pleistocene (French Bay Member, Grotto Beach Formation): San Salvador Island, Bahamas. 6 Lithified calcareous sand (calcarenite), deposited in a nearcoastal marine environment during an interglacial phase. Consisting predominantly of skeletal grains (peneroplid foraminifera), peloids, and grains that have been intensively attacked by microborers and transformed into microcrystalline calcite (dark). Late Pleistocene (Belmont Formation): Devonshire Bay, southern Bermuda. 7 Fossil beachrock composed of large micritized aggregate grains (AG) and some skeletal grains (bored shells: arrow points to borehole; foraminifera, F). In contrast to –> 8, all the intergranular pores are filled with cement. Beachrock is formed in the intertidal zone of subtropical or tropical coasts. Late Pleistocene (Grotto Beach Formation): Same locality as –> 5. 8 Juvenile beachrock consisting of well-preserved fragments of Halimeda (HA), composite grains and peloids. Note the still open voids (V). The grains are encrusted by foraminifera (arrows). Beachrock originates from the growth of aragonite and Mg-calcite cements in intergranular pores. Cementation can take place within a few years. Late Pleistocene (Grotto Beach Formation): Same locality as –> 5. Map and cross-section of Andros Island. Modified from Friedman and Sanders (1978) and Blatt et al. (1972).
Figs. 1–3: Loose grains are embedded in resin which appears white in transmitted light.
Plate 3: Modern and Pleistocene Carbonates
39
40
Studies of mid- and high-latitudinal non-tropical carbonates started rather late in the 1970s, but ‘temperate’ and ‘cold-water’ carbonates are becoming continuously better documented (Nelson 1988; Henrich and Freiwald 1995; Rao 1996; James and Clarke 1997) and have become one of the most exciting tools in carbonate research. The importance of non-tropical coral reefs was underlined by Teichert (1958). Chave (1967) pointed out that shelf carbonates could form at all latitudes, regardless of water temperature. Based on a thorough analysis of modern shelf sediments, Lees and Buller (1972) recognized the fundamental differences between tropical and temperate carbonates, and the importance of water temperature and salinity as major controls on the composition of shelf carbonates. The authors emphasized the significance of ‘grain association patterns’ for distinguishing between tropical and non-tropical carbonates. This concept provides a high potential for evaluating the climatic and oceanographic controls on ancient carbonates using microfacies data (Sect. 12.2). The last twenty years have produced a wealth of papers dealing with non-tropical coastal and shelf carbonates in the Northern and Southern Hemisphere and comprising a latitudinal range of 30 to 80° N and 30 to 70° S. Temperate and cold-water shelf carbonates were studied in the northern Atlantic (northwestern France, England, Ireland, Scotland, Norway), in the Greenland and Barents Sea, on the Canadian Scott Shelf, from various parts of the Mediterranean Sea (Tunisia, Sicily, Adriatic Sea, Eastern Mediterranean) as well as from Pacific regions (Western Canada, India). In the Southern Hemisphere extensive temperate carbonates are forming today off Australia, Tasmania and New Zealand on tidal flats, gulfs, shelves, platforms and ramps. The largest modern temperate-water carbonate shelf production occurs on the outer shelf, shelf edge and upper slope of the southern and southeastern margin of Australia (Boreen and James 1993). The eastern Tasmanian shelf is an excellent example for the deposition and diagenesis of cool-water carbonates. These carbonates have been extensively studied with regard to compositional, mineralogical and geochemical criteria and are, therefore, a reference standard for recognizing and interpreting ancient non-tropical carbonates. Informative reviews and collections of case studies are available (Nelson 1988; Henrich and Freiwald 1995; James and Clarke 1997). Rao (1997) gives an excellent overview on the present state of knowledge and includes a comprehensive bibliography. Controls on temperate carbonates: Major controls on the origin and composition of temperate carbonates are (a) the rate of dilution from ter-
Temperate and Polar Carbonates
rigenous sources, (b) appropriate sea-water temperatures (which may be relatively ‘warm’ – the Mediterranean type of parts of southern Australia, or ‘cool temperate’ – e.g. Otway shelf, southeastern Australia; Tasmania), (c) the relative abundance of calcareous biota, (d) the availability of nutrients, (e) sufficient levels of saturation of CaCO3 in seawater (allowing cementation and reducing dissolution), (f) high water energy for accumulating calcareous sands, and (g) the rate of accumulation (approximately an order of magnitude lower in temperate shelves than in tropical shelves). Many non-tropical carbonates form at low CO2 levels in regions of upwelling waters, which maintain CaCO3 saturation and provide nutrients for luxuriant growth of fauna. Lowering of sea level results in wave-base erosion and redistribution of sediments from shallow to deeper water and below the wave base. Temperate and polar carbonates: Bioclastic carbonates of shelf platforms: These sediments are represented by moderately to well-sorted accumulations of coarse and fine bioclastic sands (e.g. bryozoan-, mollusc-bryozoan or coralline algal sands; Pl. 4/2, 3, 5) with minor amounts of fines and gravels (e.g. molluscan gravel). The sands consist mainly of bioclasts, some micritic peloids and angular, rounded and bored intraclasts (reworked submarine hardgrounds or semilithified carbonate). Ooids are absent. Sedimentary particles are permanently reworked and swept basinward along the slopes, where the grains mix with autochthonous or older particles. Many sediments are a mixture of modern calcareous sands and Pleistocene relict carbonates. Australian case studies demonstrate the occurrence of two major facies zones: Above the zone of winnowing the sediment consists of coarse skeletal harsh and some relict particles. Below, the deposits are muds, rich in planktonic material and some benthics (Nelson et al. 1988, James and von der Borch 1991). Carbonate mud sedimentation occurs on the temperate shelf of southeastern Australia (Blom and Alsop 1988). The calcitic muds are produced by the accumulation and disintegration of nannoplankton, as well as from the biodegradation of skeletal grains accumulating on the sea floor. Many high-latitude carbonate muds are bioerosional and result from maceration of shells (Farrow and Fyfe 1988). Bioerosion down to about 25 m is intensive in the photic zone. Shelly substrate is affected by e.g. clionid sponges, endolithic algae, bivalves, regular echinoids, crabs, fish. Bioerosional products are angular silt- and gravel-sized chips, or mud-sized disintegration products of fecal pellets.
Cool-Water Reefs
41
Case studies in the northern Atlantic underline the great importance of subtidal kelp forests, which harbor a great number of carbonate-secreting and sedimentproducing organisms. These forests occur along the wave-exposed flanks of rocky coastal platforms. In contrast, pure coralline red algal gravel deposits (‘maerl’) develop in more protected parts of the platforms. ‘Maerl‘ is a Breton term, and refers to unattached crust-forming branched or nodular coralline algae and algal gravel, including their widely distributed deposits (Pl. 4/2). Maerl is one of the most abundant shallow-water carbonates of cool-temperate environments. Modern maerl beds are found throughout the Mediterranean Sea, and are common on terrigenous-free Atlantic coasts in depths < 20-30 m, from Norway and Denmark in the north to Portugal in the south, extending to Morocco and Mauretania on the African coast. Maerl deposits have been studied in detail in Brittany, northwestern France (Henrich and Freiwald 1995, Bader 2001), British Islands (Farnham and Jephson 1977), western Ireland (Bosence 1985) and in the Mediterranean Sea. The French term ‘coralligène’ refers to recent coralline algal buildups of the Mediterranean (Bosence 1985). These carbonates occur in intertidal and sublittoral rocky coasts in water depths between about 20 and 160 m; they are represented by reefs formed by corallinacean frameworks. The coralligène can be considered to be a good recent analog for some Miocene coralline limestones of the Tethyan region that exhibit a very similar generic composition and similar growth forms of red algae.
Skeletal grains: Common bioclasts are bryozoans (Pl. 4/5), foraminifera, molluscs, echinoids and calcareous red algae. Other organisms include barnacles, serpulids, brachiopods, ahermatypic corals and sponges. The main groups occur in various proportions, depending on the water temperature, salinity, nutrients and grain size of the sediment. Compositional differences allow ‘skeletal grain assemblages’ characterized by the dominance of particular grain types to be defined (e.g. bryozoans and mollusks = bryomol association occurring on deep shelf banks of the northern Atlantic; Pl. 4/4), which are indicative of special environmental conditions and particular parts of the shelf (Sect. 12.2.2.1). Non-skeletal grains (peloids, intraclasts) are rare, ooids are absent. Cool-water reefs: Diverse in composition and distribution: Lower slope deep-water reefs (‘lithoherms’): Neumann et al. (1977) described elongate coral bioherms in water depths of several hundred meters, that are axially aligned with the flow field of the Florida current. The upcurrent end of the mounds consists of dense thickets of branching ahermatypic corals; the middle and downcurrent sections are rocky, and strongly bioeroded with attached sponges, alcyonarians, gorgonians, and crinoids. Sediments from this zone are a mixture of planktonic foraminifers and pteropods, with skeletal grains from lithoherm organisms. Large amounts of aragonite mud are trapped in the coral thicket at the upcurrent ends of the mounds or in sponge-boring cavities within the lithoherms. The description of these
Table 2.3. Generalized comparison of modern warm-water shallow-marine reefs and cold/cool-water deeper-marine coral reefs.
Sea water temperature range Bathymetrical range Light Zonation patterns Nutrional regime Trophic modes Diversity of coral framework builders Diversity of associated fauna Diversity of associated flora Biological framework stabilization Bioerosional impact Diagenetic impact Interskeletal cementation Intraskeletal cementation Internal sediment Automicrite
Warm and shallow water reefs
Cold and deeper water reefs
20 - 28 °C -50 m Photic zone Windward-leeward Light zonation Oligotrophic Photo- and heterotrophy High High High Coralline Algae High High Very common Very common Very common Common
4 - 12 °C 250 m to >1 000 m Aphotic to dysphotic zones Current zonation Taphonomic zonation Eutrophic Heterotrophy Low High Non existent Encrusting sponges High Low Rare Low to common Very common Very common
42
structures started a fascinating hunt for Holocene and ancient ‘deep-water and cold-water reefs’ which is still going on. Non-tropical reefs differ significantly from reefs thriving in shallow warm waters (Table 2.3). High-latitudinal reefs are stressed ecosystems whose biota should be able to tolerate low water temperatures as well as sharp seasonal temperature fluctuations, salinity fluctuations and low light influx. Common reef-building organisms are corals, calcareous red algae, bryozoans, bivalves (e.g. oysters), serpulids and tube-constructing sabellarid polychaetes, vermetids, encrusting foraminifera and sponges. Non-tropical reefs comprise coral reefs, coralline algal reefs, sponge-bryozoan reefs, and mollusk reefs. Examples of coralline algal reefs and maerl pavements have been studied on the northern coast of Brittany (France), and the coast of the Troms District in northern Norway (Freiwald et al. 1991; Freiwald and Henrich 1994). Oyster reefs, vermetid reefs and serpulid reefs occur in different parts of the Atlantic Ocean (Milliman 1974). Corals contribute significantly to the formation of large and extended reef-structures in deeper cold-water settings (Lophelia reefs), but also in relatively shallow warm-temperate environments (Cladocora banks of the eastern Mediterranean: Kühlmann 1996; Taviani et al. 2001). Lophelia reefs: Huge framework reefs built by the azooxanthellate coral Lophelia cover huge areas of the deep and cold eastern Atlantic shelf bottoms (Scottish shelf: Scoffin et al. 1980; Norwegian shelf: Henrich and Freiwald 1995; Freiwald and Wilson 1998; Faroe Island: Frederiksen et al. 1992, see Pl. 4/2). The corals are preserved as frameworks or coarse coral rubble. Hovland and Thomsen (1997) discuss possible causal relations between deep-water coral reefs and hydrocarbon seeps. Polar carbonates and arctic sponge-bryozoan mounds: Parameters other than the rather stable water temperature control polar carbonates influenced by glaciomarine conditions (Domack 1988; Henrich et al. 1997): the seasonality of light levels (polar night, polar summer), winter sea ice coverage versus summer open polar conditions, seasonality in plankton production, and seasonality in terrigenous sediment discharge by glaciers, meltwater plumes, rivers and release of sea ice sediment (Henrich et al. 1992, 1997). The richness and diversity of benthic organisms are controlled by relatively constant (low but stable) temperatures, low fluctuations in salinity, low terrestrial input and fluctua-
Cool-Water Carbonates
tions in light regime, seasonal sea ice cover, ice shelves, currents and circulation patterns. The largest open-shelf cold-water carbonate platform in the Arctic is Spitsbergen Bank in the northwestern Barents Sea (Henrich et al. 1996; Pl. 4/6). The bank corresponds to a >80 000 km2 wide shallow shelf platform bordered by > 400 m deep troughs. Water depths range from less than 25 m to deeper than 120 m. The bank is completely covered by sea ice for 6 months. Melting sea ice causes strong seasonal variations in surface water salinity. The maximum summer temperature is 2 °C on the bank and 4–6 °C in the marginal zone. Another bioclastic carbonate production area is the southwestern Svalbard shelf where skeletal carbonate deposition resting on tillites takes place under glaciomarine conditions in depths between 50 and 80 m (Andruleit et al. 1996). Carbonates are produced by (a) barnacles living on kelp forests and forming mobile sands which are transported over the shallow platform, and (b) benthic communities (infaunal bivalves, balanids, hydrozoans, soft corals, sponges, bryozoans) forming buildups on the deep flanks of the platform. The seamount is covered by extensive sponge-bryozoan bioconstructions exhibiting three distinct facies belts depending on variations in down-slope food transfer. Different communities occurring in these belts are biogenic mats with sponge spicules and bryozoan fragments, sponge bryozoan-serpulid buildups, bryozoan thickets, and sponge-crinoid mounds (Henrich et al. 1992). Basic patterns of sediment distribution are coarse, high-energy bioclastic sands consisting of balanids, bivalves and benthic foraminifera on top of the bank and arctic mollusk coquina/ glaciomarine gravel lag deposits over the flanks. These carbonates differ from other high-latitude shelf carbonates with or without glaciomarine controls. Postglacial or interglacial carbonate sediments capping older glaciogenic units and not influenced by glaciomarine settings are known from the Scottish shelf, the Norwegian shelf, off Nova Scotia and Newfoundland, and western Canada. Carbonate deposition associated with glaciomarine environments occurs on Antarctic shelves (Domack 1988; Taviani et al. 1993). Significance of cool-water carbonates: 1 Quantitative and qualitative data of high-latitude cool-water carbonates are strongly needed in all speculations on regional and global carbonate budgets. 2 Because the kind of skeletal and non-skeletal grains, facies, carbonate mineralogy and diagenesis vary with sea-water temperatures and other factors, the differentiation of tropical and non-tropical carbon-
Basics: Modern Non-Tropical and Polar Carbonates
ates is an important tool in reconstructing paleoclimate and paleolatitudes. Classical concepts, based on grain associations (e.g. foramol, Sect. 12.2) must be refined, taking into account not only temperature and salinity as major controls on carbonate composition, but also the negative influence of terrigenous input as well as the importance of food supply, biogenic interactions e.g. food transfer from seasonal plankton blooms to benthic communities (benthopelagic coupling) and variable oceanic configurations (Betzler 1995). In addition, differentiation between cold-water and tropical carbonates is complicated by the fact that the occurrence and distribution of cold-water carbonates are controlled not only by lateral latitudes but also by the change in warm to cold conditions along downslope profiles (Carannante et al. 1988). These bathymetrically controlled changes may occur at any latitude and need further study. 3 In contrast to modern tropical carbonates, composed predominantly of aragonite and Mg-calcite, many Phanerozoic limestones are formed from calcitedominated carbonates. Modern cold-water carbonates, therefore, may be good analogues for ancient shelf limestones, e.g. from the Paleozoic or Jurassic (James and Bone 1989, 1991; James and von der Borch, 1991; James et al. 1992). 4 Diagenesis and reservoir potential: It is generally assumed that sea-floor and meteoric cementation are very limited in cool-water carbonates as compared with tropical warm-water carbonates. Carbonate minerals are predominantly High-Mg calcite (and minor aragonite) in shallow warm temperate carbonates and High-Mg to Low-Mg calcites in cool temperate carbonates. The mineralogical composition is controlled by water temperatures and water depths. Carbonate saturation of sea water in non-tropical settings is less pronounced or even undersaturated (Opdyke and Wilkinson 1990). Diagenesis, therefore, is predominantly destructive (Alexandersson 1978). Compared to tropical shallow-marine carbonates the diagenetic potential is low and sediments may remain essentially unlithified for a long period of time. Nevertheless, intraskeletal and intergranular carbonate cementation can be common in some temperate and cold-water carbonates (Rao 1996; James et al. 1992, 1994; Henrich and Freiwald 1995). Basics: Modern non-tropical temperate and polar carbonates Andruleit, H., Freiwald, A., Schäfer, P. (1996): Bioclastic carbonate sediments on the southwestern Svalbard shelf. – Marine Geology, 134, 163-182 Bader, B. (2001): Modern Bryomol-sediments in a cool-wa-
43
ter, high-energy setting: the inner shelf off Northern Brittany. – Facies, 44, 81-104 Bone, Y., James, N.P. (1993): Bryozoans as carbonate sediment producers on the cool-water Lacepede Shelf, southern Australia. – Sed. Geol., 86, 247-271 Boreen, T.D., James, N.P. (1993): Holocene sediment dynamics on a cool-water shelf: Otway, southeastern Australia. – J. Sed. Petrol., 63, 574-588 Bosence, D.W. (1985): The Coralligene of the Mediterranean - a Recent analog for Tertiary coralline algal limestones. – In: Toomey, D.F., Nitecki, M.H. (eds.): Paleoalgology. – 215-225, Berlin (Springer) Briggs, J.C. (1974): Marine zoogeography. – 475 pp., New York (McGrawHill) Cairns, S.D., Stanley, G.D. (1981): Ahermatypic coral banks: living and fossil counterparts. – Proc. 4th Int. Coral Reef Symp. Manila, 1, 611-618 Colantoni, P., Cremona, G., Ligi, M., Borsetti , A.M. , Cate, F. (1985): The Adventure Bank (off Southwestern Sicily): a present day example of carbonate shelf sedimentation. – Giorn. Geol. Ser. 3, 47, 165-180 Domack, E.W. (1988): Biogenic facies in the Antarctic glacimarine environment: basis for a polar glacimarine summary. – Palaeogeogr., Palaeoclimat., Palaeoecol., 63, 357372 Farrow, G.E., Fyfe, J.A. (1988): Bioerosion and carbonate mud production on high-latitude shelves. – In: Nelson, C.S. (ed.): Non-tropical shelf carbonates - modern and ancient. – Sed. Geol., 60, 281-297 Fornos, J.J., Forteza, V., Jaume, C., Martinez-Taberner, A. (1992): Present-day Halimeda carbonate sediments in temperate Mediterranean embayments: Fornells, Balearic Islands. – Sed. Geol., 75, 283-293 Freiwald, A. (1993): Coralline algal maerl frameworks - islands within the phaeophytic kelp belt. – Facies, 29, 133148 Freiwald, A. (1998a): Geobiology of Lophelia pertusa in the North Atlantic. – Habilitation Thesis University Bremen, 194 pp. Freiwald, A. (1998b): Modern nearshore cold-temperate calcareous sediments in the Troms District, Northern Norway. – J. Sed. Research, A68, 763-776 Freiwald, A., Henrich, R. (1994): Reefal coralline algal buildups within the Arctic Circle: morphology and sedimentary dynamics under extreme environmental seasonality. – Sedimentology, 41, 963-984 Freiwald, A., Henrich, R., Pätzold, J. (1997): Anatomy of a deep-water coral reef mound from the Stjernsund, West Finnmark, Northern Norway. – In: James, N.P., Clarke, J.A.D. (eds.): Cool-water carbonates. – Soc. Econ. Paleont. Min., Spec. Publ., 56, 141-162 Freiwald, A., Henrich, R., Schäfer, P., Willkomm, H. (1991): The significance of high-boreal to subarctic maerl deposits in Northern Norway to reconstruct Holocene climatic changes and sea level oscillations. – Facies, 25, 315-340 Freiwald, A., Wilson, J.B. (1998): Taphonomy of modern deep, cold-temperate water coral reefs. – Historical Biology, 13, 37-52 Gillespie, J.L., Nelson, C.S. (1997): Mixed siliciclastic-skeletal carbonate facies on Wanganui Shelf, New Zealand: a contribution to the temperate carbonate model. – In: James, N.P., Clarke, J.A.D. (eds.): Cool-water carbonates. – Soc. Econ. Paleont. Min., Spec. Publ., 56, 127-140 Halfar, J., Godinez-Orta, L., Ingle, J.C. (2000): Microfacies
44
Plate 4
Modern Non-Tropical Cool-Water Carbonates
Modern Non-Tropical Cool-Water Carbonates: High-Latitudinal Temperate and Polar Shelf and Reef Carbonates from the Northern Atlantic
Marine carbonate production is by no means restricted to tropical warm-water environments. Abundant and widespread biogenic carbonate production occurs in non-tropical, mid- and high-latitudinal temperate and coolwater settings with low or missing clastic input, and even in cold water beyond the Arctic Circle. Examples are known from the southern and northern hemispheres. Cool-water carbonates record climatic and paleoceanographic information in their sedimentary deposits and architectures. Microfacies and sedimentary architecture of ancient cool-water carbonates provide basic data for the evaluation of paleoclimate and paleolatitudes. Modern coolwater carbonates are good analogues for limestones formed in calcite-dominated Phanerozoic seas. Widely extended carbonate deposits and even reefal frameworks occur in Subarctic and Arctic seas. Most sediments are bioclastic sands and gravels, formed by balanids, mollusks, foraminifera and coralline red algae. Many of these organisms are epibionts, living on benthic macroalgae which form extended kelp forests in nearshore environments. Kelp and coralline algae are adapted to seasonal climatic perturbations. On the Spitsbergen Bank, the largest open-shelf cold-water carbonate platform of the Arctic region, the carbonate production is centered around two main factories: (a) barnacle sands, produced within kelp forests growing on the shallowest part of the platform and transported over the platform, and (b) high biogenic production by infaunal bivalves (–> 6) as well as balanid-hydrozoan-soft coral-sponge-bryozoan buildups on the flanks of the platform (Henrich et al. 1997). Marked differences in the composition of bioclastic grain associations reflect specific environmental conditions on the shelf bottoms; these ‘grain association patterns’ are powerful indications of the climatic, oceanographic and latitudinal controls of ancient cool-water limestones (see Sect. 12.2). Recognizing reefs in deep, aphotic zones is highly important for interpreting ancient analogues. Different reef types are formed by corals, coralline red algae (–> 1), siliceous sponges, bryozoans, and mollusks. These organisms, living in stressed ecosystems, are able to tolerate low water temperatures, and seasonal fluctuations of light intensity, nutrient availability and salinity. Spectacular reef mounds built by the azooxanthellate coral Lophelia (–> 3) originate at present e.g. in eastern Atlantic Ocean of northwestern Europe, at the west Florida and the Bahama Bank slopes, the southeastern Brazilian continental slopes and on eastern Pacific seamounts. Azooxanthellate deep-water corals have the potential to construct large, biologically zoned framework-supported reef mounds below the storm wave base in the aphotic zone (Freiwald 1997). The coral reefs of the Nordic Seas in water depths between 200 and 1200 m can be followed tens of kilometers at lengths and widths of several hundreds of meters. 1 Coralline red algal reef: The reefal framework covers about 125 000 square meters and grows in water depths of 6 to 16 m. Lithothamnion cf. glaciale is one of the most common framework builders. Note the high density of branching common in turbulent waters. The branches are a prominent source for the production of algal gravels and rhodoliths deposited onshore and offshore and producing rhodolith pavements. Troms District, Northern Norway. 2 Near-coast coralline algal gravel (‘maerl’) facies. Multiple fragmentation of branched rhodoliths caused by storms result in the formation of algal fragment accumulations. The maerl, a mixture of carbonate (skeletal) sand and seaweed is used as agricultural fertilizer. It characterizes cool-water temperate sediments, composed of unattached crusted-forming branched or nodular algae and algal gravel. In turbid, high-energy waters of the cool-temperate northwestern European Atlantic, maerl occurs mostly in the photic zone of wave-protected coasts at depths of 1-10 m (Freiwald et al. 1991). In the warmtemperate Mediterranean Sea, maerl is common in depths down to 40 m. Water depth 5 m. Troms District, Norway. 3 Azooxanthellate reef-forming deep-water coral (Lophelia pertusa). Azooxanthellate corals harbour no photosynthetic symbionts. The corals feed on suspended particles which are seasonally generated in the high-productivity euphotic zone. The corallites of Lophelia exhibit strong variations in the branching pattern, the shape and thickness of corallites, the diameter and the intratentacular budding patterns. These variations support the framebuilding potential of the anastomosing coral. Faroe Island Shelf. Water depth 252 m. 4 Open-shelf skeletal carbonate sediment composed of mollusks, bryozoans and serpulids. Malangsgrunnen Bank, off Troms, northern Norway. Water depth 95 m. 5 Selected bioclasts (1–2 mm fraction) of the open-shelf skeletal carbonate shown in –> 4. Most bioclasts are bryozoans. The facies corresponds to the ‘bryomol grain association’ (see Pl. 106/2, 3) which is common on modern cool-temperate shelves. Water depth 90 m. Malangsgrunnen Bank, off Tromso, northern Norway. 6 Arctic-benthic carbonate deposits of the seasonally sea ice-covered Spitsbergenbank (73°30' N and 9°10' W), western Barents Sea. Coquina-gravel lag deposit colonized by abundant bivalves (Mya truncata, – M, note the cylindrical siphones, arrow; Chlamys islandica - C), echinoids (E), hydrozoans, soft corals, balanids, sponges, bryozoans - B) attributed to the initiation of small-scale mounds. The high density of in situ organisms is due to a permanent high productivity zone at the boundary of Atlantic and Arctic surface waters over the flanks of the Spitsbergenbank. Water depth 79 m. –> 1–5: Courtesy of A. Freiwald (Erlangen), 6: Freiwald et al. 1991
Plate 4: Modern Cool-Water Carbonates
45
46
analysis of recent carbonate environments in the Southern Gulf of California, Mexico - a model for warm-temperature to subtropical carbonate formation. – Palaios, 15, 323-342 Henrich, R., Freiwald, A., Schäfer, P. (1997): Evolution of an Arctic open-shelf carbonate platform, Spitsbergen Bank (Barents Sea). – In: James, N.P., Clarke, J.A.D. (eds.): Cool-water carbonates. – Soc. Econ. Paleont. Min., Spec. Publ., 56, 163-181 Henrich, R., Freiwald, A., Wehrmann, A., Schäfer, P., Samtleben, C., Zankl, H. (1996): Nordic cold-water carbonates: Occurrences and controls. – Göttinger Arb. Geol. Paläont., Sonderband, 2, 35-52 Henrich, R., Hartmann, M., Reitner, J., Schäfer, P., Freiwald, A., Steinmetz, S., Dietrich, P., Thiede, J. (1992): Facies belts and communities of the arctic Vestrisbanken Seamount (Central Greenland Sea). – Facies, 27, 71-104 James, N.P. (1997): The cool-water carbonate depositional realm. – In: James, N.P., Clarke, J.A.D. (eds.): Cool-water carbonates. – Soc. Econ. Paleont. Min., Spec. Publ., 56, 1-20 James, N.P., Bone, Y. (1989): Petrogenesis of Cenozoic, temperate water calcarenites, South Australia: a model for meteoric/shallow burial diagenesis of shallow water calcite sediments. – J. Sed. Petrol., 59, 191-204 James, N.P., Clarke, A.D. (eds., 1997): Cool-water carbonates. – Soc. Econ. Paleont. Min., Spec. Publ., 56, 440 pp. James, N., Collins, L.B., Bone, Y., Hallock, P. (1999): Subtropical carbonates in a temperate realm: modern sediments on the southwest Australian shelf. – J. Sed. Petrol., 69, 1297-1321 Lees, A. (1975): Possible influences of salinity and temperature on modern shelf carbonate sedimentation. – Marine Geol., 19, 159-198 Lees, A., Buller, A.T. (1972): Modern temperate-water and warm-water shelf carbonate sediments contrasted. – Marine Geol., 13, M67-M73 Macintyre, I.G. (1988): Modern coral reefs of western Atlantic: new geological perspective. – Amer. Ass. Petrol. Geol. Bull., 72, 1360-1369 Messing, C.G., Neumann, A.C., Lang, J.C. (1990): Biozonation of deep-water lithoherms and associated hardgrounds in the northeastern Straits of Florida. – Palaios, 5, 15-33 Mullins, H.T. (1981): Modern deep-water coral mounds north of Little Bahama Bank: criteria for recognition of deepwater coral bioherms in the rock record. – J. Sed. Petrol., 51, 999-1013 Nelson, C.S. (1988): An introductory perspective on nontropical shelf carbonates. – Sed. Geol., 60, 3-12 Nelson, C.S., Bornhold, B.D. (1983): Temperate skeletal carbonate sediments on Scott Shelf, northwestern Vancouver Island, Canada. – Marine Geol., 52, 241-266 Nelson, C.S., Keane, S.L., Head, P.S. (1988): Non-tropical carbonate deposits on the modern New Zealand shelf. – In: Nelson, C.S. (ed.): Non-tropical shelf carbonates - modern and ancient. – Sed. Geol., 60, 71-94, Neumann, A.C., Kofoed, J.W. , Keller, G.H. (1977): Lithoherms in the Straits of Florida. – Geology, 5, 4-10 Rao, C.P. (1996): Modern carbonates, tropical, temperate, polar. Introduction to sedimentology and geochemistry. – 206 pp., Howrah (Univ. Tasmania) Rao, C.P. (1997): A colour illustrated guide to sedimentary textures, cold, cool, warm, hot; An introduction to the interpretation of depositional, diagenetic and hydrothermal
Deep-Marine Environments
temperatures. – 128 pp., Howrah, Tasmania (Rao, C.P.) Rao, C.P. (1999): Cold water polar aragonitic bivalve elemental composition, east Antarctica. – Carbonates and Evaporites, 14, 56-63 Schäfer, P., Henrich, R., Zankl, H., Bader, B. (1996): Carbonate production and depositional patterns of Bryomolcarbonates on deep shelf banks in mid and high northern latitudes. – Göttinger Arb. Geol. Paläont., Sonderband, 2, 101-110 Scoffin, T.P., Bowes, G.E. (1988): The facies distribution of carbonate sediments on Porcupine Bank, northeast Atlantic. – In: Nelson, C.S. (ed.): Non-tropical shelf carbonates - modern and ancient. – Sed. Geol., 60, 125-134 Smith, A.M. (1988): Preliminary steps toward formation of a generalized budget for cold-water carbonates. – In: Nelson, C.S. (ed.): Non-tropical shelf carbonates - modern and ancient. – Sed. Geol., 60, 323-331 Smith, A.M., Nelson, C.S., Spencer, H.G. (1998): Skeletal carbonate mineralogy of New Zealand bryozoans. – Marine Geol., 151, 27-46 Teichert, C. (1958): Cold- and deep-water coral banks. – Bull. Amer. Ass. Petrol. Geol., 42, 1064-1082 Walther, J. (1910): Die Sedimente der Taubenbank im Golf von Neapel. – Preuss. Akad. Wiss., phys.-math. Kl., Anhang Abh. 3, 1910, 49 pp. Further reading: K021, K210 (glacial carbonates)
2.4.5 Deep-Marine Environments 2.4.5.1 Settings Oceanic deep-marine settings beyond the shelf break bordering pericontinental shelves include the slope, the base-of-slope and continental rise, the abyssal plains, mid-oceanic ridges, oceanic rises and volcanic seamounts, as well as deep-sea trenches. Deep-marine environments, however, also comprise parts of the deeper areas of peri- and epicontinental open shelves and of outer ramps, lying below the storm wave water base, and below the lower boundary of the well-lit zone. The bottom of the modern oceans comprises three bathymetrically defined zones (Fig. 2.2; Box 2.6). The bathyal zone lies to the seaward of the shallower subtidal shelf zone and landward of the deeper abyssal zone. The upper limit is marked by the edge of the continental shelf. The abyssal zone seaward of the bathyal zone covers approximately 75% of the total ocean floor. It pertains the nearly flat deep-sea plains. Abyssal plains are usually covered by thin deposits of pelagic ooze or distal turbidites. The lower limit of the zone is bathymetrically only loosely defined. The deepest area of the sea bottom, the hadal zone, lies specifically within the deep-sea trenches below the general level of the abyssal zone, in depths between >6 000 and more than 10 000 m. These marginal trenches are narrow, steepsided troughs running parallel to continental margins at seaward base of a continental platform.
Sedimentation Processes
47
Fig. 2.11. Sources and pathways of modern deep-sea sediments, deposited in tropical, temperate and boreal (polar) zones. Note that only the radiolarian oozes have a fossil record starting in the Paleozoic. Coccoliths (planktonic algae) and planktonic globigerinid foraminifera have become essential constituents of deep-sea sediments since the Cretaceous. The fossil record of deep-sea carbonates formed by pteropods (gastropods adapted to pelagic life) is scanty and limited to the Tertiary and Pleistocene (Buccheri et al. 1980). The earliest records of diatoms (siliceous algae) date from the Middle Cretaceous, but the group did not become important as sediment builders before the Tertiary. Modified from Hay (1974).
In contrast to modern oceanic basins, many basins found in the geological record occur in cratonic positions and represent intra-platform or intra-shelf basins (e.g. pelagic epeiric basins). Note that the term basin as used by carbonate sedimentologists does not necessarily imply that carbonate basins are analogous to deep oceanic basins. Carbonate basins generally lie well above abyssal depth, commonly less than a few hundred meters below sea level.
ooze, sediments derived from shallow-marine sources and deposited on slopes and basins, deep-sea clays and volcanic material, and glacial-marine sediments produced by ice-rafting (Fig. 2.11). The basic classification of deep-marine sediments considers grain size and composition. Currently used compositional classifications were developed by scientists participating in the DSDP and ODP programs (Berger 1974; Kennett 1982).
2.4.5.2 Sedimentation Processes
2.4.5.3 Pelagic Sedimentation
Sedimentation in the deep sea is controlled by four major processes (Cook and Mullins 1983; Cook et al. 1983; Enos and Moore 1983; Pickering et al. 1986; Stow 1985, 1986): (1) settling of marine pelagic and windblown material from the water column, (2) bottom transport by gravity flows including turbidity currents, debris flows, grain flows, and slumping, (3) redeposition of the sediment by geostrophic bottom currents including contour currents, and (4) chemical and biochemical precipitation on the sea floor. The first process is responsible for the deposition of pelagic sediments, the second one for the formation of a broad spectrum of resediments.
Biogenic deep-sea sediments are formed by the remains of pelagic organisms contributing to the formation of carbonate or siliceous oozes and muds. The term ‘pelagic’ means ‘of the open sea’ and refers to marine nektonic or planktonic organisms whose environment commonly is the open ocean. Large terrigenous sediment input and submarine volcanism cause mixtures of biogenic material and terrigenous clays resulting in the formation of fine-grained muddy, ‘hemipelagic sediments’. The character of the biogenic sediment depends on the supply of the biogenic material, the dissolution at and near the sea-bottom and in the water column, the dilution of the sediment by non-biogenic components, and alterations during diagenesis: The supply depends on the fertility of the pelagic organisms, which in turn is related to the availability of nutrients. Both nutrients and sunlight control the primary production. The dis-
Modern deep-sea deposits comprise biogenic sediments including pelagic siliceous oozes and carbonate
48
Glossary for Deep-Water Carbonates
Box 2.6. Glossary of terms used to describe deep-water carbonates Abyssal plains: Area of deep sea bottom below 2000 m and between 4000 and 6000 m, including the nearly flat deepsea plains covered by sediment. Accretionary slope: A depositional slope, less steep than by-pass slopes. Characterized by net accumulation of sediments. Aragonite compensation depth (ACD): Level in the oceans where aragonite is dissolved. The bathymetrical position of the ACD and the CCD depends on the fertility of the surface waters and the degree of undersaturation of deep waters. Aseismic volcanic ridges: Volcanic seamounts, ridges and plateau, rising several kilometers above the deep ocean floor. Base-of-slope: The lower or lowermost part of slopes, continental rise. Bathyal: Sea bottom between about 200 (continental shelf edge, shelf break) and 2000 (4000) m, encompassing the continental slope and continental rise, down to the deep sea.. Bypass slope: Slopes on which sediment is transported from shallow to deeper water without significant deposition on the slope. Calcite compensation depth (CCD): The level in the deep oceans where the rate of dissolution of calcium carbonate (calcite) balances the rate of deposition and below which carbonate-free sediments accumulate. The level is characterized by a transition from carbonate ooze to deep-marine clay or siliceous ooze. Today, the CCD has a mean depth of about 4.5 km, but varies between ocean basins. Note that large shifts of up to several thousand of meters of the CCD have occurred during the Meso- and Cenozoic. Chalk: Fine-grained porous carbonate rock, predominantly composed of pelagic nanno- and microfossils, and deposited on deeper shelves and basins. Continental rise: Setting at the foot of the continental slope between the slope and the deep ocean basin, characterized by a gently sloping surface created by coalescing submarine fans. Sedimentation takes place by downward sediment transport and deposition at the slope base. The continental rise may be subdivided by submarine canyons and deepsea channels. Contourite: Sediment deposited on the continental rise by contour-following bottom currents. Debris flow: A sediment gravity flow. Sediment transport by a cohesive matrix. Corresponds to mudflow. Erosional slope: A steep slope. Net sediment loss. Erosion due to rockfalls, sediment gravity flows and the effect of contour currents. Grain flow: A sediment gravity flow. Grains are supported during the sediment transport by dispersive pressure arising from grain collisions. Gravity flows: Downslope flow of sediment. Comprises turbidites, grains flows, and debris flows. Guyot: A type of seamount with a platform top. Hadal: Deepest area of sea bottom, lying specifically within the deep-sea trenches below the general level of the abyssal zone, between 6000 and > 10,000 m. Hemipelagic sediments: Fine-grained muddy sediment consisting of a mixture of calcareous biogenic material and terrigenous clay. More than 25% of the fraction coarser than 5 µm is of terrigenous or volcanic origin. The sediment has undergone lateral transport. Deposited near continental margins, on slopes, abyssal plains, and ridges. Lysocline: The level or ocean depth at which the rate of dissolution of calcium carbonate just exceeds its combined rate of deposition and precipitation. The lysocline is characterized by a noticeable decrease in the percentage of deepmarine carbonate sediments and by an interval separating well-preserved from poorly preserved microfossil assemblages. The calcite lysocline is defined by planktonic foraminifera and calcareous nannofossils, the aragonite lysocline by pteropods. Oceanic basin: Comprises various deep-sea settings including abyssal plains, volcanic hills, mid-oceanic ridges and deep-sea trenches. Oceanic plateau: Topographically high area of the ocean floor that rises to within 2-3 km of the sea floor surface above the abyssal floor. Many Pacific plateaus are covered by carbonate ooze overlying volcanic rocks. Pelagic carbonates: Fine-grained calcareous sediments of the open sea, composed chiefly of the skeletal debris of calcareous planktonic organisms, and deposited in a range of deep-water settings (slopes, ocean floor, seamounts, mid-oceanic ridges). Periplatform sediments: Intermixed platform-derived sediments and pelagic carbonate muds (Schlager and James 1978). Resedimented carbonates: Carbonate rocks characterized by redeposition of shallow-water carbonate sediment into deep water. Submarine mass transport and redeposition is related to a variety of processes, including rockfalls, slides, slumps, and gravity flows. Resedimented carbonates may be additionally reworked by deep-marine bottom currents. Rockfalls: Single blocks or many clasts derived from steep slopes of the shelf margins and deposited in talus piles at the base of slopes or as blocks on the slopes. Abundant along fault-bounded platform margins and along fore-reef slopes. Seamount: Elevation of the sea floor, 1000 m or higher, flat-topped (Guyot) or peaked. Slope: Surface which connects the shelf or the shelf break with the continental rise or the deep-sea floor. Turbidite: A sediment laid down by a turbidity current. Characterized by graded bedding, moderate sorting and a vertical sequence of distinct sedimentary structures. Turbidity currents are initiated by the disturbance of the sediment near the shelf margin or on the slope; deposit originally shallow-water sediments on the slope, at the foot of the slope and on abyssal plains. Limestone turbidites are called ‘calciturbidites’ or ‘allodapic limestones’.
Resedimentation
tribution of pelagic sediments on the sea bottoms is controlled by the different dissolution of biogenic grains, produced in the surface waters of the oceans. Most dissolution of siliceous tests (radiolaria, diatoms) occurs in highly silica-undersaturated shallow surface waters. In contrast, most dissolution of carbonate tests takes place on the sea floor as bottom waters become more undersaturated in calcium carbonate. Therefore, siliceous biogenic sediments reflect the biological fertility of the surface waters, whereas carbonate biogenic deep-sea sediments chiefly reflect the preservation potential of carbonate at various depths. The proportions of siliceous and carbonate-rich oozes depend on bottom-water circulation which controls both dissolution patterns, and, through upwelling, fertility patterns (Berger 1970, 1974, 1976). Differences in the composition of pelagic calcareous plankton reflect differences in oligotrophic, mesotrophic and eutrophic nutrient environments (Henrich et al. 2001).
2.4.5.4 Resedimentation (‘Allochthonous Carbonates’) Downslope gravity transport along deep-marine slopes occurs by various processes including rockfall, slumping, sliding, and sediment gravity flows. Subaqueous rockfalls occur in areas with very steep topography producing talus and scree deposits at the base of slopes. A slump is a downslope translocation of a sediment parcel along discrete shear planes. In slides, large blocks move on only a few slippage planes. Both slumping and sliding are controlled and triggered by high rates of sedimentation creating oversteep slopes and differential compaction, structural changes (e.g. effect of salt domes) as well as by earthquakes. In sediment gravity flows, the sediment is transported downwards under the influence of gravity. The sediments are transported by different mechanisms: In ‘grain flows’, the sediment is supported during transport by direct grain to grain interactions. In ‘debris flows’ larger grains are supported by a matrix consisting of a mixture of fine sediment and interstitial fluid. In ‘turbidity currents’, the sediment is supported mainly by the upward component of fluid turbulence. The resulting deposits differ in grain size, sorting, texture and sedimentary structures. Because turbidity currents are episodic, the resulting turbidites are interbedded with finer pelagic sediments representing the background sedimentation.
49
2.4.5.5 Carbonate Plankton and Carbonate Oozes Carbonate oozes cover about 50% of the modern ocean floor. They are deposited on deep-sea plains, submarine rises, ridges and plateaus. Modern carbonate oozes consisting of the shells of minute drifting planktonic organisms comprise foraminiferal ooze, nannofossil ooze and pteropod ooze: Foraminiferal oozes are dominated by tests of sand-sized planktonic foraminifera (Pl. 73/10) occurring within carbonate mud consisting of calcareous nannofossils. Nannofossil oozes or coccolith oozes are very fine-grained muds composed predominantly of remains of planktonic algae (Chrysophyta: Coccolithophorida; Pl. 7/3-5). Pteropod oozes are restricted o shallow tropical areas in depths < 3 000 m. They consist of aragonitic shells of planktonic gastropods (pteropods and heteropods). The carbonate production is greatest in surface waters with high biological productivity caused by upwelling. Production and growh rates of pelagic foraminifera and coccolithophorids are differnt. Plankton contributes to the sedimentary record by (1) settling of individual tests, (2) fecal pelletization, and (3) the deposition of various biogenic materials (including ‘marine snow’, aggregates, flocks and particulate organic matter, POM). Individual settling may play a role in the deposition of larger microfossils (e.g. foraminifera, radiolarians), but fecal pelletization and seasonally varying sinks of detrital aggregates (e.g. phytodetritus displaying an important food resource in the deep sea, Thiel et al. 1988) play the greatest role in the transportation of smaller tests (e.g. coccolithophorids; Honjo 1976). –> A cautionary note: Comparable calcareous plankton is not known from pre-Mesozoic records and comparisons with modern analogues work for pelagic Mesozoic-Tertiary carbonates, but not for Paleozoic slope and basin carbonates (see Sect.15.8).
2.4.5.6 Preservation Potential and Dissolution Levels In modern oeans carbonate sediments are largely absent below depth of about 4500 m. High hydrostatic pressure, low water temperature and high partial pressure of CO2 lead to dissolution and significant diferences in the carbonate preservation potential of deepsa sediments. These differences have been used to distinguish two levels:
50
(1) The lysocline, defined as the depth that separates well-preserved from poorly preserved planktonic assemblages and where a noticeable decrease in the percentage of carbonate occurs. Because calcareous microfossils dissolve at different rates, three types of lysoclines are used: the pteropod or aragonite lysocline (ALy), defining significant dissolution of aragonite material occurring in relatively shallow depths; and the foraminiferal lysocline and the coccolith lysocline (calcite lysocline, CLy), defining the significant dissolution of calcite material in greater water depths. (2) The calcite compensation depth (CCD) is defined as the depths at which calcite microfossils and carbonate are largely absent. The depth of the CCD in modern oceans varies between about 4 200 and 5 400 m, with a mean depth of about 4500 m. At that level, the rate of dissolution balances the rate of accumulation. –> A cautionary note: The depth position of the CCD in Phanerozoic oceans has varied during geological time owing to fluctuations in the amount of carbonate being brought into the oceans from the land, variations in hydrothermal activity, or increased carbonate production on the shelves. Because the calcareous tests of planktonic organisms differ in dissolution susceptibilities, dissolution types can be used to record steps in the formation of diagenetic patterns. Present-day foraminifera species exhibit a dissolution ranking that is strongly controlled by the thickness, size and shape of the tests. Most coccoliths are more resistant than foraminifera with regard to dissolution. First evidence of coccolith dissolution occurs in depths above 3000 m, but changes in the composition of the assemblages become evident in depths between 3000 and about 5000 m.
2.4.5.7 Carbonate Slopes, Periplatform Carbonates and Carbonate Aprons Gravity flows, particularly turbidites and debris flow deposits, are common on carbonate slopes, and also occur in basins (Cook et al. 1983; McIlreaü992). These sediments are derived from shallow-water platform and platform-edge settings, sometimes also from the upper slope, and transported to upper, mid and lower slope settings. Carbonate slopes were categorized as: (1) ‘by-pass slopes’ on which sediment is transported from shallower to deeper water without significant deposition on the slope: gravity flows bypass the slopes nd accumulate in adjacent basins. (2) less steep ‘depositional or accretionary slopes’ where gravity flow sediments are predominantly deposited in lower slope settings, as dem-
Basics: Deep-Marine Carbonates
onstrated by excellent examples from the Holocene Bahamian slopes. The resulting fine-grained ‘periplatform carbonates’ (Schlager and James 1978; Boardman and Neumann 1984; Boardman et al. 1986; Pilskaln et al. 1989) form a sequence of resedimented shallow-marine carbonates intercalated within pelagic carbonates. The Bahamian carbonate slopes (Straits of Florida, Little Bahama Bank) offer the possibility of studying the interplay between allochthonous and pelagic sedimentation, and submarine cementation, resulting in the formation of hardgrounds. The periplatform ooze of the upper slope consists of a mixture of bank-derived material and pelagic debris. Lower slope and base-of-slope seiments are turbidites and mud- or grain-suported debris flow deposits, interbedded with periplatform ooze, and forming large carbonate aprons. Cyclic variations in the prevailing mineralogy (aragonitic and Mg-calcitic platform-derived grains vs calcitic pelagic grains) of periplatform carbonates and differences in the abundance of platform-derived and pelagic carbonates signalize climate changes, sea-level fluctuations and changes in the rate of carbonate production on the platforms which are reflected by microfacies analysis (Sect. 16.1). Another promising approach of microfacies studies is the evaluation of the controls on the different declivities of carbonate slopes which depends greatly on the primary depositional texture. Basics: Deep-marine carbonates Berger, W.H. (1991): Produktivität des Ozeans aus geologischer Sicht: Denkmodelle und Beispiele. – Zeitschrift der deutschen geologischen Gesellschaft, 142, 149-178 Cook, E., Enos, P. (eds., 1977): Deep water carbonate environments. – Soc. Econ. Paleont. Min. Spec. Publ., 25, 336 pp. Cook, H.E., Hine, A.C., Mullins, H.T. (1983): Platform margin and deep water carbonates. – Soc. Econ. Paleont. Min. Short Course, 12, 563 pp. Cook, H.E., Mullins, H.T. (1983): Basin margin environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H.(eds.): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 539-617 Coniglio, M., Dix, G.R. (1992): Carbonates slopes. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 349-373, Ottawa (Geol. Ass. Canada) Crevello, P.D., Harris, P.M. (1985): Deep water carbonates: buildups, turbidites, debris flows and chalks. – Soc. Econ. Paleont. Min. Core Workshop, 6, 527 pp. Doyle, L.J., Pilkey, O.H. (eds., 1979): Geology of continental slopes. – Soc. Econ. Paleont. Min. Spec. Publ., 27, 374 pp. Enos, P., Moore, C.H. (1983): Fore-reef slope environment. – In: Scholle, P.A., Bebout, D.G., Moore,C.H. (eds.): Carbonate depositional environments. – Amer. Ass. Petrol. Geol. Mem., 33, 507-537 Hsü, K.J., Jenkyns, H.C. (eds., 1974): Pelagic sediments: on land and under the sea. – Spec. Publ. Int. Ass. Sedimentol., 1, 447 pp.
Seep and Vent Carbonates
Leggett, J.K. (1985): Deep-sea pelagic sediments and palaeooceanography: a review of recent progress. – In: Brenchley, P., Williams, B.P.J. (eds.): Sedimentology, recent developments and applied aspects. – 95-118, Oxford (Blackwell) Little C.T.S., Campbell, K.A., Herrington, R.J. (2002): Why did ancient chemosynthetic seep and vent assemblages occur in shallower water than they do today? Comment. – Int. J. Earth Sciences, 91, 149-153 Macilreath, I.A., James, N.P. (1979): Facies models. Carbonate slopes. – Geosci. Canada Reprint Ser., 5, 189-199 Milliman, J.D., Droxler, A.W. (1996): Neritic and pelagic carbonate sedimentation in the marine environment: ignorance is not bliss. – Geol. Rundschau, 85, 496-504 Mullins, H.T. (1986): Periplatform carbonates. – Colorado School Mines Quart., 81, 63 pp. Mullins, H.T., Cook, H.E. (1986): Carbonate apron models: alternatives to the submarine fan model for paleoenvironmental analysis and hydrocarbon exploration. – Sed. Geol., 48, 37-79 Oxburgh, R., Broecker, W.S. (1993): Pacific carbonate dissolution revisited. – Palaeogeogr., Palaeoclimat., Palaeoecol., 103, 31-39 Schlager, W., Chermak, A. (1979): Sediment facies of platform-basin transition, Tongue of the Ocean, Bahamas. – Soc. Econ. Paleont. Min. Spec. Publ., 27, 193-208 Schlager, W., Camber, O. (1986): Submarine slope angles, drowning unconformities, an self-erosion of limestone escarpments. – Geology, 14, 762-765 Schlanger, S.O., Douglas, R.G. (1974): Pelagic ooze-chalklimestone transition and its implications for marine stratigraphy. – Spec. Publ. Int. Ass. Sedimentol., 1, 117-148 Shanmugam, G. (2000): 50 years of the turbidite paradigm (1950s - 1990s): deep-water processes and facies models – a critical perspective. – Marine Petroleum Geol., 17, 285342 Stow, D.A.V., Faugéres, J.C. (eds., 1998): Contourites, turbidites and process interaction. – Sed. Geol., 115, 1-386 Westphal, H., Reijmer, J.J.G., Head, M.J. (1999): Sedimentary input and diagenesis on a carbonate slope (Bahamas): response to morphologic evolution of the carbonate platform and sea-level fluctuations. – In: Harris, P.M., Saller, A.H., Simo, J.A.T. (eds.): Advances in carbonate sequence stratigraphy: application to reservoirs, outcrops, and models.– Soc. Econ. Paleont. Min. Spec. Publ., 63, 247-274 Further reading: K022, K023, K024
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Kennish 1993). The very first vent site with sea-bed dwelling large benthic organisms was discovered 1977 on the Galapagos rift in the eastern Pacific, where hydrothermal vent organisms are associated with sulfide mineralizations and surround black and white smokers within an area of a few hundreds of square meters. Today, examples of seep and vent communities are known from the Pacific (Fig. 2.12), the Monterey Bay of California, the Gulf of Mexico, the continental slope off Louisiana, the North Sea and the Black Sea (Nybakken 1993). Modern cold-seep communities are nearly always found in oligotroph waters of the aphotic zone, and occur from shallow to deep water down to about 3000 m (Callender and Powell 1999, Little et al. 2002). The success of the low-diversity vent communities (consisting of bivalves, vestimentiferan tube-worms, arthropods) depends on specialized anaerobic bacteria, which build carbon molecules using reduced inorganic compounds (particularly sulfides) that are contained in the heated water gushing from the vents. The low-diversity vent communities depend on these endosymbiontic bacteria. Bivalves (giant vesicomyid clams, infaunal lucinids) are able to exploit directly gas seepage by hosting chemoautotrophic bacteria in symbiosis in the gills. The symbionts produce organic matter via chemosynthesis using methane or sulfur from vented fluids or available in the environment. Hydrothermal vent communities and cold seep communities are phylogenetically related and show similarities in the physiology and anatomy of their dominant organisms.
2.4.6 Seep and Vent Carbonates
Cold-seep communities contribute to the formation of a specific carbonate rock type. Characteristics of these limestones are (1) carbonate beds embedded in shales or mounds exhibiting irregular textures, and occurring within a strongly differing facies, (2) low-diversity but abundant mollusks, and (3) abundant authigenic carbonate cements. Ancient vent and seep carbonates are becoming increasingly more frequently reported from the Phanerozoic rock record (see Sect. 16.5).
One of the most exciting discoveries in the last years is the recognition of authigenic and biogenic carbonates formed in deep-water settings by fluid and gas expulsions near cold seeps and hydrothermal vents (Beauchamp and von Bitter 1992). Biomass concentrations of organisms with calcareous shells occur near hot hydrothermal vents on midocean ridges and cluster around cold seeps of natural gas, brines, hydrocarbon and methane in varying depths of the oceans, particularly within faulted continental margin environments (Tunnicliffe 1992; Lutz and
Basics: Seep and vent carbonates Aharon, P. (2000): Microbial processes and products fueled by hydrocarbons at submarine seeps. – In: Riding, R.E., Awramik, S.M. (eds.): Microbial sediments. – 270-282, Berlin (Springer) Beauchamp, B., von Bitter, P. (eds., 1992): Chemosynthesis: geological processes and products. – Palaios, 7, 337-484 (with papers on modern and ancient chemosynthetic carbonates) Callender, W.R., Powell, E.N. (1997): Autochthonous death assemblages from chemosynthetic communities at petroleum seeps: biomass, energy flow and implications for the fossil record. – Hist. Biology, 12, 165-198
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Life on Seeps and Vents
Fig. 2.12. – A: Actinians, sponges and gorgonians settling on a non-active black smocker. Yellowish color of substrate is caused by iron-sulfur compounds. East-Pacific ridge, 2300 m water depth. B: Abundant white deep-sea mussels (Vesicomyidae) and Neolepas sp. together with shrimps (Alvinocaris) and gastropods (Phymorhyncus) living on a H2S rich crater floor with a slight content of methane. Hydrothermal activity changed the sediment which now contains amorphic iron-monoxyd and pyrite. Edison sea-mount (Bismarck Sea, Papua-New Guinea), 1450 m water depth. C: A dying community of deep-sea mussles (Vesicomyidae) in the crater of the Edison sea-mount. Many carrion-feeding gastropods (Phymothyncus) thrive on the distinct double valves. Same locality as 2. D: Fresh pillow lava with some echinoderms (Brisingidae). East-Pacific ridge, 2400 m depth. All photographs courtesy of Forschungsinstitut Senckenberg, Frankfurt/Main.
Callender, R., Powell, E.N. (2000): Long-term history of chemoautotrophic clam-dominated faunas of petroleum seeps in the northwestern Gulf of Mexico. – Facies, 43, 177-204 Fisher, C.R. (1990): Chemoautotrophic and metanotrophic symbioses in marine invertebrates. – Reviews of Aquatic Sciences, 2, 399-436 Hovland, M. (1990): Do carbonate reefs form due to fluid seepage? – Terra Nova, 2, 8-18 Little, C.T.S., Campbell, K.A., Herrington, R.J. (2002): Why did ancient chemosynthetic seep and vent assemblages occur in shallower water than they do today? Comment. – Int. J. Earth Sciences, 91, 149-153 Lutz, R.A., Kennish, M.J. (1993): Ecology of deep-sea hydrothermal vent communities: a review. – Reviews in Geophysics, 31, 211-242 Paull, C.K., Chanton, J.P., Neumann, A.C., Coston, J.A., Martens, C.S. (1992): Indicators of Methane-derived carbonates and chemosynthetics organic carbon deposits: examples from the Florida escarpment. – Palaios, 7, 361375
Peckmann, J., Reimer, A., Luth, U., Hansen, B.T., Heinicke, C., Hoefs, J., Reitner, J. (2001): Methane-derived carbonates and authigenic pyrite from the northwestern Black Sea. – Marine Geology, 177, 129-150 Roberts, H.H., Aharon, P., Walsh, M.M. (1993): Cold-seep carbonates of the Louisiana continental slope-to-basin floor. – In: Rezak, R., Lavoie, D.L. (eds.): Carbonate microfabrics. – 95-104, New York (Springer) Stakes, D.S., Orange, D., Paduan, J.B., Salamy, K.A., Maher, N. (1999): Cold-seeps and authigenic carbonate formation in Monterey Bay, California. – Marine Geology, 159, 93-109 Tunnicliffe, V. (1991): The biology of hydrothermal vents. Ecology and evolution. – Oceanogr. mar. Biol. Ann. Rev., 29, 319-407 Tunnicliffe, V. (1992): The nature and origin of the modern hydrothermal vent fauna. – Palaios, 7, 338-350 Venturini, S., Selmo, E., Tarlao, A., Tunis, G. (1998): Fossiliferous metanogenetic limestones in Eocene flysch of Istria (Croatia). – Giornale di Geologia, 60, 219-234 Further reading: K211
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MICROFACIES ANALYSIS This book consists of three major parts. This first part ‘Microfacies Analysis’‚ deals with the methods of microfacies studies (Chap. 3), qualitative and quantitative microfacies data (Chap. 4, 5 and 6), diagenetic criteria (Chap. 7), limestone classification (Chap. 8) and biological controls on the formation of carbonate rocks (Chap. 9 and 10).
3 Methods
This chapter summarizes field work studies, sampling strategies and laboratoy methods relevant to microfacies analysis. A precise field record of geological and paleontological data as well as sampling strategies that consider the vertical and lateral variations in these data are vital for the success of microfacies studies. Laboratory methods study microscopic features and mineralogical and geochemical data.
A highly recommended description of field methods is given by Tucker (1982). The methods used in the study of litho- and biofacies of sedimentary rocks are the subject of several textbooks (Leeder 1982; Lewis 1984; Pettijohn et al. 1987; Miall 1990; Chamley 1990; Blatt 1992, Friedman et al. 1992, Lewis and McConchie 1993; Boggs 1995; Reading 1996; Leeder 1999; Miall 2000; Einsele 2000).
3.1 Field Work and Sampling
3.1.1.1 Lithology, Texture and Rock Colors
Many criteria that are important in facies interpretation can not be recognized sufficiently in small-scale microfacies. Paleoenvironmental interpretations derived from microfacies should be controlled by lithological criteria and sedimentary structures evaluated by the high information potential provided by fossils and biogenic structures. Microfacies sampling requires an understanding of the meaning of bedding and the depositional characters reflected by sedimentary structures.
Lithology: Sedimentary rocks are classified into siliciclastic rocks (claystone/mudstone, siltstone and sandstone), conglomerates and breccias, carbonate rocks (limestone, dolomite), mixed siliciclastic-carbonates (marl, argillaceous and sandy limestones), evaporites (gypsum, anhydrite, salt), siliceous sedimentary rocks (cherts), phosphorites and organic-rich rocks. Carbonate rocks make up 20 to 25 percent of all sedimentary rocks in the geological record and are classified into limestones and dolomites (dolostones); see Fig. 3.1. Limestones consist of more than 50% CaCO3. They comprise limestones and dolomitic limestones. Dolomites are composed of more than 50% CaMg(CO3)2 and are subdivided into calcitic dolomites (50–90% dolomite) and dolomites. Although detailed differentiation of carbonate rocks is best performed in the laboratory, field distinctions can also be made, using dilute 10% hydrochloric acid (limestone will fizz strongly, dolomite will show little or no reaction). Alizarin-red-S
3.1.1 Field Observations Field work forms the basis for studying sedimentary deposits. Criteria that can be studied in outcrops and cores in facies analyses are: lithology, texture, rock color, bedding, sedimentary structures and diagenetic features, and fossils and biogenic structures.
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_3, © Springer-Verlag Berlin Heidelberg 2010
54
Fig. 3.1. Common lithologies of carbonate rocks and mixed carbonate-siliciclastic rocks. The symbols used are those from the standard Shell Legend.
in weak HCl stains limestone pink to mauve but leaves dolomite unstained. Mixed siliciclastic-carbonate lithologies include marls (a sediment consisting of 65–35% carbonate and 35–65% clay), argillaceous limestones (a limestone containing an appreciable amount of clay, less than 50%) and sandy limestones (a limestone containing an appreciable amount of quartz sand). Texture: Attention should be focused on the depositional variability of limestone beds, on depositional processes recorded by characteristic sedimentary patterns, and on the control of microfacies-based paleoenvironmental interpretations using larger-scaled criteria. The texture of major limestone types can be characterized in the field using the classification proposed by Dunham (1962); see Sect. 8.3.2. Much information is provided by the size, shape and sorting of grains, the ratio of matrix to grains, and the fabric. Details are described in Chap. 4. The textural composition of beds and laminae may be characterized by distinct vertical gradations in grain size (graded bedding). ‘Normal grading’ (coarser grains at the base, finer grains at the top) and ‘reverse or inverse grading’ (coarser particles at the top grading downwards to finer particles) are important in the interpretation of depositional processes (Pl. 17/1, Pl. 102/2, Pl. 115/2, Pl. 116/ 1, Pl. 126/2). Rock color: Many limestones are a shade of gray, but limestones can also display distinctive rock colors which need a more precisive terminology (Folk 1969). Dolomites are often creamy yellow or brown in color. Rock colors can be characterized by simple verbal descriptions (e.g. medium to dark gray) or by comparing fresh rock fracture surfaces with standard color schemes
Rock Color
(e.g. U.S. National Research Council Rock-Color Chart based on the Munsell Color System: Goddard et al. 1948). The use of the Rock Color Chart requires (a) determination of the gradation of color (hue), followed by (b) estimation of light to dark values, and (c) estimation of the degree of saturation (chroma). Colors are indicated by a code (e.g. 5 R 4/6) or by a descriptive name (e.g. moderate red). Color names used in the Rock Color Chart do not always correspond to the subjective color impression of geologists, which may be adequate for general surveys but not, for example for the characterization of carbonate rocks used in the arts or as building stones. Colorimetric analysis is another approach but relatively rarely used for limestones (Engelbrecht 1992), except when determining technological criteria. Colors of carbonate rocks are strongly controlled by depositional conditions, diagenesis and recent weathering. The amount and composition of the non-carbonate fraction is crucial for the color aspect because the adsorption of iron and manganese on the surface of finely dispersed clay minerals is often responsible for distinct colors. Fe and Mn content, rock color and depositional facies may exhibit distinct correlations, specifically in hemipelagic and pelagic sediments. Of particular significance is the oxidation-reduction balance: Yellow colors, for example, can be caused by the oxidation of FeS as demonstrated by the famous Jurassic Solnhofen Limestone of Southern Germany which appears in fresh samples bluish gray. Comparisons of photometric color measurements with geochemical data of limestones indicate correlations of bluish and yellowish colors with pyrite and goethite contents and dark or light rock color values with clay and silt contents. Green and drab gray colors may be caused by organically derived carbon and Fe2+ compounds such as sulfides in reduced sediments. Various shades of red, yellow and brown colors are due to Fe3+ compounds such as hematite and limonite in oxidized sediments. Yellowish shades of marbles can be caused by small quantities of siderite. Dark colors are also induced by aromatic hydrocarbons (pure solvent-soluble organic matter composed of carbon and hydrogen only). The amount of organic matter and the degree of light reflection are related (Patnode 1941). Black colors can also be caused by thermal heating. Sedimentary limestones may undergo considerable color change during metamorphism (Winkler 1994). Why is this limestone red? Vivid rock colors generally provoke questions which are difficult to answer.
Bedding
55
Very generally, the coloration of sediments is regarded as synsedimentary (caused by the input of detrital coloring material or by contributions of ferric bacteria; Mamet et al. 1997) or postsedimentary (caused by diagenetic changes within the sediment or coloration by overlying beds). Red limestones may serve as an example for the problems involved: The red color of marine carbonates is primarily caused by small amounts of Fe3+ (about 2%: Franke and Paul 1980). Volcanic emanations provide a direct source of iron in seawater. But the weathering of land masses associated with an input of iron to the sea (as colloids or absorbed on clay minerals) and organic matter provide the major source. Low sedimentation rates, near-bottom waters rich in oxygen, the absence of sulfate-reducing bacteria as well as the local Eh/pH potential within the sediment and at the sediment/water interface favor the development of red colors (e.g. by dehydration of goethite to hematite: Rech-Frollo 1971).The intensity of the red coloration may depend on the amount of detrital material and iron (Hallam 1967; Flügel and Tietz 1971). Because iron is transported from land to sea via several solution/precipitation-phases, the importance of climatic controls for the formation of red limestones must also be considered. The following diagenetic factors promote rock coloration: (a) Compaction facilitates the flow of oxydized pore water through the sediment. (b) Pressure solution and freeing of iron from clay minerals may be responsible for red-colored stylolite seams, often occurring together with burial dolomites (Mattes and Mountjoy 1980). (c) During dedolomitization, Fe originally incorporated in dolomite can be precipitated as FeO(OH) and oxidized to Fe2O3, producing reddish rock colors. cm
Shell Standard Legend 1995 Meter bedded
100 30
Colored paleosols (e.g. terra rossa, a residual soil, over limestone bedrock, typically in karst areas) can impact on rock colors. Reddening of paleosol horizons is caused by (a) in situ alterations of iron-rich clays, (b) reddened detritus brought from the exposed surface downward through the soil by percolating waters, (c) dehydration of goethite to limonite and hematite during subaerial exposure in a well-drained environment (Retallack 1981), or (d) concentration of reddish matrix by eluviation within the paleosol profile during alternating wet and dry climate periods (Goldhammer and Elmore 1984). Terra rossa, brought to the sea will not directly cause red colors (Hinze and Meischner 1968). Near-surface cavities in karstified reef limestones may contain red internal sediments, derived from an overlying soil zone, with deeper cavities being filled with green and gray internal sediments (Satterley et al. 1994). Changes of rock colors with depth below the surface indicate changes in the oxidation state within the vadose and phreatic zones (Esteban 1991).
3.1.1.2 Bedding and Stratification, Sedimentary Structures and Diagenetic Features Many of the structures described in this subchapter can also be recognized at a microfacies scale, particularly sedimentary fabrics and diagenetic textures (Fig. 3.4). Bedding and stratification: Carbonate rocks are stratified or non-stratified. Non-stratification is a result of (a) primary lack of bedding (e.g. in reefs), (b) bed-destructing processes (e.g. intensive burrowing), or (c) diagenetic processes (e.g. dolomitization or strong recrystallization of limestones).
Boggs 1995 Very thickly bedded Thickly bedded
Layers greater than 100 mm, thick to very thick beds
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10
Thinly bedded 3
Demicco and Hardie 1994
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Fig. 3.2. Common field classifications of stratified sediments. The Shell classification is based on scales. Many textbooks on sedimentology, e.g. Boggs (1995), recommend a classification based on thickness ranges. Demicco and Hardie use only three thickness categories, whose limits were determined by the experience of the authors studying predominantly Upper Proterozoic and Lower Paleozoic carbonates. In choosing 5 mm as the upper boundary of the thinnest category, the importance of thin microbial mats in tidal flats has been considered. Bed thicknesses between 0.5 and 10 cm and between 10 and 30 cm are common in shelf carbonates.
56
Thickness of Beds
bounded above and below by distinct surfaces are called bedsets (Collinson and Thompson 1982). Limestones consisting of thin, strictly parallel or non-parallel laminae are named laminites (Sect. 5.1.3). The descriptive term ‘Plattenkalk’ refers to sequences composed of fine-grained, laminated limestones, which are bedded in flat centimeter-decimeter thick, parallelsided units (Swinburne and Hemleben 1994).
Fig. 3.3. Geometries of carbonate rock layers. Thin sections. Left – Planar fine-laminated limestone characterized by alternating millimeter-scale lamina couplets consisting of mudstone (dark gray) and packstone (light streaks). The black line in the center and the black spots near the top are bitumina. Note the extremely fine lamination within the lamina couplets, the parallel orientation of fine shells and the concave-up position of the ostracod shell near the top of the sample. These criteria indicate suspension settle-out of mud (calcisiltite) within a quiet-water environment, alternating with sporadic influx of peloids, clay and terrigenous quartz grains transported by gentle traction currents. The sample represents an organic-rich facies formed in small isolated, tectonically induced basins within an intertidal to subtidal carbonate platform. The bitumen-rich limestones were deposited from a stratified water body with anaerobic and dysaerobic conditions (Czurda 1973; Hopf et al. 2001). Large-scale cyclicity was controlled by sea-level change, small-scale cyclicity was controlled by climatic and seasonal fluctuations. These limestones are regarded as potential carbonatic source rocks. Right – Wavy-crinkled fine laminated dolomite characterized by irregularly undulated stromatolitic layers. Black laminae are interpreted as the result of microbial activity (binding of sediment, accumulation of organic material). The sample represents the intertidal part of the Hauptdolomit platform. Both samples are from the Late Triassic Hauptdolomit of the Wiestal quarry, Gaissau, Salzburg/Austria. Scale is 5 mm.
Stratification : A bed is a tabular or lenticular layer that is sufficiently distinct from under- and overlaying strata with respect to lithology, composition and texture and separated from these strata by recognizable boundary planes. Genetically, a bed may represent a depositional episode during which physical conditions did not change significantly (‘sedimentation unit’: Otto 1938). Following the proposal of McKee and Weir (1953) the term bed should be limited to strata thicker than 1 cm. Very thin strata (thickness <1 cm) are called laminae (Campbell 1967). Laminae are the result of changing depositional conditions causing variations in the grain size, texture, mineral composition, and content of clay and organic material. Groups of laminae forming a distinct sedimentary structure are called laminasets. Groups of similar beds or cross-beds,
Field studies of bedding and stratification must consider (1) boundary planes and bedding surfaces, (2) bed thickness, (3) composition and internal structure of beds, and (4) vertical bed sequences. (1) Boundary planes and bedding surfaces: Upper and lower surfaces of beds are called bedding planes, most of which are caused by abrupt changes in depositional conditions, non-deposition or erosion. However, bedding planes also result from diagenetic processes or weathering. Pressure solution and stylolite seams parallel to depositional bedding may produce ‘pseudo-bedding‘ simulating true depositional bedding, and leading to sampling problems, because bed-to-bed sampling usually relies on the assumption of sedimentary boundary planes. Bedding surfaces are abrupt planar, abrupt irregular or gradational, with even (flat), wavy or curved surfaces that are parallel or non-parallel to each other. Differences in bedding surfaces provide information related to the time span of non-deposition (e.g. scalloped or irregular potholed surfaces, sometimes with paleosols, indicate a paleokarst stage; irregularly shaped hardgrounds may point to deep-marine solution). Sedimentary structures on top and bottom surfaces of bedding planes provide information about depositional energy and substrate features. Some of the more common structures on the top surfaces of limestones are modifications of surface topography, bed forms (ripples, dunes, hummocks), shrinkage cracks and primary current lineations, and various trace fossils. Structures seen on bottom surfaces include flute marks, groove marks, tool marks, scour structures as well as load marks. (2) Bed thickness: Terminologies for the bed and laminae thickness have been proposed by various authors (McKee and Weir 1953; Ingram 1954; Reineck and Singh 1980), but none has been generally accepted. Ingram’s scale (used by many authors) and the Shell approach (distinguishing only ‘ranges’) are to be recommended (Fig. 3.2). Demicco and Hardie (1994) distinguish three layer categories based on layer thickness, layer boundary geometry, and textural layer composition. This classification is a practical key to the identi-
Field Criteria
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Stratification Planar lamination Wavy lamination Crinkled lamination Planar thin bedding Wavy thin bedding Crinkled thin bedding Planar thick bedding Wavy-crinkled thick bedding Hummocky cross-bedding Large-scale cross-bedding Wave-ripple cross-bedding Heterolithic cross-bedding Small-scale cross-bedding Graded bedding Cracks Mud cracks Prism cracks Sheet cracks Fenestrae Breccias - Conglomerates Storm-lag lithoclast gravels Channel-lag gravels Edgewise gravels Imbricate gravels Limestone breccias Submarine hardgrounds Neptunian dykes Deformation structures Nodular structures Biogenic structures Microbial laminites Stromatolites Thrombolites Bioturbation Diagenetic structures Caliche Tepee Solution pores Evaporite minerals Chicken-wire sulfates Deformed sulfate layers
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Fig. 3.4. Common field criteria of carbonate rocks include stratification types as well as sedimentary and diagenetic structures. The processes causing many of these structures are highlighted by microfacies data (black dots). The figure considers the scheme developed by Demicco and Hardie (1994) for shallow-marine carbonates.
fication and interpretation of layering without crossstratification, and can be applied both to field and thinsection observations (see Fig. 3.3 and Pl. 18). Bed thickness analysis of limestones can be used in recognizing breaks in sedimentation, cyclic sedimentation patterns and gradual environmental changes. These investigations are based on the hypothesis that the thicknesses of stratigraphic units follows a lognormal distribution (Lumsden 1979). (3) Composition and internal structure: Beds appearing internally homogeneous (massive, structureless) in the field, may exhibit textural and structural inhomogenities when studied using staining and etching methods, ultra-thin sections, X-ray or tomography (Larsen and Theide 1971; Kenter 1989).
Many beds are internally differentiated into laminae and a number of informal units. Laminae result from short changes in the depositional conditions causing variations in grain size, composition and texture, fossil content as well as the content of non-carbonate and organic material. Small-scale laminations of fine-grained limestones assist in recognizing of yearly varves, distal turbidites or tidal rhythmites (Brown et al. 1990; Archer and Feldman 1994). Fine-scale biogenic lamination is particularly common in lagoonal and lacustrine sediments; stratification types of lake sediments were classified by Cole and Picard (1975) and Bullwinkel and Riegel (2001). Internal layers and laminae that are essentially parallel to the bedding planes constitute a planar stratification or planar lamination (Fig. 3.3). A practical key
58
for identifying and interpreting layering in carbonates without internal cross-stratification was provided by Demicco and Hardie (1994). The classification is based on layer thickness (Fig. 3.2, layer geometry (planar: Pl. 20/7; wavy: Pl. 18/2, 4, Pl. 123/1; and crinkled: Pl. 25/5, Pl. 123/3) and layer composition (using Dunham’s limestone classification). The authors present a very useful discussion of the possible origins of laminated and bedded carbonates. Layer geometry and layer composition can be evaluated by microfacies data. Layers or laminae deposited at an angle to the bounding surfaces are responsible for the formation of crosslamination. All types of cross-stratification as well as flaser and lenticular bedding known from siliciclastic rocks also occur in carbonate sediments: Small- and large-scale cross-stratification, wave-ripple cross-stratification, hummocky cross-stratification (Imbrie and Buchanan 1965; Duke 1985). Planar lamination and cross-lamination are common in shallow-marine carbonates, but cross-lamination is often obliterated. Parallel laminae form in a variety of environments (lacustrine; peritidal to deep sea) but are often destroyed by burrowing organisms and therefore better preserved in reducing or anaerobic environments. Heterolithic stratification is characterized by irregular alternations of mud and sand layers (Pl. 18/2) which result in flaser, wavy or lenticular bedding structures. (4) Vertical bed sequences can be recognized from the differences in texture, fossil content or be thickness. Thinning/fining-upward and coarsening/thickening-upward sequences related to rates of environmental change and depositional energy are the significant patterns recording changes in sea level and accomodation. Sedimentary structures and diagenetic features: These structures exhibit a wide variety of pre-, syn- and postdepositional features which allow the processes and the interpretation of paleoenvironments to be dismissed (Fig. 3.5). A highly recommended introduction to sedimentary structures and early diagenetic features of shallow-marine carbonates is the well-illustrated SEPM Atlas prepared by Demicco and Hardie (1994). Many sedimentary structures are identical with those of siliciclastic rocks (e.g. cross-bedding, ripple marks) but, because of the strong impact of biogenic and diagenetic factors on carbonate sedimentation, structures also exist which are more common in or limited to limestones (e.g. fenestral and growth cavities, tepees). Excellent overviews of sedimentary structures of non-carbonate and carbonate rocks were provided by Potter and Pettijohn (1963), Collinson and Thompson (1988), Reineck and Singh (1986), and Ricchi Lucchi
Fossils and Biogenic Structures
(1995). Allen (1984) discussed the processes controlling specific sedimentary structures. Various classifications of sedimentary and diagenetic structures differ in their more genetic and process-oriented or descriptive approaches. Demicco and Hardie (1994) differentiate physical structures, biogenic structures, and chemical features. Physical structures include stratification, breccias and conglomerates, mudcracks and other disruption structures, soft deformation structures, mechanical compaction structures and miscellaneous physical structures. Erosional structures, slump structures, diagenetic deformational structures and dikes are larger scale structures, commonly involving several beds. Slumping and sliding structures need specific studies. Instructive field and sample illustrations of sedimentary and diagenetic structures can be found in Pettijohn and Potter (1964) and Demicco and Hardie (1994). STRATIFICATION Lamination and bedding planar
wavy
crinkled hummocky
CROSS-BEDDING GRADED BEDDING
Fining upward (normal) Coarsening upward (inverse) Mudcracks
CRACKS FENESTRAE
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BRECCIAS, CONGLOMERATES Breccia
Conglomerate
NEPTUNIAN DYKES DEFORMATION STRUCTURE S load casts BIOGENIC STRUCTURES Bioturbation DIAGENETIC STRUCTURES Gypsum
Tepee
slumping convolute bedding Stromatolites Bored surface Solution pores
Anhydrite
Fig. 3.5. Standard Shell Legend symbols of abundant sedimentary, biogenic and diagenetic structures occurring in carbonate rocks. The symbols are useful for a uniform schematic representation.
Checklist for Field Description
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Box 3.1. Checklist for field description of paleontological data in carbonate rocks. Bedded limestones Bedding and sedimentary structures Bedding types (lithology, thickness, bedding boundaries, texture; see Fig. 3.1, Fig. 3.2) ‘Event beds’ ? (beds caused by physical and biological events of short duration, e.g. storms, turbidity currents, earthquakes, sudden death of organisms resulting in mass acummulations; see Sect. 5.1.3) Cyclic patterns ? (see Sect. 16.1) Sedimentary structures and diagenetic features (see Fig. 3.4) Types and preservation of fossils Classification of fossils (group, taxa) Evaluation of fossil types with regard to ecological criteria Trace fossils and/or body fossils ? Burrowing and bioturbation (see Sect. 5.1.4) Taphonomic criteria (see Chap. 10) Abundance and distribution of fossils Fossil concentrations Abundance of fossils based on bulk samples or on counts using various grids (see Sect. 6.2.1) Distributional patterns of fossils on bedding planes (random, patchy, oriented) Lateral distribution of fossils within the beds Distribution and type of fossils within underlying and overlying beds Massive limestones (e.g. reef limestones) Geometry of the carbonate body Mapping of the boundary and geometry of the buildup Mapping of major litho- and biofacies types Establishing time lines (tracing bedding surfaces in the flank facies), discontinuities and event marker beds, e.g. storm layers, synsedimentary breccias) Look for bottom-top structures (particulary at the flanks of buildups) Types and preservation of fossils Classification of fossils Taphonomic criteria Analysis of the principal fossils with regard to their inferred role in buildup formation (e.g. framebuilders, binding and encrusting organisms) Differentiation of growth forms of sessile organisms Abundance and distribution of fossils Vertical and horizontal cover (density) and size of reefbuilding organisms (using a gridded overlay, the quadrat method orthe line method, see Sect. 6.2.1) Recognition of significant vertical sequential changes in growth forms, taxonomic composition or abundance Relation between sediment, texture and fossils Description of the ‘matrix’ between and around autochthonous reefbuilders (see Sect. 16.2.5.2) Description of growth cavities, shelter structures, open-space structures and solution voids (see Sect. 5.1.5) with regard to geometry, dimension, distribution and filling (sediment, spar, crystal silt, still open) Description of fissures with regard to setting, size, orientation, filling and fossils Abundance of fine-grained sediment and carbonate cements in relation to benthic fossils Diagenetic features: see Sect. 7.45
3.1.1.3 Fossils and Biogenic Structures Fossils, fossiliferous sediments and biogenic sedimentary structures are key sources of information on paleoenvironments, sedimentation patterns and diagenesis. Strategies necessary for describing paleontological data in the field were summarized by Goldring (1991). Field strategies are different for bedded sediments and for dominantly autochthonous buildups. Box 3.1 is based on a field classification of fossiliferous sediments proposed by Goldring (1991). Of special interest are ‘fossil concentrations’, bioclastic fabrics, and ichnofabrics.
Fossil concentrations are relatively dense accumulation of biologic hardparts (irrespective of taxonomic composition, state of preservation, or degree of postmortem modification). By taking into consideration the taxonomic composition, bioclastic fabric, geometry and internal structure of the deposit, skeletal concentrations can be grouped genetically according to the inferred relative importance of biogenic, sedimentologic and diagenetic processes (Kidwell et al. 1986; Fig. 3.6). This approach is useful for interpreting ‘shell beds’ composed of mollusks and brachiopods, and understanding the processes
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Field Logs
Fig. 3.6. Field description of skeletal concentrations (Modified and completed after Kidwell et al. 1986). Accumulation of shells, echinoderm fragments and other fossils are produced by biogenic (area 1 in the triangle), sedimentologic (area 2) and diagenetic processes (area 3). An example of biogenic concentrations are oysters, whose larvae preferentially colonize sites with abundant adults causing reef-like structures in subtidal shelf settings. Photo at top right: Placunopsis, Middle Triassic, Germany. Sedimentological concentrations form through hydraulic reworking of hardparts, and/or through non-deposition of sediment, or selective removal of the sedimentary matrix. These processes result in the formation of conspicuous shell beds (lumachelles, coquinas). An example is shown in the foto at bottom left: Accumulation of looseley packed, bimodal sorted bivalve shells within a storm layer originated in inner shelf position; the inversely graded carbonate mud was infilled sub-sequent to the deposition of the shells. Late Triassic, Northern Alps, Austria. Diagenetic concentrations originate from the increase of the spatial density of fossils by compaction and selective pressure solution. The photo at bottom right displays a crinoidal limestone, formed on a carbonate ramp and strongly affected by pressure solution. Early Carboniferous: Frank’s Slide, Turtle Mountains, Alberta/Canada. The triangle can be used to estimate the position of shell beds along an onshore-offshore transect: Tidal flats are characterized by a composition corresponding to the areas 1 and 2; lagoons by the dominance of the areas 1 and 4; shallow shoals by the dominance of the area 2; inner shelf shell beds by the areas 2 and 4; and subtidal outer shelf settings below storm wave base by the abundance of concentrations that correspond to area 1.
and environments responsible for their origination. Fossiliferous carbonates yielding fossils or fossil fragments larger than 2 mm can be readily categorized into macroscopic fabric types using semi-quantitative scales for close-packing and size-sorting of coarse bioclasts (Kidwell and Holland 1991; Fig. 3.7). The fabric types are useful to narrow the possible modes of origin, e.g. accumulations of mollusk shells or crinoidal fragments. Ichnofabrics and trace fossils have a high potential for interpretating of paleoenvironments and depositional processes (Donovan 1994; Bromley 1996). A semiquantitative field classification of bioturbation based on pattern recognition of the percentage of original sedimentary fabric disrupted by biogenic reworking facilitates the explanation of bioturbation structures seen in thin sections (Droser and Bottjer 1986; see Sect. 5.1.4). Distinctive distributional patterns of trace fossils in carbonate shelf environments are useful in delineating environments of deposition and determining sea-level positions.
3.1.1.4 Field Logs and Compositional Logs Field information originating from stratigraphic sections and mapping is commonly recorded by written descriptions of observed features (field notes: Tucker 1982, Graham 1988). Graphic logs aid in documenting of rock composition and texture and consist generally of a main column showing the lithological sequence (drawn as column or imitating a weathering profile) and adjacent columns showing the occurrence and distribution of sedimentological and paleontological data. These columns may include only field data or a combination of field and laboratory data. Various standard forms for plotting graphic logs and standardized coding systems have been proposed to uniformly describe facies criteria, and to facilitate transforming field and laboratory data into digital data (e.g. Charollais and Davaud 1974; AmStrat-CanStrat system: Miall 1990). Reijers et al. (1993) recommend a hierarchically organized lithofacies coding system allowing macro-
Packing
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Fig. 3.7. Classification of coarse bioclastic fabrics characterized by closepacking of fossils (modified after Kidwell and Holland 1991). The classification can be used in the field as well as in thin section studies. Basic parameters are packing, size sorting of the bioclasts larger than 2 mm, and the biofabric which is bioclastsupported (‘densely packed’) or matrixsupported (‘loosely packed and ‘dispersed’). The fabrics are well-sorted exhibiting little variations in the size of bioclasts, bimodal sorted (with two discretes modes), or poorly sorted exhibiting great variations in bioclast sizes. The packing and sorting descriptors convey an immediate visual image of the fossiliferous limestone. The criteria have genetic implications. The categorization by close-packing and size-sorting limits the number of possible origins and leads to basic hypotheses that can be tested using taphonomic, palecologic, sedimentologic and stratigraphic features. The sample exhibits a fabric characterized by large gastropod shells and abundant smaller bioclasts consisting of small gastropods and bivalve valves. These bioclasts constitute an overall loosely packed fabric which would be classified as gastropod floatstone following Dunham’s classification (see 6.2). The finer bioclasts, however, are densely packed and these parts of the thin section would be classified as fine-bioclastic mollusk packstone. Small gastropods often are regarded as ‘dwarf’ forms pointing to unfavorable environments. It should be taken in mind, however, that the majority of gastropods is characterized by mm-scale shells. Note the strong recrystallization of the large gastropods. The originally aragonitic shell has been replaced by coarse-grained calcite crystals. The first whorls are completely obliterated; following whorls exhibit various cement generations or are infilled by micritic sediment. Late Cretaceous pebble in Quaternary deposits: Graz/Styria, Austria.
scopic features of sedimentary rocks to be objectively described and the data stored in a computer database. The standard lithologic logging system proposed by Shell International Exploration and Production (Standard Legend 1995) offers a widely used possibility for describing features recognized in outcrops, cores and microfacies samples. Lithology, texture and composition, rock color, accessory minerals, fossils, stratification and sedimentary structures, postdepositional criteria and diagenetic features as well as physical criteria (e.g. porosity and permeability) are expressed by symbols, standard abbreviations and codes. Too many symbols and criteria will mask trends in the stratigraphic sequence. For the purpose of microfacies studies the most important data should be shown in logs: e.g. depositional textures, fossils, grain types, matrix and carbonate cement. Logs are presented in a variety of scales depending on the problems involved. Detailed microfacies logs may require a scale of 1 cm = 1 m or less. Authors studying depositional patterns of carbonate rocks use a scale of 1 cm =10 m.
3.1.2 Sampling The following comments refer to outcrops and cores. Sampling of cuttings requires specific methods discussed in Sect. 17.1.4.2. Sampling depends on the nature of the project (Griffith 1967). The scope of the problem must be defined before collecting samples. Typical research objectives may be • depositional texture of carbonate rocks as an aid to classifying rock types and assessing physical properties, • type, texture and size of carbonate grains as an aid to determining depositional environment and paleocurrents, • studies of grains, matrix, cement and pores for understanding diagenesis and physical/technological properties of carbonate rocks, • paleontological criteria, for subdividing and correlating stratigraphical units or providing data which can help with paleoenvironmental reconstructions.
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3.1.2.1 Search Sampling and Statistical Sampling Size of samples Microfacies sampling should provide (1) lithological samples exhibiting textures, structures and fossils on a hand specimen scale, (2) samples for laboratory analysis (samples for geochemical investigations should consider induration and degree of weathering!), and (3) samples for paleontological studies. Because the methods used in laboratory investigations may not yet be certain at the time of sampling, large-sized samples are recommended allowing several medium to largescale thin sections to be prepared and specific questions to be studied. The sizes 5 x 4, 5 x 5, and larger up to 20 x 15 cm have proved useful as formats for thin sections used in microfacies analyses. Large and oriented samples (showing the position of the samples in spaces and marked in the field) are necessary for • determining top-and-bottom criteria, allowing the depositional and structural top to be recognized, • studying textural variations and fabrics (e.g. grain orientation; bedding, discontinuity surfaces, bioturbation, breccia), • describing biotic characters (e.g. growth forms, succession patterns, guilds). Reef limestones require large and very large samples because of the high compositional variations within centimeter and decimeter ranges. Only large samples allow synchronous and non-synchronous associations to be differentiated and sediment types and internal features (cavities, cryptic habitats) within bioherms or biostromes to be discussed. Large samples, sometimes comprising a whole bed, are also necessary for calciturbidites (Herbig and Mamet 1994). Equipment and first observations Microfacies sampling requires the use of a hammer, chisel and, sometimes specific equipment (angle grinder). Hand lenses as well as a pocket microscope for grain size measurements (Müller 1967) and textural observations in the field are important prerequisites. • Determine rock color, texture, composition and mean size of grains, induration and weathering in hand specimens. Statistical sampling Statistical sampling, based on random samples with the objective of estimating the statistical probability of facies types and patterns, is a common approach in modern microfacies studies (Carozzi 1988). Statistics require that the samples be taken according to a specific
How Many Samples?
plan, e.g. at equal distances or according to a random distribution. Common methods of sampling are: • Simple random sampling: Samples are taken at unevenly spaced intervals determined by random numbers, and the variability of the samples is checked by variance tests (Eltgen 1970). • Systematic sampling: Samples are taken at evenly spaced intervals and checked by significance tests; (McCammon 1975). Sampling within a constant vertical distance facilitates the statistical treatment of distributional data (Grötsch et al. 1994). • Stratified sampling: At least two samples are taken from each bed; the number of samples should be proportional to bed thickness. • Cluster sampling: Several adjacent samples are taken in random distribution. This approach was successfully used for microfacies studies of reef limestones and shelf carbonates. In practice, real possibilities of microfacies sampling and theoretical statistical requirements differ widely due to the outcrop situation, accessibility and physical condition of the student carrying heavy sample bags. ‘We collect what we can...’ Beware of the ‘family photo album effect’ In a family photo album extraordinary events like birthdays and holidays are commonly overrepresented as compared with photos of every day life. Similarly, many microfacies samples are often selectively taken because of a conspicuous abundance of fossils in some beds or exceptional textural types. The ‘daily life’ reflected by more monotonous lithotypes (e.g. micritic shelf limestones or hemipelagic limestones) is recorded to a minor degree, because the vertical sampling distance is larger and the spacing more irregular. The depositional ‘trends’ (e.g. cycles) derived from samples representing extraordinary events rather than ‘daily life’ may display an artificial pattern.
3.1.2.2 How Many Samples? The variability, diversity, and richness of microfacies data are related to sample size effects. An increase in sample size may lead to an increase of richness and diversity. Various analytical techniques are known to estimate sample size in geological, biological, and archeological studies. Some methods are also valuable in microfacies analysis (curves of sample size against diversity or richness, regression and simulation techniques). There is, of course, no general rule for the number of microfacies samples that are ‘truly representa-
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Advices for Sampling
tive’ of the limestone studied, but some empirical tendencies do exist: Shallow-marine bedded carbonate sequences need a detailed sampling program, because they exhibit high microfacies variability due to changing subtidal and tidal depositional sites and the common occurrence of cyclic sedimentation patterns. A very close sampling of cyclic carbonates at an interval of 0.10 m and less is recommended if high-order short-term cycles are to be detected (Buggisch et al. 1994). Thin- to medium-bedded platform and ramp carbonates are commonly sampled at intervals between 0.20 and 0.50 m (corresponding to a bed-to-bed sampling) or at intervals of several meters. Bed-to-bed sampling is advised if the study aims to understand of short-term changes in environmental controls of biota and sedimentation. Novak and Carozzi (1972) studied a 5 m thick section of Late Carboniferous-Permian platform carbonates by using thin section samples taken at intervals of 20 cm and a continuous series of samples. Although continuous sampling provided more information and allowed a better separation of microfacies types, the general trend in the sequence of microfacies types was still recognizable for sampling intervals of 20 cm. Wider intervals are used for (1) search sampling (taking samples preferentially from conspicuous beds noticeable because of a particular composition, structure or cyclic pattern), and (2) obtain pilot studies aiming to a first overall picture of the section studied and recognize lithotypes which may be evaluated later by microfacies data. Cyclothems that are evident already from field observations are often described using only a few samples. This may provide false information because shallowing- or deepening trends of the cycles could be missed. In addition, cyclic sedimentation patterns may be obscured by the amalgamation of beds, thus calling for densely spaced samples (Gischler et al. 1994). Massive to coarse-bedded carbonates (reefs, mounds) require the study of many samples in order to consider (1) the variability of sediment within and between frame-building and baffling organisms, (2) the various kinds of voids and open-space structures common in mounds and reefs, (3) lateral microfacies changes on a decimeter-scale, and (4) relations between mound facies and intermound facies. Microfacies analysis of bedded deep-water carbonates often is based on rather few samples because of the assumed ‘monotonous’ microfacies spectra of these limestones. The calculated sampling interval (profile thickness/number of samples) of most case studies lies between 1.00 and 2.50 m. Wheras the overall qualitative composition of deep-water limestones may indeed be uniform, the quantitative investigation of microfa-
cies criteria and microfossil distribution, based on a ‘one-centimeter style’ sampling exhibits hidden aspects of environmental controls (Noé 1993). Hardgrounds and condensed pelagic limestones should be studied in channel samples covering the total bed thickness and with particular consideration of the bed surfaces, because microfacies types can change within a distance of a few millimeters and centimeters. Collecting samples for paleoecological or biostratigraphical analysis must take into account the strong facies control of the frequency and distribution of fossils. Many microfacies studies aiming for biostratigraphical classifications are based on spot samples taken at regular intervals or over a regular number of event beds (e.g. calciturbidites). Channel samples continuous over the total thickness of limestone beds are necessary to understand fine-scale vertical variations in event beds. These variations describe the hydraulic changes during the formation of the calciturbidites.
3.1.2.3 Practical Advice for Microfacies Sampling 1 Become clear about the principal aim of the investigation and decide whether you will use random sampling or systematic sampling strategies. 2 Start with a test section to understand the scale of microfacies variations and choose the most appropriate vertical sampling distance. 3 Mark the orientation of the sample (sedimentary top and, if necessary dip and strike, which may be important for the interpretation of grain orientation as a current indicator) before removing the sample from the bed. 4 Take samples of an appropriate size allowing for the preparation of thin sections, the study of non-carbonates in acid residues and geochemical analyses. 5 Consider the negative effect of weathering and avoid taking samples for geochemical and diagenetic studies at or near bedding surfaces. 6 Follow a strict sampling strategy, but do not exclude sampling of beds showing rare but conspicuous criteria (e.g. extremely fossiliferous beds, beds exhibiting differences in colors). But be careful not to follow the ‘family photo album effect’ by sampling predominantly ‘extraordinary’ lithotypes. 7 Locate microfacies sampling in those parts of beds which will deliver a maximum potential on information. 8 When working with bedded carbonate sequences, do not forget to sample also non-carbonate interlayers as well.
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9 Do not restrict sampling to microfacies alone, but also collect macrofossils and specific micropaleontological samples. 10 Record the sampling locality and the beds sampled in field photographs and detailed sketches. Basics: Field work Field methods and graphic record Allen, J.R.L. (1982): Sedimentary structures. Their character and physical basis. – Development in Sedimentology, 30, 1258 pp., Amsterdam (Elsevier) Collinson, J.D., Thompson, D.B. (1988): Sedimentary structures. – 2nd edition, 208 pp., London (Allen and Unwin) Dietrich, R.V., Dutro, J.T., Foose, R.M. (1982): AGI data sheets for geology in the field, laboratory, and office. – 2nd edition, Boulder (American Geological Institute) Gagliardi, R., Kälin, O., Morreale, E., Pataca, E., Scandone, F., Sprugnole, R. (1980): Una banca dati geologici del progetto finalizzato geodinamica. – Consiglio Nazionale delle Ricerche. Progetto Finalizzato Geodinamica, Pubblicazione, 393, 1-90 Leeder, M.R. (1982): Sedimentology. Process and product. – 344 pp., Boston (Allen and Unwin) Lewis, D.W., McConchie, D.M. (1994): Practical sedimentology. 2nd ed. – 256 pp., New York (Chapman and Hall) Reijers, T.J.A., Mijnlieff, H.F., Pestman, P.J., Kouwe, W.F.P. (1993): Lithofacies and their interpretation: a guide to standardised description of sedimentary deposits. – Mededelingen Rijks Geologische Dienst, 49, 1-55 Shell International Exploration and Production B.V., The Hague (1995): Standard Legend 1995. – Document and CD-ROM version Tucker, M.E. (1982): The field description of sedimentary rocks. – Geological Society of London, Handbook Series, 2, 112 pp., London (Open University Press) Tucker, M.E. (ed., 1988): Techniques in sedimentology. – 416 pp., Oxford (Blackwell) Bedding, sedimentary and biogenic structures Bromley, R.G. (1996): Trace fossils: biology, taphonomy and applications. – 2nd ed., Dordrecht (Kluwer) Cole, R.D., Picard, M.D. (1975): Primary and secondary sedimentary structures in oil shale and other fine-grained rocks, Green River Formation (Eocene), Utah and Colorado. – Utah Geology, 2, 49-67 Campbell, C.V. (1967); Lamina, laminaset, bed and bedset. – Sedimentology, 8, 7-26 Droser, M.L., Bottjer, D.J. (1986): A semiquantitative field classification of ichnofabric. – J. Sed. Petrol., 56, 558-599 Goldring, R. (1991): Fossils in the field. Information potential and analysis. – 218 pp., London (Longman) Kendall, C.G.St.C., Warren, J. (1987): A review of the origin and setting of tepees and their associated fabrics. – Sedimentology, 34, 1007-1027 Kidwell, S.M., Holland, S.M. (1991): Field description of coarse bioclastic fabrics. – Palaios, 6, 426-454 Lewis, D.W., McConchie, D.M. (1993): Practical sedimentology. – 256 pp., London (Chapman and Hall) Lumsden, D.N. (1971): Facies and bed thickness distributions of limestones. – J. Sed. Petrol., 41, 593-598 Plummer, P.S., Gostin, V.A. (1981): Shrinkage cracks: desiccation or synaeresis. – J. Sed. Petrol., 51, 1147-1156
Practical Advices
Rock colors Wells, N.A. (2002): Quantitative evaluation of color measurements: I. Triaxial stereoscopic scatter plates. – Sedimentary Geology 151, 1-15 Folk, R.C. (1969): Toward greater precision in rock-color terminology. – Geol. Soc. America Bull., 80, 725-728 Flügel, E., Tietz, G.F. (1971): Über die Ursachen der Buntfärbung in Oberrhät-Riffkalken (Adnet, Salzburg). – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 139, 29-42 Goddard, E.N., Trask, P.D., De Ford, R.K., Rove, O.N., Singewald, J.T., Overbeck, R.M. (1948): Rock-color chart. – Geological Society of America Patnode, H.W. (1941): Relation of organic matter to color of sedimentary rocks. – Amer. Ass. Petrol. Geol. Bull., 25, 19211933 Wright, W.D. (1969): The measurement of color. – 340 pp., London (Hilger) Sampling Griffith, J.C. (1967): Scientific method in analysis of sediments. – 508 pp., New York (McGraw-Hill) Schindel, D.W. (1980): Microstratigraphic sampling and the limits of paleontologic resolution. – Paleobiology, 6, 408426 Further reading: K003, K004, K014, K015, K016
3.2 Laboratory Work: Techniques Originally strongly focused on the study of thin sections, peels or slices with the aid of low to medium power binocular and petrographical microscopy, the combined use of conventional thin-section data and more sophisticated detailed microscopy (SEM, cathodoluminescence and fluorescence microscopy) as well as mineralogical and geochemical studies have proved to be of high potential in the application of microfacies analysis. Valuable summaries of techniques are found in Müller (1967), and especially in Tucker (1981). The latter book includes excellent overviews of diverse microscopical methods, cathodoluminescence microscopy, scanning electron microscopy, X-ray powder diffraction and chemical analysis of sedimentary rocks. Texts covering the paleontological aspects of microfacies studies are listed in Sect. 10.4. 3.2.1 Slices, Peels and Thin Sections The composition, fabric and texture of carbonate rocks are studied in slices, peels, or thin sections which may be etched or stained in various phases of preparation in order to emphasize specific criteria. Various techniques and methods are described by Adams et al. (1984) and Miller (1988). Slices: Oiling and gentle etching of cut and smoothened carbonate rock faces exhibit otherwise cryptic fab-
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Laboratory Work
rics or show fabrics within apparently homogeneous limestones (Bromley 1981). Peels: The imprint of etched surfaces on transparent plastic films is the most rapid and simplest method for studying cut, polished, and etched samples. Acetate replicas can be made using a dry-peel technique (McCrone 1963; Germann 1965) or by making peel sheets by pouring solutions (ethyl cellulose in trichloroethylene, nitrocellulose) on to glass plates (Davies and Till 1968). The procedure of preparing peels includes (1) cutting, grinding (and polishing) the sample, (2) drying and etching the cut surface, (3) flooding the surface with a solvent and (4) pressing the film onto the surface; the film will settle down into irregularities in the etched surface and produce the replica. After pulling off, the peel should be pressed or mounted between glasses. Peels have many advantages: • They offer a rapid and inexpensive method to obtain a first observation of the characteristics of the sample. • Peels can be used not only for carbonates but also for cherts, siliceous clastics and sulfates rocks (Mandado and Tena 1986). • Peels can be studied just like thin sections using a binocular microscope or a simple microfiche equipment and they can be placed in the negative carrier of a photographic enlarger, producing large negative prints by projecting the image directly to bromide paper. These prints may be used as working copies exhibiting locations of areas of particular interest. • Replicas can be made of large surfaces, even up to several tenths and hundredths of square centimeters. • Serial peels can be produced by regrinding surfaces after each peel, thus allowing the three-dimensional reconstruction of open-space structures or fossils only known from thin sections (Honjo 1963). • Staining peels prior to the flooding of rock surfaces allows minerals and fabrics to be differentiated (Katz and Friedman 1965). Since earlier stains can be removed by repolishing, different stains indicating different minerals can be used with good success. • Peels are suited for high-magnification studies and phase contrast microscopy (Frank 1965; Honjo and Fischer 1965) and for SEM studies (Brown 1986). • Acetate peels have proved to be of value in studying carbonate cements and textural criteria (e.g. grainsize distribution: Gutteridge 1985). Disadvantages of peels are the susceptibility of acetate films to contamination and the danger that porous samples might be difficult to work with, because the film may bulge into the pores, causing blisters. Ac-
etate films are isotropic; the main disadvantage of peels, therefore, is that minerals can not be identified by optical properties, such as relief or birefrigence. Thin sections: The investigation of the composition and fabric of surface and subsurface samples requires standard thin sections with a thickness of approximately 30 µm. These thin sections are usually cut perpendicular to the bedding. Initially, the thin sections should not be covered, allowing later more detailed investigations (staining etc.). Very fine-grained micrites are studied in high-quality ultra-thin sections with a thickness between 0.5 and 5 µm (Lindholm and Dean 1973), allowing observations at magnifications of between x 150 and x 1000. Thin sections may be subject to etching and staining (Friedman 1959; Houghton 1980; Rassineux and Beaufort 1987). These procedures and other more detailed studies (e.g. cathodoluminescence microscopy: Mugridge and Young 1984) only can be applied to uncovered thin sections. Rapid covering can be done using a plastic spray (Moussa 1976). Microfacies thin sections are generally larger and thicker than regular petrographic slides. Commonly used sizes are from 5 x 5 up to 15 x 20 cm. Large thin sections are necessary for studying reef limestones, coarse-grained detrital limestones and carbonates deposited during events (e.g. turbidites, tempestites). • Note: Samples are generally cut perpendicularly to bedding planes. Therefore, a thin section comprising a vertical interval of 5 cm records a set of ‘microsequences’ corresponding to shorter or longer time intervals within a scale of hundreds, thousands or more years.
3.2.2 Casts, Etching and Staining Casts: Impregnation of carbonate rocks and fossils with low viscosity epoxy resin under vacuum and under pressure has a high potential in recognizing micron- and submicron-scaled structures in porous but also moderately porous carbonates and evaporites. Impregnation of carbonate rocks with a mixture of fluorescent dye and epoxy resin enhance the ability to study micro- and macroporosity. SEM observations of artificial casts aid in the study of • microporosity and pore geometries of limestones and chalks (Gardner 1980; Beckett and Sellwood 1991; see Pl. 6), • modern and ancient microborings (Golubic et al. 1975; see Pl. 52/1-4). Etching and staining: Etching of slices or thin sections with hydrochloric acid (1–5% by volume) or acetic acid (about 20% by volume) causes partial dissolu-
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tion that emphasizes textural criteria and allows subsequent staining, which is one of the most useful techniques in carbonate rock studies. Comparing staining techniques with other methods, e.g. X-ray diffraction, may exhibit discrepancies in estimates of dolomite in limestones (Gensmer and Weiss 1980). The success of etching and staining depends on (a) the strength and temperature of the reagents, (b) etching and staining times, (c) the pre-treatment of the sample, and d) the composition and fabric. An advantage of staining is the possibility using the method already used during field studies (Warne 1962). Staining assists in (1) identifying minerals common in carbonates (aragonite, Mg-calcite, calcite and dolomite, gypsum and anhydrite, feldspars), (2) studying dedolomitization (Evamy 1963), (3) recognizing microporosity (Yanguas and Dravis 1985).
3.2.3 Microscopy The microscopic study of thin sections of carbonate rocks provides the basic microfacies data. The combination of microfacies analysis with the criteria derived from fluorescence, photoluminescence, cathodoluminescene and fluid inclusion studies provide key information on sedimentary and diagenetic processes.
3.2.3.1 Petrographic Microscopy Good polarizing microscopes and binocular microscopes are essential tools for microfacies studies. Microscopes should have photographic equipment low and high power magnification and should be equipped with mechanical or computer-interfaced stages for examining modal composition and fabrics (Kobluk and Vyas 1989). The speed, thoroughness and accuracy of thin-section analysis can be increased by using integrated computerized systems (Kobluk and Kim 1991). To provide an overview of the whole thin section, the use of digital film scanners (De Keyser 1999) and microfiche instruments is strongly recommended. Microfacies observations should always start with low-power examinations and continue to high-power magnification. Primary textural features in dolostones and recrystallized limestones can be made visible by using a simple light diffuser (translucent plastic or plain white paper), placed directly under the thin section (Delgado 1977). There are many excellent and colored texts on mineral identification in thin sections (e.g. Scholle 1979; Adams et al. 1984; Carozzi 1993).
Casts, Etching, Staining
3.2.3.2 Stereoscan Microscopy Utilization of Steroscan Electron Microscopy (SEM) in the study of ancient and modern carbonate rocks has brought about major breakthroughs in our understanding of depositional and diagenetic products and processes (Welton 1984; Trewin 1988). The chief advantages of the SEM are the great depth of focus, producing excellent photographs of extremely small (micron- and submicron-sized) three-dimensional surfaces, the high power of resolution and the possibility of very high magnifications in the range between x 10 up to x 100 000. Transmission Electron Microscopy (TEM) of replicas, however,which has been used prior to SEM in the study of carbonate rocks (Fischer et al. 1967), is still a valuable tool in carbonate investigation (e.g. biomineralization modes, microbial contribution to the formation of carbonates). Today, SEM studies are a part of integrated analyses which combine SEM with energy dispersive X-ray (EDX) data, back-scattered electron images and image analysis, cathodoluminescence microscopy and conventional studies of (stained or etched) peels and thin sections. Peels and thin sections can be used directly for SEM studies. EDX studies show element composition and element distribution. Element mapping is used to distinguish mineralogical phases and to elucidate Mg and Sr distributional patterns in cements and fossils. Carbonate SEM samples should be unaffected by weathering. Nonporous limestones and dolomites are studied using untreated fractures of rock chips or slices (size about 5 by 10 by 10 mm) which are polished and etched with dilute hydrochloric acid or formic acid, or Titriplex-III. Samples are mounted on stubs by doublesided sticky tape or various glues. Porous carbonates may need specific sample treatment (Trewin 1988). The major applications of SEM in investigating carbonate rocks include studies of the • matrix of fine-grained limestones (micrite, chalk) and the composition of modern carbonate muds, • composition and ultrastructure of fossils and nonbiogenic carbonate grains, • breakdown of fossils, including the importance of endolithic borings in the destruction of grains and the formation of micritic coatings, • type, sedimentological and biostratigraphical role of micro- and nannofossils, • distribution and geometry of micropores, • mineralogy, morphology and spatial relations of carbonate cements, • stages of dolomitization,
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Microscopy
• type and distribution of non-carbonate minerals, • changes taking place during burial diagenesis and incipient metamorphosis. • 3.2.3.3 Fluorescence, Cathodoluminescence and Fluid Inclusion Microscopy Fluorescence is a form of luminescence and represents the property of material to emit light when exited by visible or ultraviolet light (Rost 1992). Fluorescence microscopy is a standard technique in coal petrography, paleobotany, and the organic petrography of sediments of mainly organic origin or with a high organic content (e.g. petroleum source rocks). Fluorescence in limestones is probably caused by organic matter (Machel et al. 1991). Samples are studied as polished slices and thick thin sections (Clausing 1991). Epifluorescence in carbonates is used for • recognizing hidden microfabrics (Dravis and Yurewicz 1985; Klotz 1989), • studying organic constituents of fine-laminated limestones (e.g. cyanobacteria, algae, pollen and spores), • differentiating automicrites and allomicrites (Neuweiler 1995), • understanding biomineralization processes. Cathodoluminescence (CL) microscopy of carbonate rocks, starting with the work of Sippel and Glover (1965), is a significant tool in the petrography of marine and non-marine carbonates (Marshall 1988; Barker and Kopp 1991; Barbin et al. 1991) and in the investigation of primary or diagenetic microstructures of fossils (Amieux 1987; Elorza et al. 2001). CL microscopy stimulates luminescence on polished thin sections (thickness < 30 µm) and polished rock surfaces by electron bombardment. Luminescence depends on the material characteristics of the excited solid. These characteristics comprise the chemistry, crystal structure, lattice defects and other factors (Hemming et al. 1989). The main activator elements of carbonates are Mn2+ and Pb+2. Activation of calcite and dolomite by Rare Earth elements is also a common feature of Phanerozoic marine limestones and sinter calcites (Habermann et al. 1996). The subjective determination of CL colors should be supplemented by spectrometric analyses. CL studies are often combined with petrographic microscopy (Gregg and Karakus 1991) and rare element microanalyses, and have become a significant part of microfacies analyses. The major applications of CL microscopy for carbonate rocks (see Pl. 5) are for • observing and interpreting of diagenetic phases (e.g. zonar structures within crystals reflecting changes in
•
• •
the chemical environment and/or in the growth speed; early diagenetic and burial cements reflecting burial depths; mineral transformations, such as the alteration of calcite to dolomite (Richter 1984), evaluating pore water chemistry, reflected by sector zoning in calcite crystals (Bruckschen et al. 1992), and interpretation of the diagenetic history of reservoir rocks, reconstructing diagenetic events which identify stages of cementation and provide a ‘cement stratigraphy’ that are used to establish cement-based stratigraphic correlations (Meyers 1991), recognizing fabrics (e.g. distinction of recrystallized structures, recognition of microcracks), studying the diagenesis of fossils.
Fluid inclusions in minerals have been studied during the last twenty years with respect to the genesis and migration of hydrocarbons, mineral deposits, diagenesis of sedimentary rocks and problems of structural geology and petrology (Roedder 1984, 1990; Lattanzi 1991). Fundamental for the interpretation of fluid inclusion data is the assumption, that the fluid inclusions still have the same composition and volume as at the time of formation in a closed system. Alterations may be caused by changes in geometry and volume or by a temporary or permanent opening of the inclusions. Carbonate crystals with fluid inclusions appear dark in thin sections and polished sections. Cloudy inclusions of aqueous liquid and vapor within carbonate crystals provide information on the diagenetic environment of cements (vadose zone, low-temperature phreatic zone, high-temperature burial zone) and their thermic history. The size of the inclusion ranges from < 10 µm to 30 µm. Fluid inclusions are caused by primary engulfment of fluids within growing crystals or by secondary entrapment of fluids in microfractures that are subsequently healed. Samples should not have been subjected to significant overheating beyond the temperature of entrapment or the temperature of homogenization (Barker and Goldstein 1990) and should have been reequilibrated during recrystallization and deformation. Calcite crystals should display distinct growth zones. Genetic interpretations and classifications are based on the composition, form, and position of the inclusions as well as their relationship to crystal structures (Roedder 1984; Goldstein 1993). Investigations require a petrographical microscope (with UV epifluorescence capabilities for differentiating aqueous and hydrocarbon inclusions) and a heating and freezing stage. Fluid inclusion studies are com-
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Cathodoluminescence
Plate 5 Cathodoluminescence (CL): A Possibility to See More Cathodoluminescence microscopy of thin sections has become an important tool of microfacies analysis.It assists in observing diagenetic phases recorded by different cement types (–> 1, 2, 8) and establishing a ‘cement stratigraphy’as well as recognizing zonation patterns of calcite crystals and the mode of porosity occlusion (–> 3). CL reveals the original structures of fossils and their growth modes (–> 4) and facilitates the recognition of diagenetic not or little altered shells which can be used for the determination of stable isotopes. Generally CL accentuates diagenetic and sedimentary textures (–> 6, 7), thus allowing a more precise observation of microfacies criteria. The combined use of CL and petrographic microscopy is strongly recommended. 1 and 2 Primary interparticle pore in transmitted light (–> 1) and under CL (–> 2). Note the different appearance of cement successions. The succession consists of a light-brown luminescent radialfibrous cement (A) circumcrusting a grain followed by non-luminescent dogtooth cement (B) and banded yellow to brown luminescent burial coarse mosaic calcspar cement which fills the remaining pore space. Early Jurassic: Carbonate platform of Jbel Bou Dahar, High Atlas, Morocco. 3 Zoned crystal growth. The luminescence pattern exhibits the distinct zonation of a large crystals filling the pore space in a travertine. Quaternary: Tivoli near Rome, Italy. 4 Ulltastructure and growth modes of fossils: Shell of the inoceramid species Antarcticeramus rabotensis (Crame and Luther). Under cathodoluminescence a distinct yellow luminescent lamina is visible which interrupts the normal sized prisms constructing the shell. A recovery to normal crystal sizes is rather slow as seen by the smaller prisms (here) left of the yellow lamina. This interruption in shell growth is interpreted as the result of a cooling interval or even a freeze shock which influenced the growth of prisms in the shell. Growth direction of the shell is from right to the left in this picture. Late Cretaceous (Campanian): James Ross Island, Antarctica. 5 Grade of diagenetic alterations of shells. Punctate brachiopod. The brachiopod shell appears dark (unaltered) with some luminescent parts (diagenetically altered). The punctae are filled by orange luminescent diagenetic calcite. Sediment in the lower part shows bright orange luminescence due to Mn2+ incorporation (‘activator’ ion) during the diagenetic transformation of metastable aragonite and High-Mg calcite to the stable phase Low-Mg calcite and are therefore relatively stable to diagenetic alteration. Unaltered shells usually show no luminescence. The shell belongs to the genus Schizophoria. Devonian: Iowa, USA. 6 Diagenetic stages: Relictic ooids. The ooid layers were dissolved mainly during late burial diagenesis; the preserved core consists of recrystallized shell fragments which show a weak red to orange luminescence. Open moldic porosity appears dark. The outline of the ooids is traced by orange to red luminescent early diagenetic micritic envelopes and fine crystalline cements. Interparticle porosity between ooids is filled by coarse crystalline sparite showing bright orange luminescence. This is interpreted as the result of late diagenetic neomorphism of an original early diagenetic marine-phreatic sparite. Late Cretaceous: Sinai Peninsula, Egypt. 7 Dolomitization and compaction. Dolomite rhomboids. The rhomboids have a non-luminescent to weakly luminescent core surrounded by a bright orange luminescent zone. The weak luminescence of the cores indicates iron-rich diagenetic fluids and the incorporations of Fe2+ in the dolomite lattice (‘quench’ ion) during crystal growth. By contrast, the final bright orange luminescent zone indicates changing pore fluid chemistry with more Mn2+ incorporation in the crystal lattice. Broken cores cemented by the outer luminescent zone indicate ongoing compaction during dolomitization. The rhomboids are interpreted as baroque dolomite typical of late burial diagenesis. Same locality as –> 6 8 Diagenetic filling of a fossil (brachiopod shell not visible in figure). Left part shows sediment (micrite) with red to orange luminescence due to diagenetic transformation to Low-Mg calcite and the incorporation of Mn2+ in the calcite lattice. Remaining porosity was cemented by large calcite crystals (right) formed during late diagenesis. The distinct zoning of the crystals with weak and bright orange luminescent zones indicates slow crystal growth, changing pore water chemistry and/or changing redox conditions during crystal formation. The latest cements exhibits bright orange to yellow luminescence and document the final stage of the diagenetic cements. Middle Devonian: Anti-Atlas, Morocco. –> 1, 2: Blomeier and Reijmer 1999; 3: Courtesy of R. Koch (Erlangen); 4: Elorza et al. 2001; 5-8: Courtesy of R. von Geldern (Erlangen)
Plate 5: Cathodoluminescence
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70
Trace Elements and Stable Isotopes
monly combined with petrographic and cathodoluminescence microscopy and isotope analysis (Sotak and Lintnerova 1994). Fundamental techniques have been described by Hollister and Crawford (1981), and Goldstein (1993). Fluid inclusions in carbonates are a powerful tool for the • history and timing of shallow and burial diagenetic and low-grade metamorphic processes ( Brockerhoff and Friedman 1987; Liedmann 1992), • conditions of dolomitization (Goldstein et al. 1990; Montanaze and Read 1992), • migration processes of paleofluids (Burruss et al. 1985; Ricken 1992; Muchez et al. 1994), • uplift and subsidence history of sedimentary rocks (McLimans 1987), • subaerial exposure and karst phenonema (Goldstein et al. 1990), • data on pore water chemism and temperature during the formation of geodes, vugs and veins.
3.2.4 Mineralogy and Geochemistry Carbonate minerals include minerals composed of CO32- and one or more cations. Approximately sixty minerals occur in nature with a CO3 group in common. Of these the trigonal CaCO3 (calcite) and its orthorhombic polymorph (aragonite) are the most common in modern sediments. Calcite and dolomite are by far the most common carbonate minerals in ancient carbonate rocks. Several cations can substitute in varying amounts for the Ca2+ in the crystal structure: Mg2+, Fe2+ and Mn 2+ are more readily accepted in the hexagonal calcite structure, and Sr2+ and Ba2+ by the aragonite structure. The minor element composition of calcite and aragonite is determined by biochemical and physical-chemical processes.
The most important carbonate minerals are LowMg calcite (LMC), High-Mg calcite (HMC), aragonite and dolomite (Fig. 3.8). The dividing line between LowMg calcite and High-Mg calcite is commonly drawn at 4 mol % MgCO3. These groupings can be subdivided into those containing a variable amount of MgCO3 (4–12 mol %) and those containing a consistently high amount of MgCO3 (12–28 mol %, e.g. corallinacean red algae, echinoderms). The latter group represents the HMC sensu stricto (hHMC = high High-Mg calcite). The amount of Mg substitution for Ca in the calcite structure is determined by X-ray diffraction analysis. The composition of carbonate sediments is studied by petrographical microscopy and X-ray diffraction (XRD), which is particularly useful for analyzing finegrained carbonate rocks. The qualitative and quantitative methods of XRD whole rock analysis and clay mineral analysis are described by Hardy and Tucker (1988). XRD data provide information on the chemical composition of carbonate minerals, allow carbonate minerals to be distinguished, and are used to determine the ordering of dolomite crystals and the proportions of calcite and dolomite. The ‘carbonate bomb’ (Müller and Gastner 1971) offers the possibility of rapidly determinating the contents of CaCO3 and insoluble residues in limestones. 3.2.5 Trace Elements and Stable Isotope Analysis Compare also Sect. 13.1 for the use of geochemical studies of ancient limestones and the advantage of combined investigations of microfacies, stable isotopes and trace elements. Trace elements: The elemental analysis of carbonate rocks provides important data on the sedimentary and diagenetic history. Atom Adsorption Spectrophotometry
System
Mol % MgCO3
Stability
Calcite CaCO3 Low-Mg calcite (LMC)
trigonal
<4
stable
Mg-calcite CaCO3 High-Mg calcite (HMC)
trigonal
>4 to ~30 Mg content correlated with water temperature
metastable
Aragonite CaCO3
orthorhombic
very low
metastable, alters readily in calcite under aqueous conditions
Dolomite
trigonal CaMg(CO3)2
40 - 50
stable
Fig. 3.8. Common carbonate minerals.
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Mineralogy Components Skeletal Non-skeletal
Bulk sediments
Isotope signal Preservation potential
Mineralogy of components
HIGH
Pristine aragonite
Mollusks
Marine cements
Good chance of preservation of carbon and oxygen isotope signals
Pristine Low-Mg calcite fossils, grains or cements
Brachiopods, belemnites, foraminifera, bivalves
Marine cements, Low-Mg calcite ooids
Pelagic sediments, particularly coccolith oozes
Phosphatic fossils
Conodonts, fish teeth
MODERATE
Secondary calcites
Carbon isotope signals may be preserved, oxygen isotope signals are commonly altered
(stabilized in relatively closed systems with low water/rock ratio)
Molluscs, foraminifera, corals, echinoderms, calcareous algae
Marine cements, ooids, peloids, intraclasts
Some micrites, some shallowwater carbonates, some dolomites
LOW Carbon and oxygen isotope signals very likely to have been altered
Secondary calcites (stabilized or cemented in relatively open systems with high water/rock ratio)
Limestones altered by near surface meteoric diagenesis of intensive cementation or recrystallization during burial; many dolomites
Fig. 3.9. Preservation potential of carbon and oxygen isotopes in ancient carbonate sediments and skeletal and non-skeletal constituents, and phosphatic fossils. Modified after Marshall (1992).
(AAS) is widely used in the analysis of limestones and dolomites because of elements hidden in minor amounts within the carbonate lattice. Inductively-coupled Plasma Atomic Emission Spectrometry Analysis (ICP-AES) and Neutron Activation Analysis (NAA) are used for determining the contents of trace elements and rare earth elements in carbonate rocks by whole-rock and selective analyses (Fairchild et al. (1988). Minor elements in carbonate rocks and in fossils are important paleoenvironmental indicators. The combined use of microfacies data, trace elements and isotope geochemistry in facies analyses continues to evolve in current studies (see Chap. 13). Stable Isotopes: The isotope ratios of 18O/16O and 13C/ are very often used to trace processes involved in the production of carbonate grains and their subsequent conversion into limestones. Isotope studies are performed as whole-rock bulk analyses or analyses of single components (e.g. isolated fossils, carbonate cements) or as microsamples of individual parts of fossils (e.g. shells) or cement crystals. The preservation potential of isotope signals depends on the preservation of the primary mineralogy and the degree of diagenetic alterations (see Fig. 3.9). Diagenesis can obscure the original composition of constituents of carbonate rocks, and if undetected, can lead to erroneous conclusions. All samples, therefore, must be tested prior to isotope analysis with respect to potential diagenetic overprinting, using petrographic and CL microscopy, SEM criteria and chemical data. Because shells con12C
sisting of Low-Mg calcite are rather resistant to diagenetic alteration, isotope analyses of brachiopods are enjoying increasing popularity. Stable isotope analyses of ancient carbonates provide proxies for paleoenvironmental constraints (e.g. temperature and salinity variations of the sea water, terrigenous input), unravel changes in carbon cycling of ancient ocean-atmosphere systems and document global fluctuations in the climate. Isotope abundance measurements for geochemical research are determined on a mass spectrometer. Basics: Laboratory work Techniques and microscopical methods Beckett, D., Sellwood, B.W. (1991): A simple method for producing high-quality porecasts of carbonate rocks. – Sed. Geol., 71, 1-4 Brown, R.J. (1986): SEM examination of carbonate microfacies using acetate peels. – J. Sed. Petrol., 56, p. 538 Delgado, F. (1977): Primary textures in dolostones and recrystallized limestones: a technique for their microscopy study. – J. Sed. Petrol., 47, 1339-1341 Folk, R.L. (1987): Detection of organic matter in thin-sections using a white card. – Sed. Geol., 54, 193-200 Friedman, G.M. (1959): Identification of carbonate minerals by staining methods. – J. Sed. Petrol., 29, 87-97 Gardner, K.L. (1980): Impregnation technique used colored epoxy to define porosity in petrographic thin sections. – Canadian Journal of Earth Sciences, 17, 1104-1107 Germann, K. (1965): Die Technik des Folienabzuges und ihre Ergänzung durch Anfärbemethoden. – Neues Jahrbuch für Geologie und Paläontologie Abhandlungen, 121, 293-306 Gies, R.M. (1987): An improved method for viewing micropore systems in rocks with the polarizing microscope. – Soc. Petroleum Engineers Formation Evaluation, 1987, 209-214
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Gutteridge, P. (1985): Grain-size measurement from acetate peels. – J. Sed. Petrol., 55, 595-596 Katz, A., Friedman, G.M. (1965): The preparation of stained acetate peels for the study of carbonate rocks. – J. Sed. Petrol., 35, 248-249 Kobluk, D.R., Kim, K.H. (1991): A new microcomputer-assisted approach to the study of thin sections of carbonate rocks. – Bull.Canadian Petroleum Geology, 39, 252-259 Mandado, J.A., Tena, J.M. (1986): A peel technique for sulfate (and carbonate) rocks. – J. Sed. Petrol., 56, 548-549 Miller, J. (1989): Microscopic techniques. I. Slices, slides, stains and peels. – In: Tucker, M. (ed.): Techniques in sedimentology. – 86-107, Oxford (Blackwell) Moussa, M.T. (1976): Plastic spray thin-section cover. – J. Sed. Petrol., 46, 252-253 Mugridge, S.J., Young, H.R. (1984): Rapid preparation of polished thin sections for cathodoluminescence study. – Canadian Mineralogists, 22, 513-515 Müller, G. (1964): Methoden der Sedimentuntersuchung. – 303 pp., Stuttgart (Schweizerbart) Murphy, C.P. (1986): Thin section preparation of soils and sediments. – 149 pp., Berkhamsted (Academic Publ.) Ruzyla, K., Jezek, D.I. (1987): Staining method for recognition of pore space in thin and polished sections. – J. Sed. Petrol., 57, 777-778 Trewin, N. (1988): Use of the scanning electron microscopy in sedimentology. – In: Tucker, M. (ed.): Techniques in sedimentology. – 229-273, Oxford (Blackwell) Tucker, M. (ed., 1988): Techniques in sedimentology. – 394 pp., Oxford (Blackwell) Yanguas, J.E., Paxton, S.T. (1986): A new technique for preparation of petrographic thin-sections using ultraviolet-curing adhesive. – J. Sed. Petrol., 56, 539-540 Warne, J. (1962): A quick field or laboratory staining scheme for the differentiation of the major carbonate minerals. – J. Sed. Petrol., 32, 29-38 Wolf, K.H., Easton, A.J., Warne, S. (1967): Techniques of examining and analyzing carbonate skeletons, minerals and rocks. – In: Chilingar, G.V., Bissell, H.J., Fairbridge, R.W. (eds.): Carbonate rocks: physical and chemical aspects. – Developments in Sedimentology, 9B, 253-341 Fluorescence and cathodoluminescence microscopy Amieux, P. (1982): La cathodoluminescence: méthode d’étude sédimentologique des carbonates. – Bull. Centre Rech. Explor.-Prof. Elf-Aquitaine, 6, 437-483 Barbin, V. (1995): Cathodoluminescence of carbonates: new applications in geology and archaeology. – Scanning Microscope Supplement, 9, 113-123 Barker, C.E., Kopp, A.C. (eds., 1991): Luminescence microscopy and spectroscopy. Quantitative and qualitative applications. – SEPM Short Course, 25, 195 pp. Fairchild, I.J. (1983): Chemical controls of cathodoluminescence of natural dolomites and calcites: new data and review. – Sedimentology, 30, 579-583 Gregg, J.M., Karakus, M. (1991): A technique for successive cathodoluminescence and reflected light petrography. – J. Sed. Petrol., 61, 613-614 Hemming, N.G., Meyers, W.J., Grams, J.C. (1989): Cathodoluminescence in diagenetic calcites: the roles of Fe and Mn as deduced from electron probe and spectrophotometric measurements. – J. Sed. Petrol., 59, 401-411 Machel, H.G., Mason, R.A., Marino, A., Mucci, A. (1991):
References Laboratory Work
Causes and emission of luminescence in calcite and dolomite. – In: Barker, C.E., Kopp, O.C. (eds.): Luminescence microscopy and spectroscopy: qualitative and quantitative applications. – SEPM Short Course, 25, 9-25 Marshall, D.J. (1988): Cathodoluminescence of geological materials. – 146 pp., London (Unwin Hyman) Marshall, D.J. (1991): Combined cathodoluminescence and energy dispersive spectroscopy. – SEPM Short Course, 25, 27-35 Meyers, W.J. (1991): Calcite cement stratigraphy: an overview. – In: Barker, C.E., Kopp, O.C. (eds.): Luminescence microscopy and spectroscopy: qualitative and quantitative applications. – SEPM Short Course, 25, 133-148 Miller, J. (1988): Cathodoluminescence microscopy. – In: Tucker, M. (ed.): Techniques in sedimentology. – 174-190, Oxford (Blackwell) Neuser, R.D., Bruhn, F., Götze, J. Habermann, D., Richter, D.K. (1996): Kathodolumineszenz: Methodik und Anwendung. – Zentralblatt für Geologie und Paläontologie. Teil I, 1995, 287-306 Pagel, M., Barbin, V., Blanc, P., Ohnenstetter, D. (2002): Cathodoluminescence in geosciences. – 514 pp., Berlin (Springer) Ramseyer, K., Fischer, J., Matter, A., Eberhardt, P., Geiss, J. (1989): A cathodoluminescence microscope for low intensity luminescence. – J. Sed. Petrol., 59, 619-622 Richter, D.K., Zinkernagel, U. (1981): Zur Anwendung der Kathodolumineszenz in der Karbonatpetrographie. – Geologische Rundschau, 70, 1276-1302 Rost, F.W.D. (1992): Fluorescence microscopy. Volume I. – 253 pp., New York (Cambridge University Press) Mineralogy, geochemistry and stable isotopes Arthur, M.A., Anderson, T.F., Kaplan, I.R., Veizer, J., Land, L.S. (1983): Stable isotopes in sedimentary geology. – SEPM Short Course, 10, 435 pp. Clauer, N., Chaudhuri, S. (eds., 1992): Isotopic signatures and sedimentary records. – Lecture Notes in Earth Science, 43, 529 pp. Dickson, T. (1990): Carbonate mineralogy and chemistry. – In: Tucker, M.E., Wright, V.P: Carbonate sedimentology. – 284-313, Oxford (Blackwell) Fairchild, I., Hendry, G., Quest, M., Tucker M. (1988): Chemical analysis of sedimentary rocks. – In: Tucker, M. (ed.): Techniques in sedimentology. – 274-354, Oxford (Blackwell) Faure, G. (1986): Principles of isotope geology. – 589 pp., New York (Wiley) Hardy, R., Tucker, M. (1988): X-ray powder diffraction of sediments. – In: Tucker, M. (ed.): Techniques in sedimentology. – 191-228, Oxford (Blackwell) Hoefs, J. (1997): Stable isotope geochemistry, 4th edition. – 201 pp., Berlin (Springer) Moore, C.H. (1989): Carbonate diagenesis and porosity. – Developments in Sedimentology, 46, 338 pp., Amsterdam (Elsevier) Morse, J.W., Mackenzie, F.T. (1990): Geochemistry of sedimentary carbonates. – Developments in Sedimentology, 48, 707 pp., Amsterdam (Elsevier) Veizer, J. (1983): Trace elements and isotopes in sedimentary carbonates. – In: Reeder, R.J. (ed.): Carbonates: mineralogy and chemistry. – Reviews in Mineralogy, 11, 265-299 Further reading: K005 to K012
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4 Microfacies Data: Matrix and Grains
The purpose of this chapter is to outline and discuss the criteria used in the description of matrix and grains. Microfacies criteria reflected by fabrics and quantitative microfacies data are discussed in Chap. 5 and 6. Major constituents seen in thin sections of carbonate rocks are ‘grains’ and ‘groundmass’, the latter developed as fine-grained carbonate ‘matrix’ or as sparry calcite representing pore-filling cements or recrystallized matrix. A sparry calcite groundmass is conventionally called ‘sparite’ (calcite spar), and discussed in Chap. 7. Grains and matrix are separated at 62.5 µm in terrigenous sediments but at much lower values in carbonate rocks. Here the lower limit is defined differently as larger than 4 µm or >20 µm or >50 µm. SEM studies reveal the arbitrary character of particle size boundaries between grains and matrix, but a conventional definition is necessary for thin-section studies. The physical arrangement of particles, including their texture and structure, forms a fabric which reflects the depositional and diagenetic controls on the formation of carbonate rocks.
4.1 Fine-Grained Carbonate Matrix: Micrite, Microspar, Calcisiltite Fine-grained matrix constitutes a significant part of the carbonate record. Despite the importance of microcrystalline limestones in basin analysis and as reservoir rocks and limestone resources, classifying and genetically interpretating microcrystalline carbonates still involve many problems. However, the advent of the scanning electron microscope and the data provided by geo- and biochemical analyses now allow us to take a closer look at modern carbonate muds and ancient finegrained limestones. Definition: The term matrix denotes the interstitial material between larger grains. Most definitions refer to the relatively small size of the crystals and/or particles bordering intergranular pores. Hence the matrix of micro-grained limestones includes fine-grained ma-
terial such as micrite, microspar and calcisiltite. A common generalization is that ‘matrix’ is synonymous with ‘groundmass’, but some authors use these terms to designate both fine-grained interstitial material (e.g. micrite) as well as coarse interstitial crystal fabrics (e.g. sparite formed by cementation or neomorphic processes, see Sect. 7.6). This contradicts the limitation of the term matrix to ‘mechanically deposited material between particles – as distinct from precipitated cement’ (Bathurst 1975). One should be cautious about regarding all fine-grained interstitial material as synsedimentary matrix: Intensive micritization, compaction and neomorphism of mixed carbonate/siliciclastic deposits as well as alterations of peloids may lead to the formation of a carbonate pseudomatrix (Dickinson 1970; Geslin 1994). Techniques: Genetic interpretation and the understanding of the properties of fine-grained carbonates require the examination of polished and/or etched rock surfaces, thin sections, stereoscan microscope and transmission microscope studies, the investigation of geochemical and isotopical signatures as well as the use of biomarkers (Sect. 13.3). Conventional thin sections used in microfacies studies are commonly too thick (about 30 µm) to reveal distinct crystal boundaries. Peels and ultra-thin sections offer better data (Honjo and Fischer 1965; Wolf et al. 1967; Lindholm and Dean 1973; Murphy 1986; Tucker 1988). Micrite textures described only on the basis of light microscopy may be inadequate. Epifluorescence microscopy assists in the discussion of organic material preserved in limestones (Dravis and Yurewicz 1985, Neuweiler and Reitner 1995). SEM studies of broken, polished and etched surfaces provide an insight into the variability of micrites. Polished and etched SEM samples deliver better results than broken surfaces. The main data which can be retrieved from SEM studies of fine-grained carbonates are summarized in Box 4.1. Major categories: Micro-grained limestones exhibit mean sizes commonly between 2 µm to about 30 µm.
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_4, © Springer-Verlag Berlin Heidelberg 2010
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Micrite Terminology
Box 4.1. SEM criteria used in the distinction of fine-grained microcrystalline limestones of different origin and diagenetic history (based on Fischer et al. 1967; Flügel 1968; Wise et al. 1970; Minoura 1974; Moshier 1989; Peszat 1991; Lasemi et al. 1993; Munnecke 1997). Microfabric Grain size: Is measured by counting the longest diameter of 100 to 200 grains. Allows attribution to grain size-based classifications. ‘Sorting’ in micrites, as the result of competition between space and grain growth, can be used as measure of the textural maturity of fine-crystalline carbonates. Grain size ranges may reflect differences in the depositional environment (e.g. reef micrites and micrites in limestone/shale sequences). Differences in predominant crystal sizes provide indications of the potential aragonite-dominated or calcite-dominated precursors of calcite micrites. Grain contacts: Grain contacts seen under the SEM are straight, curvilinear or serrate. The percentage of specific grain contacts reflects the degree of diagenetic changes (e.g. recrystallization). Grain shape: Differentiation of anhedral, subhedral and euhedral micrite grains. Anhedral and subhedral grains forming a hypidiotopic fabric appear to be common in micritic rocks composed of grains > 4 µm. Irregular, interlocked grains form an ‘amoeboid mosaic’, in contrast to block-like grains forming a ‘pavement mosaic’ common in limestones affected by thermal stress. Most micrite grains are angular to subangular. Subrounded and rounded grains result from intergranular solution and crystal overgrowth. Inclusions in crystals: Liquid and gas inclusion pits are common in microspar crystals as well as in larger spar crystals. The composition of the inclusions indicates the conditions under which recrystallization took place. Microfenestrae: Some micrites contain round voids, 40 to 80 µm in size, which are filled with blocky calcite. The preservation of this microfenestral fabric is used as evidence against a strong mechanical compaction of the sediment prior to lithification. Boundary between calcite matrix and fossils: Diagenetic stages are distinguished by differences in the boundaries between matrix and fossils: (a) Sharp boundaries between matrix and Low-Mg calcite fossils (e.g. brachiopods, bryozoans, ostracods, trilobites) with well-preserved ultrastructures; (b) moderately sharp boundaries between matrix and fossils (recrystallized tabulate and rugose corals, stromatoporoids); and (c) boundaries between matrix and fossils (e.g. mollusks) which are difficult to recognize. Microporosity: SEM studies reveal numerous types of micropores (5-10 µm wide) in micritic limestones, including both primary and solution-enhanced intercrystalline pores, as well as micromolds, microvugs and microchannels (Pl. 7/6; see Chap. 5). Microporosity is developed in many subsurface micritic limestones, including significant petroleum reservoirs in Mesozoic and Tertiary deposits of the Middle East, North Africa, and the North American Gulf Coast. Skeletal grains A main advantage of SEM studies of micrites is the possibility of recognizing the occurrence and abundance of planktonic nannofossils (Pl. 7/3-5, Pl. 77/23) in mudstones and wackestones in both deep-water and shallow-water settings. SEM differentiation of silt-sized and smaller bioclastic grains derived from the breakdown of shells helps to define micrites formed by the accumulation of disintegrated invertebrate skeletons (group 7: Fig. 4.1).
Three categories can be distinguished according to differences in grain size and composition: 1 Micrite, characterized by a crypto- to microcrystalline crystal texture. The original definition by Folk (1959) sets a grain-size limit of <4 µm. 2 Microspar, characterized by a fine-grained calcite matrix consisting of crystals ranging from 5 to about 30 µm in diameter, most commonly about 5 to 7 µm, and often exhibiting a rather uniform size distribution. 3 Calcisiltite (calcite silt), characterized by a finegrained matrix composed of detrital silt-sized calcite particles. Silt comprises a size grade between 2 and 62 µm (Assally et al. 1998). The upper size limit is approximately the smallest size which can be differentiated with the unaided eye.
4.1.1 Micrite Terminology: The term ‘micrite’ is the abbrevation of ‘microcystalline calcite’. Originally, ‘micrite’ was proposed as a genetic term referring to lithified mechanically deposited lime mud (Folk 1959). Most authors, however, now use the term in a non-genetic descriptive sense for a rock composed of fine-grained calcite crystals and particles formed in place or by the accumulation of fine-grained pre-existing carbonate material. This broad definition even includes ‘microcrystalline cement’ and ‘micrite cement’, precipitated as micron-sized Mg-calcite in reefs, beachrocks, hardgrounds and caliche, and occurring in primary mud-free, grainsupported sediments of carbonate slopes (Wilber and Neumann 1993). Friedman (1985) has strongly opposed the use of the term micrite for microcrystalline calcite
Classification of micrites
precipitated in marine and nonmarine diagenetic environments (see Milliman et al. 1985 for discussion). In general, micrite is understood to be the finegrained matrix of carbonate rocks and the fine-grained constituent of carbonate grains. The crystal size of micrites ranges from cryptocrystalline (not resolvable with the petrographic microscope) to microcrystalline (individual crystals recognizable with a petrographic microscope but not visible with a binocular microscope and hand-lens). The original micrite definition relies strongly on crystal size limits which can only be recognized in petrographical thin sections (Pl. 6) or the use of SEM (Pl. 7). An example of a ‘typical’ micrite corresponding to Folk’s definition are the famous Jurassic lithographic limestones of Solnhofen in southern Germany (Fig. 4.4; Pl. 46/1). Terms often used more or less synonymously with micrite are lime mud, lime ooze as well as lime mudstone (Pl. 46/1), calcimudstone and calcilutite. The last three terms are useful field terms. Microfacies thin sections are approximately 30 µm thick, and are commonly thicker than petrographical thin sections. This may cause difficulty in measuring true crystal sizes. Therefore, upper crystal size limits for micrite have come into use: 10 µm (Elf-Aquitaine 1975), 20 µm (Dunham 1962), and 30 µm (Leighton and Pendexter 1962). Bosellini (1964) differentiated ‘micrite I’ (crystal size <4 µm) and ‘micrite II’ (4– 30 µm). The broadening of the definitions has resulted in being subsumed micrite and microspar into one category. A conspicuous micrite type is minimicrite (Folk 1974), consisting of Mg-calcite and calcite crystals with a size range mostly below 1 µm (Pl. 6/6). Minimicrite appears black in transmitted light and white in direct light. Minimicrite occurs in lime mudstones, wackestones, and microbialites, and constitutes the tests of porcelaneous foraminifera. It is found in calcimicrobes (Pl. 7/6) and in the skeletons of coralline red algae (Sect. 10.2.1.2). Bernier (1994) stressed the fact that lithographic limestones composed of very fine micrite (crystals frequently as small as 2 µm) and characterized by specific controls on their origin should be clearly separated from other micrites by the designation lithographic micrite. Classification of micrites: Few topics in carbonate sedimentology have been debated as controversially as the origin and classification of fine-grained carbonate rocks. Today, fine-grained carbonate muds originate in non-marine environments (e.g. pedogenic and lacustrine settings; Sect. 2.4.1), and marine environments including shallow marine inter- and subtidal settings
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(e.g. tidal channels, algal mats, lagoons, platforms), reefs, and deep-marine ocean floors. Main hypotheses starting with Sorby (1879) explain the origin of microcrystalline calcite in limestones as (1) in-place formation of fine-grained carbonate triggered by biochemical and physicochemical factors, (2) post mortem disintegration of calcareous algae, (3) physical or biological abrasion of skeletal material, (4) accumulation of pelagic calcareous plankton, and (5) a result of diagenetic processes including cementation and recrystallization (Fig. 4.1). Honjo (1969) differentiated micrites formed by the growth of carbonate in place or resulting from the destruction, accumulation and deposition of pre-existing carbonate particles characterized by subhedral polygonal calcite grains meeting at interfaces (orthomicrites), and micrites composed of calcareous pelagic biota (nannoagorite). These primary micrites are often contrasted with the secondary micrites and pseudomicrites resulting from diagenetic processes (Wolf 1965). Several genetically defined terms are in use to designate the major modes of micrite genesis: Automicrite (Wolf 1965) refers to ‘autochthonous’ micrite interpreted as the result of in-place formation of fine-grained calcite or aragonite on the sea floor or within the sediment as an authigenic product triggered by physicochemical, microbial, photosynthetic and biochemical processes (Pl. 6/3, 5; Pl. 50/2, 4). The term was revived in the context of studies of mud mounds and biocalcification (Reitner and Neuweiler 1995, Neuweiler 1995). Automicrite exhibits specific fabrics and has a uniform mineralogical and chemical composition. Automicrite is fundamentally different from allomicrite (Wolf 1965), meaning ‘allochthonous’ micrite and interpreted as the sedimentary microcrystalline matrix derived from the breakdown of various carbonate grains and deposition of fine-grained material (Pl. 6/1, 6; Pl. 7). Allomicrites may contain mineralogically and chemically different constituents. The term ‘orthomicrite’ (Honjo 1969) designates micrites made of subhedral polygonal clay-sized calcite grains meeting at interfaces. Modern carbonate muds and ancient fine-grained carbonates: Modern carbonate muds originate in marine and non-marine sites, in warm water and cold water environments and in shallow and deep settings. Mudproducing processes which can be found in modern carbonate environments and which were also operating in the formation of ancient carbonates are: • the effects of metabolic and other processes of bacteria, cyanobacteria and algae on the precipitation of carbonates, Text continued on p. 80
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Micrite and Microspar
Plate 6 Micrite and Microspar: What do they Look Like in Thin Sections? The former carbonate mud in limestones occurs as fine-grained microcrystalline calcite matrix (micrite). The originally very strict definition (crystals < 4 µm) was broadened for practical reasons and because of the difficulties in genetically interpretating micrite. Micrite represents an accumulation of allochthonous sedimentary particles forming allomicrites (e.g. calcareous nannoplankton, foraminifera or parts of algal skeletons; –> 1, 6 and Pl. 7) or a product formed in place on the sea floor (–> 5) or within the sediment (–> 3, 4) by physicochemical, microbial and biochemical processes (automicrite). Regardless of an autochthonous or allochthonous origin, the occurrence of micrite is commonly regarded as an indicator of low-energy depositional environments in protected shallow areas or in deep water below the fair-weather wave base. Many problems are involved with this paradigma. The term microspar (microsparite) refers to a calcite matrix with rather uniform crystal size and equant crystal shape. The origin is explained by the recrystallization of micrite-sized calcite crystals during recrystallization or as one-step process of cementation and calcitization of aragonite precursors (see Fig. 4.5). There exist six main criteria which should be considered in a thin-section analysis of fine-grained limestones and which are answered in the description of the figures of this plate: (A) - Sea-floor micrite or internal micrite? (B) - Overall fabric? (C) - Textural differences in the sample? (D) - Crystal size and crystal fabric (E) - Fossils and biogenic textures? (F) - Non-carbonate material? 1 Micrite types: Note the textural difference between the rounded micrite clast, composed of homogeneous micrite, and the inhomogenous finely bioclastic ‘allomicrite’ matrix (questions B and C). Bioclasts are shell debris (question E). The sample represents an intraclast wackestone (see Pl. 16/3) formed by episodic storms on a carbonate ramp (question A). See Pl. 102 for further criteria for storm deposits. The crystal sizes of the micrite clasts are about 10 µm and exceed the 4 µm limit (question D). Middle Triassic (Hauptmuschelkalk): Künzelsau, southwestern Germany. 2 Microsparite: Equant fine-grained texture interpreted to have originated by recrystallization in the vadose zone of former micrite (see Pl. 47/2). Late Jurassic carbonate ramp (see Pl. 136): Subsurface, Kinsau, southern Germany. 3 Automicrite: Cauliflower structure formed within cavities of a carbonate mud mound. The inhomogenous structure (question B) represents a succession of ‘peloidal container organomicrites’ (black, crystal size about 10 µm, question D) and accretionary thrombolites (irregular peloidal fabric formed within relics of soft sponges; see questions A and E). Thrombolite structures are also shown on Pl. 50/5 and Pl. 131/5. Early Cretaceous (Albian): Soba reef area, Mazo Grande, northern Spain. 4 Internal micrite (question A) within synsedimentary cavities in a micritic peloidal bindstone. Note gradation, geopetal filling and coarse sparite at the top. The cavities occur within dark micrite formed by sediment entrapment by tiny donezellid algae (arrows; see Pl. 60/1; Question E) which contribute to the framework of decameter-sized mud mounds. Micrite infillings of synsedimentary cavities often exhibit textures which are distinctly different from the enclosing matrix (Pl. 17/3, Pl. 20/1, Pl. 50/4). Significance of the sample: Evidence of the timing of sedimentary and diagenetic processes. The internally precipitated micrite alters the mud/grain ratio of the pre-existing sedimentary deposits and questions the use of micrite content an indicator of low water energy. Late Carboniferous (San Emiliano Formation, Westfalian): Cantabrian Mountains, northern Spain. 5 Sea-floor automicrite (question A). Hemispheroids of dense micritic/fenestral microbialites (question B) contributing to the formation of decameter-sized platform-margin reefs (Neuweiler 1993). The finely laminated structures disappear with increasing sediment supply. Autochthonous formation of microbial micrite (originally Mg-calcite) by in-situ calcification is indicated by the purity of the micrite (no non-carbonate material; question F) and the protrusional outer margins (arrows). Slight changes in sediment input are shown by the alternation of black micrite and micrite rich in fine bioclastic debris (predominantly sponge spicula; question E). Note circular cross section of burrows pointing to a rapid alternation of soft and hard bottom conditions (question E). Early Cretaceous (mid-Albian): Soba reef area, Vasco-Cantabrian basin, northern Spain. 6 Pelagic sea floor allomicrite with abundant planktonic foraminifera (questions A and E). Note differences in size and packing of the calcitic foraminiferal tests which may point to burrowing in the sediment. The sediment is dominated by foraminiferal tests most are sand-sized. The fine-grained micrite component often results from the deposition of nannofossils and only to a minor degree from the degradation of foraminiferal tests. SMF 3-FOR. Open-marine deep-water environment. Early Tertiary (Eocene): Apennines, Italy. –> 3: Reitner and Neuweiler 1995; 4: Hensen et al. 1995; 5: Neuweiler 1993
Plate 6: Micrite and Microspar
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Micrite: SEM Studies
Plate 7 A Closer Look at Micrites: Genetic Interpretations Based on SEM Studies SEM criteria (Box 4.1) are useful in classifying allomicrites according to crystal size, biotic composition and textural feature and assist greatly in interpretating the possible origin of fine-grained limestones. Many modern carbonate muds and ancient micrites are accumulations of disintegrated calcareous skeletal elements. Two models focus on different aspects of this: The ‘Halimeda model’ (Fig. 4.3) focuses on a bioclastic origin of the fine-grained matrix of carbonate rocks. Calcareous green algae (–> 1-2) are regarded as a major potential source of carbonate mud originating in tropical and subtropical shallow-marine warm-water environments (e.g. Bahamas, Florida). Muds are deposited in protected parts of platforms and are exported to periplatform areas. The skeleton of Halimeda (Pl. 57/1, 2, Pl. 149/ 4) and other lime-secreting green algae consist of minute aragonite needles which are set free on the death of the plants. These needles (–> 2) are similar in morphology, size and geochemical signatures to needles found in the aragonite muds. Phanerozoic halimedacean green algae became abundant during the Mesozoic (Pl. 57/3, 6) but were already common elements in Devonian platform carbonates (Pl. 57/4, 5). The ‘Solnhofen-Oberalm model’ (Fig. 4.4) underlines the rock-building capacity of Low-Mg calcite skeletons of planktic algae (mostly Coccolithophorida, –> 3-5) living in the upper hundred meters of the water column of the open oceans, but also occurring in shallow lagoonal waters. The accumulation of these nannoorganisms resulted in the formation of deep-marine basinal carbonates since the Mesozoic. Micrites composed predominantly of the constituent shields of coccolithophorids and related groups appear in the Triassic, are common in the Jurassic and became very abundant during the Cretaceous (e.g. in chalks, –> 3) and the Tertiary. Many taxa are useful for biostratigraphic correlations of Late Cretaceous and Cenozoic open-marine pelagic carbonates. With some exceptions (e.g. Schizosphaerella, –> 7) these ‘nannofossils’ are generally too small to be recognized in regular thin sections and must be studied in SEM. Note that the samples of Figs. 4 and 5 on this plate would be easily classified as ‘lime mudstone’ or ‘unfossiliferous micrite’ if studied in thin sections only (Pl. 44/1, Pl. 46/1). SEM images on the other hand define a characteristic packstone fabric. Extremely fine-grained ‘mini-micrites’ are common features of microbially induced automicrites and of micritic Mg-calcite skeletons of cyanobacteria (–> 6) and some foraminifera. 1 The udoteacean green alga Halimeda is one of the most effective carbonate producers in modern tropical and subtropical reefs and lagoons (mode 5 of Fig. 4.1); cf. Fig. 4.2. You are looking into the organic filament here. Recent: Bahamas. 2 Close-up of fused Halimeda aragonite needles. Algal needles occurs as blunt-ended laths which are larger than precipitated crystals. Recent: Bahamas. 3 Coccolith chalk (nannomicrite; mode 9 of Fig. 4.1). Chalk is a pelagic sediment composed predominantly of calcitic fossils including coccoliths, with many planktonic foraminifera and calcispheres (Zijlstra 1995). Important macrofossils are ammonites, belemnites and crinoids. A significant benthic fauna (echinoids, bivalves, bryozoans, sponges, brachiopods) occurs in shelf-sea chalks. The depositional depth of chalks varied between a hundred meters to a few hundred meters for shelf-sea chalks, but to many hundred meters and perhaps thousands of meters for deep-sea chalk. The sediment consists of diagenetically altered coccolith plates, smaller micrite grains and large euhedral calcite crystals resulting from pressure solution overgrowth. Coccolith shields are deposited parallel to the bedding. Diagenesis is variable. Much of the chalk is only weakly lithified (soft chalk) and retains a high porosity in spite of deep burial. The low cementation potential of shelf chalks is due to the primary low calcite mineralogy of most skeletal grains, the lack of meteoric pore fluids and the absence of deep burial. Retarded diagenetic change in chalks, also caused by large amounts of clay minerals, is responsible for high reservoir qualities. Late Cretaceous: DSDP Site 463, core 44, western equatorial Pacific. 4 Coccosphere. The preservation of coccospheres is a proxy for bottom water energy conditions. Intact coccospheres are evidence of minimal post mortem disturbance. Note overgrowth of small calcite crystals on the coccoliths. Close-up of –> 5. 5 Coccolith limestone. Note the co-occurrence of well preserved coccospheres (bottom) and isolated and broken coccoliths (right). Cretaceous (Maastrichtian): Denmark. 6 Microbial micrite, characterized by extremely small equant calcite crystals. These minimicrites were formed within the peripheral mucus of cyanobacteria. The circular open structures correspond in size and distribution to structures interpreted in the Solnhofen limestone as coccoid cyanobacteria (Keupp 1978). Small crystal size as well as the absence of pitted surfaces and inclusions suggest a High-Mg calcite precursor mineralogy. Note the high intercrystalline microporosity. Late Jurassic (Tithonian): Kapfelberg/Kelheim, Bavaria, Germany. 7 Deep-water limestone consisting of Schizosphaerella (S). This calcidinophycean algal genus is a common rock-building constituent of Tethyan Jurassic hemipelagic basinal limestones. The tiny globular fossil is constructed of two hemispherical to cap-shaped valves. Early Jurassic (Sinemurian): Adnet near Salzburg, Austria. –> 3: Courtesy A.Wetzel; 7: Böhm 1992
Plate 7: Micrite: SEM Studies
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• the influence of microbial mats on the formation of autochthonous fine-grained carbonates, • the processes involved in the breakdown of the skeletons of calcareous algae and invertebrates into finegrained particles, • the role of sessile organisms in the control of the deposition of fine-grained material (trapping, baffling, binding), • the variations in burrowing activities creating specific carbonate mud microfabrics and • the ways in which the deposition, dissolution and accumulation of calcareous plankton takes place. The overwhelming majority of modern shallow- and deep-marine carbonate muds is now regarded as biogenic or biologically induced in origin. Initially, microcrystalline carbonate matrix was considered as precipitated directly from seawater (Cloud 1962). Later, spurred by observations of Lowenstam (1955) and Lowenstam and Epstein (1957) and supported by the forward-looking suggestions of Sorby (1879), most workers began to accept shallow-marine lime muds as being primarily the product of a breakdown of skeletal biota, particularly calcareous algae. ‘Whitings’ (floating patches of lime mud), as evidence of physicochemical precipitation, fell into disrepute and were considered to be caused by biologically induced carbonate precipitation (Robbins and Blackwelder 1992). But, more recently, evidence for the direct precipitation of aragonite lime muds emerged (Shinn et al. 1989; Milliman et al. 1993; Friedman 1994). The increase in micrite carbonates during the Early Proterozoic is thought to be a result of blooms of cyanobacteria in phosphate-rich oceans. Strong photosynthesis combined with high population densities could have induced the precipitation of mud via whitings. The finegrained sediment produced in whitings was fixed in benthic cyanobacterial mats. In waters, highly oversaturated with respect to calcium carbonate, the mats develope to stromatolites (Merz-Preiss 2000). Very recent investigations suggest an intricate bio-physicochemical coupling between microbial activity, water current circulations, and water chemistry (Thompson 2000). The aragonitic mud on the Great Bahama Banks has long been of great interest as a possible analogue for ancient fine-grained carbonate rocks. As a result, the term lime mud has become more or less synonymous with aragonite needles in the minds of many geologists. It must be stressed, however, that many modern carbonate muds consist of Mg-calcite (e.g. South Florida and Belize: Taft and Harbaugh 1964, Enos 1977) and need an alternative genetic interpretation. Mic-
Formation of Micrite
ritized skeletal grains are potentially a major source of Mg-calcite mud (Reid et al. 1992). Unlithified deep-sea carbonates formed at the sea bottom were reported from the Mediterranean Sea and the Red Sea. Supersaturation with respect to calcium carbonate due to relatively high water temperatures and salinities promotes ‘inorganic’ precipitation of Mg-calcite carbonates (Milliman et al. 1969; Milliman and Müller 1977). Another hypersaline basin where inorganically precipitated carbonates have been reported is the Arabian Gulf. The large quantities of fine-grained aragonite in nearshore sediments off Saudi Arabia and the Trucial Coast are explained by inorganic precipitation, since aragonite-producing calcareous green algae do not grow in the Arabian Gulf (Kinsman 1964). However, no evidence for inorganic precipitation of deep-sea Mg-calcite carbonate muds was found by Ellis and Milliman (1985). These muds are formed by suspended carbonate grains of biogenic or detrital origin.
4.1.2 Modes of Formation of Micrite and Other Fine-Grained Matrix Types The discrimination of micrites generated by different processes is of high-ranking importance for paleoenvironmental interpretations, basin analysis and the understanding of the properties of fine-grained carbonate rocks. Fig. 4.1 summarizes twelve modes which can lead to the formation of micrites. ‘Inorganic’ micrites: (1) Physicochemical precipitation: Initially, the microcrystalline matrix of limestones was considered a ‘chemical’ rock, precipitated directly from seawater. This view was based on the interpretation of carbonate muds in hypersaline lagoons of the Bahamas as an exclusively physicochemical precipitation of Mg-calcite and aragonite in conjunction with high salinity and extreme temperature (Cloud 1962, De Groot 1965, Loreau 1982, Dix 2001), but was also contested by geochemical data (Husseini and Matthews 1972). Calculations of sedimentation rates, carbonate mud budget and stable isotope data indicate that a considerable amount of the aragonite and perhaps also Mg-calcite muds of the Great Bahama Bank, observed as ‘whitings’ could be chemically precipitated (Shinn et al. 1989), but a biochemically induced precipitation triggered by planktonic cyanobacteria cannot be ruled out. Whitings believed to have been triggered by plankton blooms are reported from marine shallow-water areas as well as from lakes. Inorganic precipitation of Mg-
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Automicrites
Processes Abiogenic (‘inorganic’) Automicrite (autochthonous micrite) formed in place at the seabottom or within the sediment
Biologically induced
1 Physicochemical precipitation triggered by salinity and water temperature fluctuations 2 Carbonate precipitation mediated by organic matrices (Ca-binding organic macromolecules), causing organomineralization and formation of organomicrite 3 Metabolic processes of heterotroph and chemolithotroph bacteria and other microbes causing microenvironmental changes which induce carbonate precipitation
Biologically controlled
4 Metabolic processes of phototrophic cyanobacteria and algae causing carbonate precipitation 5 Disintegration of benthic calcareous algae into submicroscopic fragments (Halimeda model)
Allomicrite (allochthonous micrite), deposition of disintegrated skeletal material and of fine erosional detritus
Disintegration of predominantly benthic biota
7 Disintegration of invertebrate skeletons 8 Bioerosion causing detrital abrasion and microborings causing ‘micritization’
Disintegration of pelagic biota Erosion and abrasion
Pseudomicrite Diagenetic ‘micrite’
6 Disintegration of epibionts living on seagrass and macro-algae
9 Accumulation of calcareous plankton (foraminifera; coccolithophorids and other nannofossils causing ‘nannomicrites’) 10 Mechanical erosion of limestones, e.g. at coasts 11 Micro- and cryptocrystalline carbonate cements
Diagenesis
12 Recrystallization and ‘grain diminution’ (replacement of former larger crystals by tiny crystals)
Fig. 4.1. Terminology and genetic modes of origin of carbonate muds and limestones composed of microcrystalline calcite (micrite). The numbers correspond to the numbers used in the text.
calcite muds has been suggested for carbonate muds along the Trucial Coast and in the Dead Sea. Interpretation of limestones: Various authors studying ancient carbonate rocks have assumed that carbonate mud is inorganically precipitated. This is mainly based on a facies association comparable to the Bahama Bank or related to the grain sizes of micritic limestones, similar to those of modern aragonite crystals. Automicrites Automicrites (autochthonous micrites) are microcrystalline carbonates formed at the sea floor or within the sediment. As a primary product of in-place mineralization, automicrite is fundamentally different from allomicrites which are a microcrystalline sedimentary matrix formed by the deposition of disintegrated skeletal material and fine erosional detritus. Automicrites include biologically induced, biologically controlled as well as ‘inorganic’ carbonates as discussed above. The formation of autochthonous micrites is triggered by: • Organic matrices (Ca-binding organic macromol-
ecules) mediating organomineralization (Reitner 1993, Reitner et al. 1995), and causing the formation of ‘organomicrites’. Carbonate mineralization concepts based on the interaction with Acidic Organic Macromolecules (AOM) first used to explain the skeletal structures of organisms (Degens 1979; Addadi and Weiner 1989; Lowenstam and Weiner 1989; Simkiss and Wilbur 1989), and were further developed in the context of automicrite formation (Reitner et al. 2000). • Metabolic processes of heterotroph and chemolithotroph bacteria and other microbes causing changes in the microenvironment surrounding the bacterium and inducing calcium carbonate precipitation. Bacterial controls on carbonate precipitation were at one time controversially discussed but have now been confirmed by experimental studies as well as direct field observations (Sect. 9.1). • Metabolic processes of phototroph cyanobacteria and algae, that reduce CO2 to organic carbon compounds, and thus shift the solubility equilibrium toward the precipitation of carbonate. This mode of automicrite
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Lizard Island Model
>>> Fig. 4.2. The Lizard Island model (Reitner 1993, Reitner et al. 1995). Lizard Island is located in the northern part of the Great Barrier Reef. The granitic island is fringed by coral reefs. The microbialites growing in cryptic habitats of dark reef caves offer a model for describing the biological controls on the formation of automicrites. The microbialites are products of matrix-mediated calcification via specific Ca-binding acid organic macromolecules and baffled detritus. The model shown in the figure underlines the interaction between biofilm activity, sponge development, and the development of microbialites. Microbes and sponges played a central role in the formation of ancient automicrites from the beginning of the Cambrian (Pl. 8/7, Pl. 82). The interaction between sponges and microbial biofilms is not restricted to cryptic niches in tropical reefs, but also occurs in cool-water and polar sponge mounds. A: The surface of the microbialites is covered by microbial biofilms (1) which control the input of Dissolved Organic Matter and of Dissolved Organic Carbon. Biofilms are complex structures consisting of variably diverse microbial single- or multilayered sheets (see part B). The biofilms control the distribution of sponges (2) and, via biochemical signals, the settlement of sponge larvae (3). The sponge profits from the supply of microbial nutrients and DOM from the water column (4) and DOC from the biofilm. Metabolic waste (5) is recycled. The products of symbiotic heterotrophic bacteria living within the sponges are important sources of nutrients; they also control metabolic processes and enhance calcification. Microbial biofilms are responsible for the degradation of organic matter into Ca-binding macromolecules, and therefore for the production of free calcifying mucilages. Calcified microbes and partly mineralized fungal hyphae have been observed in some biofilms (6). Fe/Mn bacteria contribute to the formation of Fe/Mn crusts (7) which are also common in many ancient automicrites. The sediment-trapping property of biofilms is important for the growth of microbialites consisting of thrombolitic crusts forming irregular pillar-like structures at the top of the microbialite (Pl. 8/1). These crusts, dome-shaped micritic and irregular peloidal textures are major constituents of ancient mud mounds (Pl. 8/2, 6). Calcium crystal growth requires increase in alkalinity. In the Lizard Island reef caves increase in alkalinity is triggered by the weathering of granitic silicates which form the island. Microbialites grow at the surface (8) or in erosional pockets (9) forming ‘accretionary organomicrites’. ‘Container organomicrites’ are represented by the carbonates precipitated within the decaying sponge. The vertical sequence (10) starts with a coralgal facies followed by a coralline algal crust and the microbialite facies. This succession corresponds to the lightdependent horizontal succession seen in the caves. B: Biofilm model after Wilderer and Charaklis (1989) adapt to microbialite formation. The biofilms are about 100 µm thick. They act as a reactor to take up dissolved organic carbon and organic particles via EPS channel systems. The reactor filters and produces DOM and DOC, e.g. as waste products, and release the organic matter back to the seawater or into the open system of microcavities and irregular pocket structures of the upper microbialite. Some of the products are acidic macromolecules which enhance mineralization and the formation of automicrite. DOC - Dissolved Organic Carbon, DOM - Dissolved Organic Matter, EPS - Extracellular Polymeric Substances, IPS Intracellular Polymeric Substances. Modified from Reitner (1993).
formation has been discussed for some time, especially in regard to micrites formed in peritidal ‘algal mats’. The formation of organomicrites is controlled by organic macromolecules (Fig. 4.2). The organic matter is characterized by a surplus of specific compounds derived from microbes, free organic matter, and decaying organisms. Evidence of Ca-binding Acidic Organic Macromolecules (AOM) have been found in modern and ancient automicrites by chromatography and fluorescence microscopy (Neuweiler and Reitner 1995, Russo et al. 1997). Fluorescence microcopy is of particular value in the determination of residual organic matter (Dravies and Yurewicz 1985, Machel et al. 1991). Two sources of Acidic Organic Macromolecules can be distinguished: (1) Internal production of AOM by eucaryotic and procaryotic organisms via metabolic enzymatic and decaying processes, (2) External origin of AOM as a part of dissolved organic matter in the
open water column. The macromolecules become enriched within immature microbialites or at the interspace water/mature sediments. They are produced and/ or fixed by organisms (e.g. in biofilms, Fig. 4.2B) or trapped in microcavities and pockets. Differences in the composition of these compounds may result in different carbonate fabrics: Enrichment in acidic amino acids (Asp + Glu) presumably is linked to the formation of stromatolite and thrombolite fabrics; proteinrich substances within confined spaces (e.g. microcavities) result in the formation of peloidal pockets, peloidal coatings and peloidal stromatolites; and the decay of soft sponge tissue, enriched in bacteria, leads to micropeloidal structures mirroring parts of the former sponge bodies. The characteristic criteria of automicrites are biolaminated structures, clotted peloidal microfabrics or extremely fine-grained cryptocrystalline textures. A common automicrite microfacies of automicrites formed in slope settings are boundstones composed of small peloids that are associated with cement layers
Lizard Island Model
(Keim and Schlager 1999). Modern automicrites as well as peloidal and microcrystalline carbonate cements consist of Mg-calcite and aragonite. The same mineralogies have been inferred for ancient automicrites but a Low-Mg calcite composition might be a possibility, too. (2) Organomicrites (Reitner et al. 1995; Pl. 8): The formation of modern Mg-calcite automicrites precipitated in isotopic equilibrium with the ambient seawater, and controlled by a specific composition of organic material was studied by Reitner (1993) in indurated microbialites occurring in reef caves off Lizard Island, Great Barrier Reef (Fig. 4.2; Pl. 8/1). The microbialites are formed under dark conditions. Phototrophic organisms are absent. The most important factor controlling the formation of organomicrite is the presence of large
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amounts of specific organic macromolecules with the ability to fix and bind divalent cations, mainly Ca2+ and Mg2+. The increase in alkalinity required for calcium crystals is triggered by ammonification (microbial decay of organic matter, e.g. sponges), sulfate reduction (in anaerobic layers of stratified water masses) or via carbonate and silicate weathering (Kempe et al. 1989; Reitner and Neuweiler 1995). Neuweiler et al. (2000) discuss the controls of soluble humid substances on the in situ precipitation of microcrystalline calcium carbonates. Organomicrites are differentiated according to the modes and loci of their formation (Reitner 1993): • Accretionary organomicrites (Pl. 6/5, Pl. 8/1, 2; Fig. 4.2) are vertically growing structures exhibiting thrombolitic, stromatolitic or massive hardground-type
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textures. Borings are proof of early lithification. The textures are a function of sediment supply. Very finegrained automicrite (0.5–1 µm; minimicrite) is formed by calcifying microbial biofilms within thin calcifying mucus sheets below the biofilms, and in pockets of trapped organic mucus and sedimentary grains. These conspicuously dark, aphanitic micrites contain in-situ peloids and in-situ ooids and are characterized by stromatolitic or/and massive fabrics. • Container organomicrites (Pl. 6/3, 4, Pl. 50/4) develop as fill-up structures of semi-closed microcavities within encapsulated decaying tissues of metazoans, mostly sponges, as well as in voids made by boring organisms. Many sponges represent ‘bacterial containers’ (Fig. 4.2) because the bacterial biomass within the sponge is higher than that of the sponge tissue itself. The decay of bacteria-rich tissue result in the formation of a clotted micropeloidal microfabric which can also be observed in ancient sponge reef carbonates. (3) Microbial precipitation controlled by heterotroph and chemolithotroph organisms (bacteria, fungi): Bacterial decay of organic matter is probably the main source of energy in the early diagenesis of sediments (Pl. 50/1, 2, 3). Bacteria are important agents in the formation of ferric hydroxide, iron sulfides, calcium phosphates as well as calcium carbonate. Starting more than ninety years ago (Drew 1911; Bavendamm 1932), an increasing number of authors have demonstrated and suggested that life processes of marine bacteria and the decomposition of organic matter by bacteria cause physicochemical changes in the microenvironment that result in calcium carbonate precipitation. This has been proved in laboratory experiments and observed in various modern carbonate-producing environments (soils, freshwater and marine realms, particularly in lagoons; see Buczynski and Chafetz 1993; Castanier et al. 1999, and Sect. 9.1). Interpretation of limestone (Box 4.2): A bacterial (not a cyanobacterial) origin of ancient micrites has been argued by many authors (Salomon 1914; Hadding 1958; Maurin and Noel 1977; Riding and Awramik 2000). Because most bacteria are indifferent to light, bacterially-controlled precipitation of automicrite is not restricted to shallow environments, but also is possible in subtidal settings, cryptic habitats as well as deep restricted basins. A microbial production of the micrite forming the bulk of ‘mud mounds’ is advocated by many authors (Sect. 9.1.5.2; Pl. 8/6). The assumed position of the Paleozoic mounds in deep-water environments below the photic zone creates some problems for this interpretation (Bourque and Boulvain 1993) except the bacteria involved are regarded as chemo-
Formation of Micrites
synthetic (Warnke and Meischner 1995). Microbial micrites were described from Tertiary reefs (Saint Martin et al. 1993). (4) Precipitation controlled by phototroph organisms (cyanobacteria and algae): The removal of CO2 by assimilating cyanobacteria and algae, linked to photosynthesis and augmented by bacterial activity, triggers the precipitation of Mg-calcite and aragonite. The production of lime mud by cyanobacteria is known in modern peritidal environments (Bahamas, Florida Bay, Persian Gulf, Red Sea sabkhas) and in highly saline lakes (e.g. Coorong, Australia; Lake Tanganyika, eastern Africa; Antarctic oasis lakes). The importance of filamentous cyanobacteria in the formation of micrite is documented by the calcification and preservation potential of benthic cyanobacteria (Golubic et al. 2000; Merz-Preiss 2000). The LowMg calcite crystals produced within the extracellular organic sheets of benthic cyanobacteria are set free after the death of the cyanobacteria and the degradation of the organic material; this may result in the formation of carbonate mud. This model is applied to ancient micritic rocks, e.g. to the Late Triassic Hauptdolomit of the Northern Alps which represents a huge platform connected with the hinterland and affected by runoffs of freshwater (Zankl and Merz 1994). Kazmierczak et al. (1996) and Kazmierczak and Iryu (1999) postulate that many Jurassic micritic and peloidal limestones are products of a variously intensive calcification of mats of benthic coccoid cyanobacteria, and that many microcrystalline cements described from ancient reefs may be products of in vivo calcified cyanobacteria. Cyanobacterial micrites are often developed as minimicrites with crystal sizes <2 µm (Pl. 7/6). The in situ calcification of benthic coccoid cyanobacteria mats, observed, for example in the highly alkaline soda lake Lake Van, Turkey, may represent a model for cyanobacterial controlled formation of micrite and peloidal micrite (Kempe et al. 1991; Kazmierczak 1996). Coccoid cyanobacteria mats are permineralized with micritic aragonite precipitating in vivo on and within the mucilage sheets surrounding the cells. Variations in the intensity of the calcification of the mats controlled by the level of calcium carbonate saturation create homogeneous micrite and peloidlike bodies, producing a pelmicrite texture. Disintegration into individual peloids and reworking by currents result in a pelsparite texture. Micritic peloidal grains reminiscent of those of the hypersaline Lake Van are common in cyanobacterial mats formed in hypersaline perimarine environments (Baffin Bay, Texas: Dalrymple 1965; Laguna Mormona, California: Horodyski Text continued on p. 88
Microbial Carbonates
85
Box 4.2. Microbial Controls on Carbonates. Interest in the bacterial fossil record is expanding owing to the recognition of the fundamental microbial controls on sedimentation and Earth’s biogeochemical evolution (Neuweiler et al. 1997, Ehrlich 1999). Bacterial sulfate reduction plays a paramount role in the formation and decomposition of minerals, and the control of early diagenetic processes in carbonates (Machel 2001). Bacteria and other microbes are able to form carbonate rocks (limestones: Castanier et al. 1999; dolomite: Wright 1999) and minerals on all scales. The significant role of microbes in the formation and degradation of shallow- and deep-marine as well as non-marine carbonates becomes increasingly more evident (Riding and Awramik 2000). Microbes (cyanobacteria, photosynthetic and non-photosynthetic bacteria, green and red algae, fungi) and ‘microbialites’ (Burne and Moore 1987) play a crucial role in the genetic interpretation of micrite, carbonate cements and biomineralization processes. Microbialites are laminated or nonlaminated organosedimentary deposits formed under the control of benthic microbial biofilms, organic matter degrading microorganisms, benthic metazoan communities, free organic matter, and sediment trapping. The formation and construction of inter- and subtidal microbial mats, intertidal and subtidal stromatolites, mud mounds and reefs, and vent- and seep carbonates is highly controlled by microbes. Marine microbialites are common structures in cryptic niches of modern and ancient reefs (see Facies, vol. 29, 1993). Mud mounds are regarded as a specific carbonate factory that differs fundamentally from the tropical and from the cool-water carbonate factories (Schlager 2000). The geologic record of the microbial contribution to framebuilding in deep-water mud mounds is inferred from (1) problematic microfossils, (2) laminated structures, and (3) clotted and fenestral micrite (thromboids). Bacterial calcification occurs in pedogenic (e.g. caliche: Loisy et al. 1999), sabkha environments, cave carbonates and lacustrine carbonates, and in hot-water travertines and beachrocks. Microbial activity has a high impact on diagenesis (Sedimentary Geology, vol. 126/1-4, 1999), controls lithification of tidal deposits and reefs and the formation of cements (Gonzales-Munoz et al. 2000), and contributes to the formation of manganese nodules, concretions (Coleman 1993) and iron sulfide and phosphorite deposits. Fe-encrusted bacteria and fungi produced the red color of Devonian mud mounds (Boulvain et al. 2001) and hematite staining of carbonate grains and microstromatolites in Jurassic carbonate ramps (Préat et al. 2000). Most heterotrophic bacteria are able to induce calcium carbonate precipitation as a by-product of physiological activity in laboratory experiments (Buszynski and Chavetz 1993). Bacteria act as nucleation sites and may become embedded in growing carbonate particles (Castanier et al. 2000). The calcification in procaryotic cyanobacteria depends on water chemistry and photosynthetic bicarbonate uptake, and the existence of suitable sheets. Physicochemical precipitation at high levels of supersaturation leads to a dense encrustation of the filaments, forming solid micritic tubes (Pl. 53/2). Photosynthetic bicarbonate take-up in less saturated waters can lead to micrite precipitation within the sheets (Merz-Preiss 2000). Globular grains and peloids occurring in modern microbial mats and ancient stromatolites (Pl. 123/ 3) are formed by bacterially induced calcium carbonate precipitation around dead cyanobacterial threads as proved by laboratory experiments. Microbial mats are vertically laminated organosedimentary structures, which develop on solid surfaces and have steep gradients of oxygen and sulfur (Van Gemerden 1993). They are dominated by cyanobacteria, colorless and purple sulfur bacteria, and sulfate-reducing bacteria. The driving force of most microbial mats is photosynthesis by cyanobacteria and algae. Microbial signatures in thin sections are in-situ formed benthic peloids and ooids (Chavetz 1986, Gerdes et al. 1994) within mats, thrombolitic, clotted and fine-peloidal fabrics, laminated and stromatolitic fabrics, fossils interpreted as cyanobacteria (see Pl. 53) and microborings (Pl. 52/1-4, 6). Thrombolites (Kennard and James 1986; Pl. 6/3, Pl. 50/5, Pl. 131/5) characterized by clotted microtexture and the absence of lamination occur in a variety of settings ranging from freshwater environments to deep shelf, and were important in the formation of framework reefs from the Cambrian to the Cretaceous (Leinfelder and Schmid 2000; Pl. 50/5). SEM observations may exhibit nannobacteria, and organic biomarker and isotope data assist greatly in the evaluation of microbial contributions to the formation of carbonate rocks. Nannobacteria in the 0.03 to 0.2 µm range occur in modern and ancient carbonates. Folk and Chavetz (2000) stress the importance of these microbes for carbonate precipitation in non-marine and marine environments, but also say that ‘one needs faith and optimism’ to see these extremely tiny grains. Southam and Donald (1999) argue that these very small spherical or rod-shaped grains may represent solid, inorganic precipitates.The geological record of calcified bacterial mats and microbial carbonates starts in the Archaean. Microbial mats are the oldest structured ecosystems on earth, known already from 3.5 Ga old carbonates (Sumner 2000). The predominance of microbes in Precambrian ecosystems has no counterpart in the Phanerozoic.
86
Microbial Carbonates
Plate 8 Microbial Contribution to the Formation of Micrites: Some Examples 1 Modern microbialite (accretionary organomicrite) forming cm-sized thrombolitic columns without distinct laminations. The structure consists of microcrystalline and peloidal Mg-calcite. Note the numerous inclusions of detrital grains (arrows). These automicrites exhibit specific soluble organic matrices which are regarded as responsible for the nucleation of calcite crystals. Bacteria within biofilms contribute to the degradation of organic matter (see Fig. 4.2). Reef cave, Lizard Island, Great Barrier Reef. 2 Bacterial contribution to the formation of mud mounds: Crudely laminated stromatolite interpreted as organic matrix mediated automicrite (organomicrite, see Figs. 4.1 and 4.2). Note the morphological similarity with –> 1. Settlement by sponges, foraminifera and polychaetes (not shown in the figure) and the criteria discussed on Pl. 6/5 indicate calcification at the surface. Mud mound. Cretaceous (Albian): La Sia, Vasco-Cantabrian basin, Spain. 3 Bacterial contribution to the formation of non-marine oncoids: Cyanobacteria contribute significantly to the formation of oncoids in non-marine, transitional and marine environments (see Pl. 12/1, 5). The taxonomic composition of the oncoids reflects differences in environmental conditions (e.g. salinity). The figure displays details of the cortex of an oncoid formed in a freshwater environment (Leinfelder 1985). The vertical sequence of the cyanobacterial morphotypes tells us the environmental history of the oncoid: a - Bacinella? sterni Radoicic (alga?, intergrowth of alga and nubeculariid foraminifera?) interwoven with cyanobacteria; b - Scytonema/Calothrix; c - crust consisting of Phormidium; d - bush-like tufts Schizothrix; e - another Phormidium crust. The oncoid-forming organisms can be compared with morphotypes of modern calcifying cyanobacteria. Changes in the morphotypes reflect response to changing microenvironments. The Scytonema/Calothrix morphotype points to an episodic toleration of subaerial exposure, the Phormidium morphotype indicates increased hardwater conditions of moderate to high energy levels, and the Schizothrix type represents fairly agitated freshwater conditions. The oncolite horizon occurs within terrestrial siliciclastica and is interpreted as deposits of an ephemeral hardwater lake with varying salinities. Late Jurassic: Albuquer, Portugal. 4 Bacterial contribution to the formation of micritic deep-water limestones: Dendrolite microbialite. Erect or pendant micritic mm-sized calcitic shrubs occurring in cryptic habitats of aphotic environments (Frutexites Maslov; arrow), and interpreted as being bacterial or fungal in origin, are often present in ferromanganese crusts of ancient deep-water stromatolites (Myrow and Coniglio 1991; Böhm and Brachert 1993). Frutexites is an important element of deep-water and peritidal carbonates formed during intervals of low sedimentation. Shallow-water examples appear to represent coelobites, deep-water examples grew within the sediment or on the sea floor of dysaerobic basins. Note the penetration of the sediment and the downward growth (indicated by the internal infill) of the shrubs. Note the change in shape of the more globular fibrous Frutexites growing into the cavity as compared with the black opaque shrubs within the sediment. The arrow points to the erosional boundary of Frutexites on the cavity ceiling. Morphologically similar bacterial shrubs form in a variety of inhospitable marine and non-marine environments (Chavetz and Guidry 1999), including travertines. An inorganic origin of some of these structures, however, can not be ruled out. Early or Middle Jurassic (Obersee Breccia): Lunz, Austria. 5 Bacterial contribution signalized by ‘black/white’ peloids: Tiny micritic peloids (arrows) surrounded by a clear rim of acicular calcite cement crystals are interpreted as bacterially induced or as microbial fossils (e.g. Baccanella, see Pl. 99/ 8). Similar peloids have been described from numerous shallow marine reefs, both modern and ancient (Macintyre 1985, Chavetz 1986). In-situ ‘benthic’ peloids occur as cavity-fill precipitates or as laminated crusts, both between and within skeletal material. Most peloids range from 10 to 60 µm in diameter, have cloudy centers, and clear exterior rims. In modern peloids, the rim consists of dentate High-Mg calcite crystals. The cores are interpreted as bacterially induced precipitates. The sample represents small cavities within boundstones forming a sponge/algal/carbonate cement reef. Middle Permian: Straza near Bled, northwestern Slovenia. 6 Bacterial/algal contribution indicated by thrombolitic fabric: ‘Peloidal micrite’, exhibiting a distinct clotted fabric, is a common constituent of ancient deep-water mud mounds. Formation and accumulation of the sediment by binding effects is explained as a result of the binding effects of bacteria and/or algae. Similar textures originate from the bacterial decomposition of sponges (see Fig. 4.2). Note the fenestral clotted packstone fabric (F: fenestrae). A clotted microstructure is taken as having been influenced by organically mediated calcification, either forming clots of bacteria or cementing peloids together. Central part of an algal mound. Late Carboniferous (Westfalian): Cantabrian Mountains, Northern Spain. 7 Cryptic microbialite containing small encrusting foraminifera (arrow). The composition of micro-encruster associations allows paleoecological conditions and sub-environments to be distinguished (Schmid 1996). Sessile foraminifera are often attached on macerated skeletons of sponges. Note faint relicts of a siliceous sponge (S) at the bottom. The decomposition of the tissue of siliceous sponges and subsequent microbial carbonate precipitation can result in the formation of peloidal and minipeloidal micritic fabrics (Warnke 1996). The association of sponges and microbes is a common feature of Phanerozoic mud mounds and reef mounds. Upper Jurassic reefs and massive buildups characterized by the abundance of siliceous sponges (Pl. 78) and calcified microbes are common on the European northern Tethyan shelf (Portugal; Spain; Southern Germany; Poland; Romania). Most mounds were formed within peri-/epicontinental seas on homoclinal mid and outer ramp settings. The mounds were composed of early lithified automicrites and of allomicrites. Late Jurassic (Oxfordian) sponge reef: Dobrogea, Romania. –> 1, 2, 7: Keupp et al. 1993; 3: Leinfelder 1985; 4: Böhm and Brachert 1993; 5: Flügel et al. 1984; 6: Hensen et al. 1995
Plate 8: Microbial Carbonates
87
88
and Von der Haar 1975; Shark Bay, Australia: Monty 1976; Gulf of Aqaba: Friedman et al. 1973, Gerdes and Krumbein 1987). Interpretation of limestones: The disintegration of externally calcified cyanobacteria is regarded as an important source of micritic reef limestones, especially of Paleozoic age (e.g. Pratt 2001). Allomicrites Allomicrites (allochthonous micrites) originate from the destruction of calcareous algae and invertebrate skeletons, bioerosion, accumulation of calcareous plankton as well as from the erosion and abrasion of lithified carbonates (Fig. 4.1). In shallow waters, tropical storms and ebb-tidal currents causing a mixing of bank and ocean waters can lead to a redistribution of genetically different carbonate muds. Carbonate muds produced on shallow-marine platforms are also a major source for periplatform sediments deposited on slopes or in basins (periplatform carbonates, see Sect. 2.4.5.7) (5) Disintegration of benthic calcareous algae: The Halimeda model. The disintegration of the skeletons of modern calcareous green algae causes the formation of sand- to clay-sized particles. Processes involved in the disintegration, deposition and redeposition of the algal particles are exemplified by the udoteacean alga Halimeda (Fig. 4.3; Pl. 7/1-2, Pl. 57/1-2). The ultrastructural elements of the calcareous skeleton of this and other green algae are aragonite needles. These needles are set free after decomposition of the organic matter. Morphologically similar needles are abundant in carbonate muds of the Great Bahama Bank and Florida Bay as well as in Pacific lagoons (Lowenstam and Epstein 1957; Stockman et al. 1967; Neumann and Land 1975; Nelsen and Ginsburg 1986). The algal origin of the needles was supported by (a) comparable stable isotope data of muds and algal needles, (b) SEM photomicrographs (Pl. 7/2) indicating morphological similarities between the aragonite produced by udoteacean algae and the aragonite needles occurring within the aragonite muds, and (c) excess productivity of udoteacean algae in the Bight of Abaco (Little Bahama Bank) where the algae produce more than twice the amount of fine-grained aragonite that has accumulated in the Bight (Neumann and Land 1975). Aragonite muds occurring in the deeper Northwest Providence Channel are, therefore, explained as the result of the extended offbank export of algal muds. These points, however, have to be discussed in the light of new data: (a) Isotope data could also indicate chemi-
Allomicrites
cal precipitation of the aragonitic muds, (b) SEM studies show that the aragonite needles in the Great Bahama Bank muds are bladed and have poorly developed pointed ends, in contrast to algal needles with blunt-ended prisms (Loreau 1982; Macintyre and Reid 1992). Nonskeletal aragonite needles are apparently precipitated ubiquitously over the Bahama Bank; they accumulate within the low-energy bank leeward of Andros Island and are exported to the deeper leeward slope (down to -90 m), and even to deep settings, as indicated by the position of the aragonite lysocline in a depth of 1500 m (Droxler et al. 1988), (c) Present-day standing stocks of green algae and carbonate production appear too low to account for the volume of Holocene aragonitic mud on top of the Bahama Bank. Algal mud production is estimated to be greater than an order of magnitude less than the precipitation rate calculated for the mud on the bank. Of course, a significant algal contribution to the submicroscopic needles of the muds can not be ruled out and has to be re-evaluated. Measurements of growth rates of Halimeda and other calcified udoteaceans suggest that these algae renew their standing stocks within a few weeks (Wefer 1980; Multer 1988; Hillis 1991, 1999) and could, therefore, produce tremendous amounts of fine-grained carbonate sediment in shallow lagoonal and reef environments. Halimeda mounds and banks, formed in deeper and shallow slope settings in eastern Australia, Indonesia and the Caribbean, document the enormous sediment-producing potential of these algae (see publications in Coral Reefs, vol. 6, 1988). The disintegration of algae, not only udoteaceans, is an additional important source of fine-grained sediments in modern reefs. Breakage and subsequent diminution of exposed calcified filaments of endolithic (boring) green algae (e.g. Ostreobium) and of filamentous cyanobacteria result in the production of silt- and micrite-size material (Kobluk and Risk 1977). Interpretation of limestones: Green algae, closely comparable with modern Halimeda, have been known since the Late Permian (Poncet 1989), were more common in the Late Triassic (Flügel 1988; Pl. 57/3), and became rock-building constituents of warm-water platform carbonates during the Cretaceous and in the Tertiary (Hillis 2001). Halimediform algae exhibiting a highly identical internal structure, however, occur as early as in the Ordovician and are more common in the Devonian (Pl. 57/4, 5). (6) Disintegration of epibionts living on seagrass and macro-algae: Seagrass (Thalassia, Posidonia, Zostera and others) is abundant in shallow subtidal and inter-
Halimeda Model
89
Fig. 4.3. The Halimeda model. Halimeda, a common calcareous green alga in lagoonal and reef environments, is one of the most prominent producer of modern carbonate muds and sands. The erect plant is several centimeters high and attached with a holdfast of root-like filaments to the sediment (B). The alga is segmented and branching. The segments possess an internal structure composed of interwoven tubular filaments which are differently arranged in the central and peripheral parts of the segments (C and D; Pl. 57/1, 2). Calcification starts at the surface between the filaments (D to F)) and progresses inwards. The skeleton consists of minute needles of aragonite (5 µm long and 0.3 µm wide; G; Pl. 7/2) which, on the death of the plant and disintegration of the segments, disperse in the sediment and provide a source for lime mud. The mud is often redeposited within the shallow-water area or in deeper environments on the slope. Isolated algal segments (Pl. 3/5) are a major component of calcareous back-reef sands and occur in the talus of many Cenozoic reefs. A: The first species of Halimeda was documented from the Mediterranean Sea in 1599 by the Italian naturalist Imperato. His figure illustrates superbly the characteristic segmentation, the branching pattern and the inconspicuous holdfast of Halimeda tuna. The holdfast of the pantropical species forms a mat that fixes the plant onto the rocks or corals. The width of the thallus is about 10 cm. B: The basic holdfasts of the sand-growing species act as anchors in times of storms. C: Internal structure of a segment. Note the different branching pattern and orientation of the filaments in the central part (medulla) and the peripheral part (cortex). D: Differentiation of cortical filaments. Calcification (gray pattern) takes place around and between the filaments. E: Calcification on and between medullar (left) and cortical filaments. F: Close-up of the cortex exhibiting remains of organic membranes; see also Fig. E. G: Aragonite needles. Sources: A, B and C - modified from Hillis-Colinvaux (1980); D - modified from Hillis (1959).
tidal settings. Seagrass can tolerate a wide range of salinities, it occurs most prolific in water depths between 2 and 15 m. The mud banks of Southern Florida, sometimes regarded as an analogue of ancient mud mounds (Bosence 1985), document four important controls of seagrass on carbonate mud production and energy baffling: (1) The long and dense root systems (rhizomes), extending a few tens of centimeters below the sea bed, stabilize and bind the sediment. (2) The long blades have a baffling effect, can trap passing suspension-trans-
ported particles and reduce current speeds (Ginsburg 1957; Scoffin 1970; Almasi et al. 1987). Depending on the density of the meadows, large amounts of muddy sediment (packstones and wackestones) will accumulate. (3) Seagrass stabilizes the substrate permitting colonization by stable-bottom sediment-producing and sediment-reworking communities, and (4) very importantly, sea-grass blades provide a constantly renewable substrate onto which benthic organisms can attach and grow (Ginsburg and Lowenstam 1958).
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Epibionts are integral members of seagrass communities. The disintegration of the epibionts is a major source of lime mud in Florida Bay (Turmel and Swanson 1976; Nelsen and Ginsburg 1986) and other areas (Land 1970; Patriquin 1972). Epibionts comprise a variety of encrusters such as coralline algae, foraminifera, bryozoans, bivalves, gastropods and serpulids (Brasier 1975; Ivany et al. 1990). Seagrass is not restricted to tropical environments but is also common in warm- and cool-temperate carbonate settings ( Gostin et al. 1984) where epibiontic bryozoans and foraminifera contribute significantly to the production of mud-, silt- and sand-sized particles (Collins 1988). Floating seaweeds (large brown macro-algae) provide another habitat for epibionts which contribute to the formation of fine-grained carbonates. The epibionts (mostly bryozoans, coralline algae) occurring on Sargassum and other kelps produce significant amounts of skeletal material, which add much to nearshore and offshore carbonate muds and sands (Pestana 1985; Wehrmann 1998). Interpretation of limestones: Seagrass is known as far back as the Late Cretaceous and floating macroalgae at least since the Jurassic, but perhaps from the Devonian. Fossil seagrasses are extremely rare in the rock record, but can be inferred by examining characteristic lithologies (carbonate lenses enclosed within shales: Petta and Gerhard 1977; fining upwards sequences: Wanless 1981), invertebrate assemblages and associated foraminifera (Eva 1980). A contribution of epibionts to the formation of fine-grained limestones of inner shelf settings, therefore, should not be ruled out. (7) Disintegration of invertebrate skeletons (bioclastic micrites): Sorby recognized as early as 1879 that invertebrate skeletons disintegrate after death by ‘maceration’ processes into microscopic elements forming calcareous muds and silt- to sand-sized particles. Destructive processes include: (a) Destructive diagenesis (physicochemical dissolution due to carbonate undersaturation), common in colder sea water at higher latitudes and characterized by dissolution and breakdown of solid carbonates on the sea floor and in contact with sea water. (b) Microbial degradation of skeletal-binding organic matter under aerobic conditions at the sediment/water interface or just below it (Henrich and Wefer 1986; Freiwald 1995). The susceptibility to dissolution of carbonate grains depends on the amount of skeletal binding organic matter, as well as on the size, shape and roughness of the crystallites (Walter 1985). The impact of degrading bacteria on biogenic carbonates is controlled by the distribution and supply of organic
Bioerosion
compounds on and within the skeletons, and the microstructure of the infested particle. (c) Degradation by microborers (see point 8). (d) Mechanical erosion and abrasion in high-energy environments (Folk and Robles 1964; Swinchatt 1965; Matthews 1966; Stieglitz 1972). Tidal currents and wave action in lagoons may be effective in grinding grains. Experiments (Chave 1964; Folk and Robles 1964) as well as studies of the mud and silt fraction of modern carbonates (Mathews 1966; Stieglitz 1972; Futterer 1979) prove that mechanical abrasion and biological processes work together in the production of muds, whereby bioerosion appears to be of greater importance in low-energy environments. Nearly all marine invertebrates are affected by this destruction process. Impressive examples were reported from benthic foraminifera, corals, bivalves and gastropods, echinoderms, barnacles as well as coralline red algae (Alexandersson 1979; Fitzgerald et al. 1979; Canfield and Raiswell 1991). Shells of modern benthic reefal foraminifera break down into micron-sized needles and rods through combined effects of corrosion, dissolution, abrasion and erosion (Kotler et al. 1992; Debenay et al. 1999). Interpretation of limestones: The formation of allomicrites by the disintegration of invertebrate skeletons appears to be a common mode both in shallow-marine and deep-marine settings. Fine-grained micritic sediments may still show the disintegrated basic skeletal elements as demonstrated by fine-grained Cretaceous limestones consisting of calcite prisms of shells of the bivalve Inoceramus (Pl. 87/8). Bioclastic micrites originating from the disintegration of corals, stromatoporoids, tabulates and crinoids were described from Silurian reefs (Watts 1981). (8) Bioerosion and microborings: Bioerosion by marine and non-marine boring and grazing organisms is a major process in the production of clay- and silt-sized carbonate particles contributing greatly to the formation of micrite and calcisiltite. Biological erosion destroys organic skeletons, grains and carbonate substrates, and creates cavities. Processes involved in bioerosion are corrosion by endolithic organisms and loosening of hard substrates by dissolution. Organisms which are able to disintegrate hard carbonate substrates and skeletons include microbes, algae, foraminifera, sponges, near-shore bryozoans, worms (sipunculids, polychaetes), mollusks, arthropods (cirripedians) and echinoderms (regular echinoids). Chemical-mechanical borings by clionid siliceous sponges in skeletons or rocks, produce large quantities of fine-grained carbonate sediment (up to 30% of the total sediment) in deeper
Accumulation of Plankton
intertidal and subtidal zones; most boring chips are siltsized (Futterer 1974, Moore and Shedd 1977). Clionids are common in tropical regions, but also occur at high latitudes. Boring bivalves (e.g. Lithophaga, Pl. 52/ 8) and snails occur preferentially in shallow seas. Of particular importance are heterotroph bacteria, cyanobacteria, green and red algae as well as fungi (Pl. 52/1-4) producing microborings and causing micritization (Pl. 12/2, Pl. 13/6). Bacterial attacks of organic matter in shells lead to the disintegration of skeletal grains into ultrastructural units. Borings of bacteria and fungi have diameters <4 µm, often 1-2 µm. Microborers produce species-specific boring patterns are used as paleobathymetrical markers (see Sect. 9.3.4). Bioerosion is responsible for significant carbonate production in reefs (Peyrot-Clausade et al. 1995; Van Treek et al. 1996). Abrasion and bioerosion of totally micritized skeletal grains provide the dominant source of Mg-calcite lime mud in some modern lagoonal environments (Reid et al. 1992). Carbonate mud production triggered by bioerosion is equally effective in tropical and non-tropical environments (Farrow and Fyfe 1988; Gaillard et al. 1994). Grazing and rasping organisms (e.g. mollusks, echinoderms, fishes) produce large amounts of fine-grained carbonate sediment in intertidal and subtidal environments. A combination of bioerosion caused by microborers, predominantly cyanobacteria, and subsequent grazing by gastropods and sea-urchins on endolithic microflora within the inter- and supratidal zone is one of the main causes of the erosion of limestone coasts (Torunski 1979). (9) Accumulation of calcareous plankton on the deep shelf and in basins: Modern pelagic carbonate muds deposited largely in water depths of several hundred to several thousand meters are dominated by (1) accumulations of calcitic planktonic foraminifera associated with tests of nannoplankton (foraminiferal or Globigerina mud), (2) accumulations of calcitic nannoplankton consisting predominantly of planktonic algae (coccolithophorids; see Winter and Siesser 1994 for an overview; coccolith mud), and (3) deposition of pelagic gastropods (aragonitic pteropods and heteropods; pteropod mud). Planktonic foraminifera evolved at the beginning of the Mesozoic and diversified through the Cretaceous and Cenozoic. The earliest quantitatively important calcareous nannofossils appeared in Late Triassic Tethyan marginal seas. Pteropods have been known since the Tertiary (Eocene). A pelagic sediment consisting of calcareous plankton is chalk (Pl. 7/3) deposited in Cretaceous deep shelf and basinal settings. The texture is similar to micrite,
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but chalk is white in color, less dense, and less compact than micrite. The sediment consists of diagenetically altered coccolith plates, smaller micrite grains and large euhedral calcite crystals resulting from pressure solution overgrowth. The low cementation potential of shelf chalks due to the primary Low-calcite mineralogy of most skeletal grains and the lack of meteoric pore fluids are responsible for the high reservoir qualities of the North Sea chalks (Scholle 1977; Zijlstra 1995). Planktonic foraminifera micrite (Pl. 6/6) Although only a small part of the original life planktonic assemblages become incorporated in the sediment, tests of planktonic foraminifera contribute significantly to the particle flux from shallow to deep waters and to the composition of pelagic carbonate muds (Berger 1971, 1975). The transition from death assemblage to sediment assemblage is strongly controlled by dissolution of calcite on the sea floor which in turn is a function of the fertility of the overlying waters and the aggressiveness of the bottom waters. Both factors depend on oceanic circulation patterns. Planktonic foraminifera are deposited with nanno-organisms and other calcareous plankton, e.g. calciodinoflagellate cysts (Pl. 73/2). Nannofossil micrite (nannomicrite, nannoagorite, ultramicrite; Pl. 7/3-5, 7) The contribution of calcitic nanno-organisms to the formation of very fine-grained micrites depends on primary productivity, depositional conditions and the course of diagenesis. Nannofossils include coccolithophorids as well as other nanno-organisms (cone-shaped nannoconids, schizospherids, spherical thoracosphaerids and others) belonging to systematically different planktonic algal groups. Modern coccolithophorids flourish in open-marine oceanic waters but also occur in lagoonal environments (Kling 1975) and reefs (Conley 1979). The cone-shaped nannoconids were abundant in Jurassic and Lower Cretaceous strata, but also occur in Late Cretaceous open-marine microcrystalline carbonates. They are commonly associated with radiolaria, calpionellids, and ammonites and may form almost pure nannoconid limestones (Pl. 77/23). Schizosphaerella (Pl. 7/7) is a common rock building fossil in Liassic pelagic carbonates (Kälin and Bernoulli 1984; Bombardiere 1993). Although there are a few reports of isolated coccolith-like fossils from Paleozoic sediments and the Middle Triassic as well (Vecsei and Duringer 1998), the earliest quantitatively important coccolithophorids appear in Late Triassic (Late Carnian) Tethyan marginal seas (Di Nocera and Scandone 1977; Bellanca et al. 1995).
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The appearance of calcareous plankton microorganisms in the Mesozoic has markedly moved the location of major carbonate deposition towards an oceanic setting. Nannofossils were abundant in Late Jurassic-Early Cretaceous fine-grained carbonates formed in shallow epicontinental and shelf settings (e.g. Solnhofen limestone, Flügel and Franz 1968; Keupp 1978; Noel et al. 1993), and also were deposited in deep-sea basins (Fig. 4.3). During the latest Jurassic the sink for carbonate deposition shifted from shallow seas to deep oceans. This is reflected by the common association of calpionellid limestones, ribbon cherts and pure nannofossil carbonates, first recognized by Steinmann (1925). Nannofossil productivity was sufficiently high in the pelagic regime to depress the carbonate compensation depth in the Tethyan Ocean. Owing to the Late Cretaceous transgressions, the pelagic carbonate sedimentation (Chalk facies) spread to epicontinental seas occurred. Many Cretaceous and Early Tertiary chalks consist almost exclusively of coccolithophorids. Post-Tertiary oceanic sediments have a smaller percentage of coccoliths and larger percentage of pelagic foraminiferal remains. The evidence of nannofossil contribution to micrite genesis requires detailed SEM studies of the ‘nannofacies’ (Noel 1969) as well as more sophisticated investigations of the orientation patterns of coccolith elements (Ratschbacher et al. 1994). Sedimentation and diagenesis of nannomicrites (Fig. 4.4) have been studied by many authors (e.g. Noel and Melguen 1978, Flügel and Keupp 1979). The good correlation of the coccolithophorid assemblages within the water masses and those found in surface sediments provide a powerful tool for the reconstruction of paleo-oceanographic patterns. Preservation patterns of coccolithophorids are important proxies for paleo-bottom currents and differential solution at the sea floor and within the sediment (McIntyre and McIntyre 1971; Schneidermann 1973). How do plankton skeletons become pelagic ooze? The sedimentation of calcareous plankton occurs by settling of individual tests, fecal pelletization, and inorganically formed flocs of biogenic material (‘marine snow’: Lampitt 1996). At the sea bottom plankton is affected by bottom currents and bioturbation. Although only a small part of the original life planktonic assemblages become incorporated in the sediment, tests of planktonic foraminifera contribute significantly to the particle flux from shallow to deep waters and to the composition of pelagic carbonate muds. The transition from death- to sediment-assemblage is strongly controlled by dissolution of calcite
Diagenetic Micrites
on the sea floor which in turn is a function of the fertility of the overlying waters and the aggressiveness of the bottom waters. Both factors depend on oceanic circulation patterns. The disintegration of plankton foraminiferal skeletons is poorly understood (Palmer-Julson and Rack 1992). Some planktonic foraminifera disarticulate into individual chambers and the chamber walls may dissociate into crystallites (Lowenstam and Weiner 1989). But most constituents of foraminiferal carbonate muds range from silt- to sand-size. The microcrystalline component of the foraminiferal mud consists of accumulations of coccoliths. In limestones and modern coccolith ooze, coccoliths are commonly observed as disarticulated plates and are rarely intact. The plates themselves may disassociate into crystallites. (10) Mechanical erosion of limestones: In high-energy environments limestone coasts are prone to mechanical erosion causing abrasion of clay- and silt-sized carbonate particles. Examples have been described from the Adriatic coast (Schneider 1977) and from the coasts of the Persian Gulf and Puerto Rico (Pilkey and Noble 1966; Kukal and Saadallah 1973). Abrasion is facilitated by coeval bioerosional loosening of rock surfaces. Rivers as well as wind transport the fines. Interpretation of limestones: Indications for micrites formed by abrasion are poor sorting and the association of fine-grained carbonate and non-carbonate particles (Lindholm 1969), a very small crystal size (<2.4 µm), as well as rounding and good sorting (Peszat 1991). Because of the lack of calcareous plankton in the Paleozoic, some authors have argued that Paleozoic micrites formed in deeper ramp, slope and basinal positions, may represent fine-grained carbonate eroded and winnowed from shallower platforms or carbonate coasts. The ten modes of possible origins of micrite described above correspond more or less to the original definition given by Folk (1959, 1962) which focuses on crystal size and inferred sedimentary origin. Both crystal size and the different clues to the origin of micrites are more successfully studied by means of scanning electron microscopy and geochemical data than solely by investigating thin sections alone. Diagenetic micrites (pseudomicrites) The recognition of the diagenetic character of microcrystalline carbonate rocks is of crucial importance, particularly for interpreting ancient reefs and speculating on the energy level during the deposition of finegrained sediments on platforms and ramps.
Solnhofen-Oberalm Model
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Fig. 4.4. The Solnhofen-Oberalm model. Biotopes, depositional settings and diagenetic environments of coccoliths and other calcareous nanno-organisms. A: Most coccolithophorids flourish in water depths down to about 100 m. Concentration maxima occur in the tropical zone at depths of about 50 m, in temperate zones at 10-20 m. Major controls on the distribution of coccolithophorids are light, water temperature and salinity. Different assemblages occur in tropical, temperate and subarctic zones. B: Most coccoliths are transported to the sea bottom enclosed in fecal pellets produced by zooplankton and fish. Free and isolated coccoliths are dissolved at the hydrographic lysocline. C: Increasing dissolution at the water/sea bottom interface occurs at depths below 3000 m. The diversity decreases strongly at the ‘coccolith lysocline’ (LY) near 4000 m. Selective dissolution patterns of coccolithophorids characterized by increasing corrosion and fragmentation are used as paleobathymetrical indicators. The zone of complete coccolith dissolution (coccolith calcite compensation depth) lies at depths of about 5300 m. Diagenetic processes involving dissolution of parts of the coccoliths and syntaxial overgrowth start at the sea bottom and are continuous within the sediment. D: Grain-to-grain pressure solution related to increasing burial causes further dissolution at burial depths > 300 m, and provides a source for the carbonate needed for the growth of porefilling calcite. Another source is the dissolution of aragonitic shells. The figure addresses the importance of nannofossils for the formation of allomicrites both in Late Jurassic deep-water and shallow-water settings. Coccoliths occur in the lagoonal Solnhofen Limestone of Bavaria with an abundance of up to 80% on bedding planes, and in the Oberalm Limestone in the Austrian Alps (abundance > 50% within the beds). The carbonate deposition in the Solnhofen lagoon was controlled by an alternation of hypersaline conditions favoring the growth of cyanobacteria mats and normal marine conditions with the sedimentation of ultraplankton and the input of allochthonous carbonates. The light-colored bedded Oberalm Limestone (Pl. 76/1) represents a common Late Jurassic to Early Cretaceous bathyal facies that is widely distributed in the Alpine-Mediterranean region. The limestone is also known as Aptychus limestone, Majolica, or Biancone. The carbonate production was controlled by the deposition of coccoliths and nannoconids as well as by the input of fine-grained carbonate derived from nearby platforms into open-marine basins. Adapted from Flügel and Keupp (1979).
(11) Micro- and cryptocrystalline calcite cements: Detailed studies reveal that many ‘micrites’ are in reality crypto- and microcrystalline calcite or aragonite cements precipitated in very different settings including reef cavities (Friedman 1985; Reid et al. 1990; Scoffin 1993), slope environments (Wilber and Neumann 1993), ocean floors, beachrock (Pl. 31/3) or caliche crusts (Sect. 2.4.1.1). A growing number of authors calls these micro- and cryptocrystalline cements ‘micrite’ or ‘micritic cements’, not taking into consideration the continuing struggle for clearness in sedimentological terminology (Friedman 1985; Milliman et al. 1985).
Micro- and cryptocrystalline cements are formed in place and, therefore, correspond to ‘automicrites’ as described above; see (1) to (4). Differentiating these micritic autochthonous microfabrics from allochthonous micrites is essential for accurately interpreting of depositional environments. A ‘micritic matrix’ occurring in intergranular pores of carbonate slope packstones or in voids of reef carbonates should not be considered a priori as an indication of low energy conditions (Reid et al. 1990; Wilber and Neumann 1993). Bahamian slope carbonates demonstrate the formation of autochthonous ‘muddy’ carbonates by interstitial precipitation of fine-grained Mg-calcite at the
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Microspar
thermocline-layer depths (200 to 800 m) and in areas of strong currents. This process can transform sedimentary grain-support fabrics into diagenetic mudsupport fabrics, and thus give the paleoenvironmental interpretation a wrong direction. A cyanobacterial origin of micrite cements has been described from Pleistocene carbonates (Kazmierczak and Iryu 1999). (12) Recrystallization and grain diminution: Recrystallization of grains (e.g. peloids, see Sect. 4.2.2) and grain-diminution (Orme and Brown 1963) of skeletal grains are further possibilities for the formation of a ‘micrite’ fabric. Grain diminution (degrading neomorphism) takes place both during early diagenetic and burial diagenetic phases (Dixon and Wright 1983). It results in a replacement of former larger crystals by tiny crystals due to changes in crystal texture or changes in mineralogy. Examples are known from modern and ancient algal carbonates, where early diagenetic microcrystalline cements are formed within the algal framework (Alexandersson 1977). The dissolution of sparry cements in grain-supported limestones and the precipitation of microcrystalline cement (Wardlaw 1962; sparmicritization, Kahle 1977) can also cause the formation of diagenetic micrite. Alteration of skeletal grains due to a combination of micritization and recrystallization lead to the formation of modern lagoonal muds composed of micrite-sized clumps of equant anhedral Mgcalcite crystals about 0.5 µm in size (Reid et al. 1992).
4.1.3 Microspar Microcrystalline carbonate rocks display a wide range of mean crystal sizes. For examples see Pl. 6/2; Pl. 47/2, Pl. 91/7). Folk (1959, 1965) distinguished between ‘micrites’ (<4 µm) and ‘microspar’ (4–30 µm). Definition: The term refers to a fine-grained calcite matrix, characterized by rather uniformly sized and generally loaf-shaped, subhedral and euhedral calcite crystals ranging from 5 to more than 20 µm in diameter (Folk 1959). Other authors use different crystal size limits: 4–10 µm (Tucker 1981) or 5–50 µm (Bathurst 1975). Further criteria are a mosaic-like microtexture, equant grain shapes and boundaries, pits within the crystals (interpreted as a result of the dissolution of former aragonite needles enclosed in poikilotopic calcite crystals), impurities of clay or organic matter between the crystals, sometimes a patchy distribution within the micrite, and the association of microsparitic limestones with shaley interbeds. Microsparite is a limestone
Fig. 4.5. Two models explaining the formation of microspar textures in fine-grained limestones. A: Aggrading neomorphism (slightly modified from Folk 1959, 1965). The microspar texture is formed by the conversion of smaller High-Mg calcite crystals (top) to larger calcite crystals (bottom) by gradual enlargement maintaining a uniform crystal size at all given times. B: Transformation of aragonite-dominated lime mud to microspar without a micrite stage (modified after Lasemi and Sandberg 1984). The sketch, based on the investigation of microcrystalline Silurian limestones, illustrates the processes leading to the formation of microspar (after Munnecke 1997). The precursor carbonate mud (top, A) consists of aragonite needles, Low-Mg calcitic and High-Mg calcitic and aragonitic bioclasts, and clay minerals. Cementation (B) in intergranular pores starts with the precipitation of microspar crystals that enclose smaller aragonite needles and bioclasts in the sediment close to the sea bottom. The boundaries of the microspar crystals are accentuated by clay seams. The dissolution of aragonite needles causes pit-like structures (C). Aggrading neomorphism (D and E) destroys larger bioclastic grains and produces the final microspar fabric. Note that the size and the shape of microspar crystals are determined in a one-step process by cementation (B), in contrast to the Folk model according to which the microspar fabric originates at the end of a long-lasting process.
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Calcisiltite
whose fine-grained matrix is developed as microspar. Sometimes the term is used simultaneously with microspar. Folk (1959) described a gap in crystal sizes somewhere between 3 and 4 µm between micrite and microspar. The data of Lasemi and Sandberg (1993) suggest that the ‘micrite curtain’ (Folk 1965) is an artifact of lumping two mineralogically different groups of micrites. Munnecke (1997), studying Silurian fine-grained carbonates, also found no gap. Origin: Several explanations for the formation of microsparitic limestones have been suggested (Fig. 4.5): (1) Recrystallization (aggrading neomorphism: Folk 1965) of micrite calcite subsequent to the removal of Mg2+ ions which form a ‘micrite curtain’ and are responsible for the gap in crystal size somewhere between 3 and 4 µm between micrite and microsparite (Folk 1974). Breaking through the ‘Mg-cage’, which in turn initiates recrystallization, is triggered by freshwater input related to freshwater aquifers, the influence of rain water, and weathering near the surface (Longman 1977). Interbedded clay could attract magnesium ions from adjacent micrites to facilitate microspar formation through aggrading neomorphism. (2) One-step neomorphic process of cementation and replacement (calcitization) of aragonite-dominated precursors to microspar without a micrite stage, also connected with the infiltration of meteoric waters. Porosity reduction occurs without compaction. Examples: Holocene and Pleistocene (Steinen 1978, 1982; Lasemi and Sandberg 1984; 1993), Pennsylvanian (Wiggins 1989) and Silurian (Munnecke and Samtleben 1995; Munnecke 1997). (3) Neomorphic growth from deep surface fluids (Brand and Veizer 1981; Suchecki and Hubert 1984). Wiggins (1989) demonstrated that saturation with respect to calcite appears to be the only requirement for microspar formation. This situation can also be met independently of freshwater if organic acids react with aragonite in marine environments (Baker et al. 1982). (4) Recrystallization of silt-sized carbonate grains (Duringer and Vecsei 1998). Significance: Depending on the model used, the occurrence of microsparitic limestones has been regarded as evidence for meteoric diagenesis, for the influence of low-salinity fluids or for subaerial exposure. Honjo (1969) pointed to the frequency of microsparitic limestones in the Paleozoic and the abundance of micritic limestones in the Mesozoic. The author discusses the possibility of controls by changes in global sea-water chemistry. A more likely explanation is that the longer time involved in the diagenesis of Paleozoic
Fig. 4.6. Polygenetic carbonate mud production in a shallowmarine platform environment. Lime mud budget for the Bight of Abaco, Little Bahama Bank. Numbers indicate possible modes of origin (see Fig. 4.1). The Bight of Abaco is a semi-closed, 7 m deep shallow carbonate lagoon. Temperature and salinity vary seasonally from 20 °C to 30 °C, and from 34 to 39‰, respectively. Skeletal aragonite and Mg-calcite muds and sands are the predominant sediments. The main source of fine-grained carbonates are sub-microscopical elements of calcareous green algae. Additional sources include fine-grained material produced by the bioerosion, abrasion and breakdown of carbonate grains, and by bacterially-induced and possible chemical precipitation of aragonite and Mg-calcite. An important source is the import of inter- and supratidal muds to the lagoon. Lagoonal pellets disaggregate and return to the fined fraction; the disintegration of cemented pellets formed near the bank edge provides a source for lagoonal muds. The export of lagoonal muds to tidal flats and onshore mud transport are triggered by storms. Some lagoonal carbonate is exported in solution. The Abaco example demonstrates that more carbonate mud is produced by algal disintegration and other sources than can be accommodated on the platform top, and that mud transport moves on to adjacent tidal flats and into deeper waters around the platform. Based on Neumann and Land (1975).
carbonate rocks has favored the formation of microspar fabrics. Global chemical abd biochemical controls have been suggested for the abundance of microspar in the enigmatic Proterozoic ‘molar tooth structures’ (Long 2003) 4.1.4 Calcisiltite Definition: The name was proposed by Kay (1951) for limestones consisting predominantly of detrital calcite particles of silt size (2–63 µm). Calcisiltite embraces the coarser part of ‘calcilutite’ (Grabau 1904) which designates a fine-grained rock composed of clayand silt-sized carbonate particles. Later Lindholm
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Check List for Micrites
Box 4.3. Check your sample! A) Sea-floor micrite or internal micrite? Check your sample to discover whether you are looking at a sea-floor micrite formed at the bottom/water interface (the more common case), or at an internal micrite formed within the cavities and intergranular voids of the sediment. This distinction is of major importance in the paleoenvironmental evaluation of micritic rocks, because internally precipitated micrite alters the mud/grain ratio and may invalidate interpretations that relate the amount of micrite matrix to water energy levels (Reid et al. 1990). Internal micrites form (a) in the water column or on the sea floor, the skeletal or non-skeletal grains are transported into the cavities. Or, (b) the internal micrite forms within the cavities. The sediment consists of skeletal debris or unattached precipitates, or of attached non-skeletal precipitates. The last two cases correspond to the autochthonous ‘container organomicrites’ (Fig. 4.2) and are characterized by common peloidal fabrics (Pl. 6/4). The micrite within cavities differs in color, microtexture, crystal size and included grains from the micrite of the host rock. Infilling of micrite may be totally or restricted to the bottom part of the cavity, resulting in a geopetal structure (e.g. in ‘stromatactis’, Pl. 20/1, 8). Alternation of micrite and grains may result in normal and reverse gradation structures. Internal micrite should not be confused with the diagenetically formed internal vadose silt (see Sect. 7.4.2.1). B) Describe the overall microfabric! Microfabrics give indications of the possible origin of the fine-grained matrix and assist in attributing your sample to one (or several) of the groups summarized in Fig. 4.1. Homogeneous/massive matrix (Pl. 6/1; Pl. 90/1): Appears dense in thin sections, dark in transmitted light, commonly without any or with only a few grains, often without fossils, and with no or only vague lamination. May point to group 1 of Fig. 4.1, but also to a diagenetic pseudomicrite matrix (modes 11 and 12). Inhomogeneous matrix (Pl. 19/1, 6; Pl. 25/4): Characterized by irregularly distributed micrite areas differing in color (dark and light gray), microfabric (with or without peloids and bioclasts) and the degree of reworking, which is often caused by burrowing and bioturbation. Peloidal matrix (Pl. 8/6): Fine-grained matrix with abundant very small peloids (no fecal peloids), commonly densely packed, sometimes exhibiting a clotted fabric, appearing in various shades of gray in transmitted light. Often associated with sponges. Common fabric of internal micrites and organomicrites. Points to modes 2 and 3. Fine-bioclastic matrix: (Pl. 6/1, Pl. 87/8-9): Micrite containing abundant very small bioclasts, which appear white in transmitted light. Bioclasts originate from the disintegration of various invertebrates, and often cannot be identified in thin sections. Points to the modes 6 to 8. Laminated matrix: Fine, curved or plane laminae, consisting of homogeneous and/or peloidal micrite, sometimes alternating with spar- and grain-bearing zones. Are laminae boundaries sharp or gradual? Lamination may reflect (a) differences in automicrite production (autochthonous growth of microbialites; modes 2, 3 and 4; Pl. 18/4; Pl. 50/2, 4, Pl. 123/ 3), (b) differences in sediment deposition (e.g. caused by tides or bottom currents; Pl. 18/2, Pl. 20/8), or (c) pressure solution seams (see Sect. 7.5.2; Pl. 37/6). Siltitic matrix: Is the bulk of the matrix within the silt-sized range, and do small angular lithoclasts occur? This might point to the abrasion of former carbonate rocks (mode 10). C) Are there conspicuous textural differences in your sample? Marked differences in texture, rock color and the type of common fossils (e.g. encrusting and boring biota) may indicate changes or breaks in sedimentation associated with hardgrounds (common in micrites of mode 9). D) Try to evaluate crystal size and crystal fabric! Petrographical thin sections allow the measurement of crystal sizes as well as the recognition of crystal shapes (an-, subor euhedral) and fabrics which may be equant (consisting of nearly equidimensional crystals) or non-equant (composed of different sized crystals). Grain size is the conventional measure in distinguishing micrite and microspar. Fabrics made of non-equant crystals may point to recrystallization (Pl. 38/3; mode 12). E) Take a look at the fossils and biogenic textures occurring within the micrite matrix! Fossils should be differentiated according to their life habits (planktonic, benthic; soft or hard bottom dwellers), frequency and preservation. Borings and fragmentation of calcareous green and red algae may indicate mud production related to the disintegration of benthic algae living on the sea bottom or on seagrass and soft macro-algae (modes 5 and 6). Strong fragmentation and intensive borings of invertebrate skeletons points to the modes 7 and 8. An abundance of calcareous plankton suggests mud formation from the accumulation of pelagic organisms (mode 9). F) Non-carbonate material Many microcrystalline limestones contain varying amounts of clay, quartz and other non-carbonate minerals. Care should be taken in differentiating the mineralogical composition and distribution (random, in zones) of these constituents. The percentage of clay in micritic limestones is a major control on recrystallization. Interbedded clay could attract magnesium ions from adjacent micrites and facilitate microspar formation through aggrading neomorphism (Sect. 7.6).
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Genetic Interpretation of Micrites
Box 4.4. Microfacies criteria useful for the genetic interpretation of micritic limestones include texture, fabric, and the preservation and association of fossils. Differences in dominating rock colors are caused by variations in non-carbonate constituents (e.g. terrigenous input, organic matter, authigenic minerals), pH conditions, and the crystal size of the finegrained carbonate matrix. References to settings concern only the more common cases. The numbers refer to the modes of origin summarized in Fig. 4.1.1. 1 Lime mudstones. Homogeneous very fine-grained ...................................................................................... Pl. 125/1 2 Boundstones, mudstone, wackestones, packstones. Inhomogeneous matrix. Some dark, very fine-grained micrite. Peloidal, thrombolitic and stromatolitic fabrics. Close association of sponges and in situ microbial structures. Sponge structures replaced by micropeloidal fabrics. Differentiated microbial fabrics developed at the surface of the sediment (crusts and vertical pillars, encrusting organisms, bored hardgrounds) or within semi-closed cavities. In situ peloids and ooids. Gray and dark rock colors within laminae. Common in ramp, deep shelf, mud mound and reef settings. ........................................................................................................... Pl. 6/5; Pl. 10/1; Pl. 28/1; Pl. 50/1, 2, 4 3 Bindstones, mudstones. Inhomogeneous matrix. Indistinct wavy lamination. Very small peloids exhibiting white calcite cement rims. Inferred bacterial structures within cavities and voids. Common in ramp, deep shelf and mud mound settings. Also in non-marine environments. ........................................................ Pl. 8/4-5; Pl. 18/1; Pl. 121/3 4 Bindstones. Very fine-grained matrix (minimicrite). Relicts of cyanobacteria, commonly filaments (Porostromata). Tidal and shallow subtidal settings. Also in non-marine environments. ........................................... Pl. 7/6; Pl. 124/1 5 Wackestones. Inhomogenous matrix. Abundant fragmented udoteacean calcareous green algae or gymnocodiacean algae. Shallow-marine subtidal settings, reefs, upper slope. ....................................................................... Pl. 57/3, 6 6 Wackestones containing abundant transported epibionts (foraminifera, serpulids, small mollusks, red algae). Characteristic biotic associations. Spatially restricted lenses consisting of fine-grained carbonates and occurring within shaly limestones. Tidal and shallow subtidal settings. .................................................................................... Pl. 67/2 7 Fine-bioclastic wackestones and mudstones. Often dark, very fine-grained micrite. Skeletal debris derived from fragmented and disintegrated shells of invertebrates. Association of debris and large bioclasts and fossils. Dark and variegated rock colors. Shelf and deeper ramp settings, base-of-slope and basinal settings. Also in non-marine environments. ..................................................................................................... Pl. 91/2; Pl. 118/1; Pl. 144/8-9; Pl. 147/4 8 Wackestones, packstones. Abundant micritized grains. Microborings in skeletal and non-skeletal grains. Gray and dark rock colors. Deep subtidal shelves and ramps. .................................................................................... Pl. 10/3, 5 9 Wackestones, packstones, mudstones. Abundant pelagic microfossils (planktonic foraminifera, calpionellids, ‘calcispheres’) and nannofossils. Light and dark rock-colors. Common in open-marine environments. Deep shelf, slope and basinal settings. ......................................................... Pl. 29/5; Pl. 73/2, 3, 10; Pl. 76/1, 7, 8; Pl. 77/23; Pl. 99/10 10 Wackestones, packstones. Inhomogeneous calcisiltite matrix. Mixing of silt-sized and larger angular and subrounded lithoclasts in association with poorly preserved bioclasts and non-carbonate clasts. Various rock colors. Shelf and ramps as well as slope settings. ..................................................................................................................... Pl. 141/2 11 Common in boundstones. Fine-grained inter- and intragranular calcite cement, dark in transmitted light. Homogeneous fabric. Microspar texture. Formed during shallow and deep burial diagenetic environments. Affects marine and non-marine carbonates (e.g. caliche). ....................................................................................................... Pl. 47/2 12 Patchy recrystallization of micrite to microspar and in places to pseudospar. The recrystallization affects both the fine-grained matrix and the grains. The crystals of bioclasts (coralline algae, mollusks) may exhibit diminution and homogenization of size. Formed during shallow and deep burial diagenetic environments. Affects marine and nonmarine carbonates. ....................................................................................................................................... Pl. 38/3, 6
(1969) described a Devonian limestone, using the term calcisiltite for fine-grained, poorly-sorted carbonates (grain sizes 5–15 µm) which probably resulted from the breakdown of skeletal material. This fabric originating from mechanical deposition should not be confused with microspar which describes a diagenetic fabric. Bioclastic calcisiltites, produced by abrasion and bioerosion processes, are common in modern shallow-ma-
rine lime mud environments (Folk and Robles 1964; Swinchatt 1965; Fütterer 1978). Microbioclastic calcisiltites, however, also occur in deeper-water ramp, slope and basinal settings. Abiogenic or biogenic breakdown of a calcite substrate following the initial diagenesis (microsparitization: Pomar et al. 1975; sparmicritization: Kahle 1977) can produce a ‘diagenetic sediment’ consisting of silt-sized crystals. This crystal silt or vadose silt
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(Dunham 1969; Pl. 34/7) should not be confused with mechanically deposited calcisiltite (see Sect. 7.4.2.1).
4.1.5 Practical Aids in Describing and Interpreting Fine-Grained Limestones How do we describe fine-grained carbonate rocks? Box 4.3 and Box 4.4 summarize the main questions which should be answered in a thorough description of micritic limestones. The answers are necessary for tackling the difficult question: ‘What do you think about the origin of these micrites?’ (see Fig. 4.6). How do we interpret fine-grained carbonates? The genetic interpretation of micrites is difficult but can however, be facilitated by a combined use of microscopic, submicroscopic and geochemical criteria. Fig. 4.6 summarizes some key criteria, but it must be emphasized that, particularly in shallow-marine and also in non-marine environments, several genetic modes work together, as exemplified by the calcareous mud budget of the Bight of Abaco, Bahamas. This example as well as many others demonstrate the central importance of mud import and mud export from shallowwater settings to deep-water settings. Export from shallow platform areas to deep shelf settings was a common process during the Paleozoic and Mesozoic (e.g. Pittet et al. 2000).
4.1.6 Significance of Fine-Grained Carbonates Micritic carbonates offer a wealth of information on the controls of the origin and physicochemical properties of fine-grained limestones. Box 4.5 lists some of these points. Further explanations are provided in the chapters indicated in the Box. Basics: Matrix and micrite Bernier, P. (1994): For a reinstatement of ‘lithographic’, a precise word to define a precise limestone. – Geobios, 16, 307-311 Borza, K., Harman, M. (1970): Lithologisches and elektronenmikroskopisches Studium von Kalksteinen des Oberjuras und der Unterkreide. – Geologicky Zbornik, Geologica Carpathica, 21, 139-151 Castanier, S., Le Métayer-Levrel, G., Perthuisot, J.P. (2000): Bacterial role in the precipitation of carbonate minerals. – In: Riding, R.E., Awramik, S.M. (eds.): Microbial sediments. – 32-39, Berlin (Springer) Charaklis, W.J., Marshall, K.C. (1990): Biofilms. – 316 pp., New York (Wiley) Fischer, A.G., Honjo, S., Garrison, R.E. (1967): Electron mi-
Significance of Fine-Grained Carbonates
crographs of limestones and their nannofossils. – Monographs in Geology and Paleontology, 1, 1-141 Flügel, E. (1967): Elektronenmikroskopische Untersuchungen an mikritischen Kalken. – Geologische Rundschau, 56, 341-358 Flügel, E., Franz, H.E., Ott, W.F. (1968): Review on electron microscope studies of limestones. – In: Müller, G., Friedman, G.M. (eds.): Recent developments in carbonate sedimentology in central Europe. – 85-97, Berlin (Springer) Flügel, E., Keupp, H. (1979): Coccolithen-Diagenese in Malm-Kalken (Solnhofen/Frankenalb, Oberalm/ Salzburg). – Geologische Rundschau, 68, 876-893 Folk, R.L. (1959): Practical petrographical classification of limestones. – Amer. Ass. Petrol. Geol. Bull., 43, 1-38 Folk, R.L. (1962): Spectral subdivision of limestone types. – In: Ham, W.E. (ed.): Classification of carbonate rocks. – Amer. Ass. Petrol. Geol. Mem., 1, 62-84 Folk, R.L. (1965): Some aspects of recrystallization in ancient limestones. – In: Pray, L.C., Murray, R.C. (eds.): Dolomitization and limestone diagenesis. A symposium. – Soc. Econ. Paleont. Min. Spec. Publ., 13, 14-48 Honjo, S. (1969): Study of fine grained carbonate matrix: sedimentation and diagenesis of ‘micrite’. – Paleontological Society of Japan Special Paper, 14, 67-82 Kazmierczak, J., Coleman, M.L., Gruszczynski, M., Kempe, S. (1996): Cyanobacterial key to the genesis of micritic and peloidal limestones in ancient seas. – Acta Palaeontologica Polonica, 41, 319-338 Kazmierczak, J., Iryu, Y. (1999): Cyanobacterial origin of microcrystalline cements from Pleistocene rhodoliths and coralline algal crusts of Okierabu-jima, Japan. – Acta Palaeontologica Polonica, 44, 117-130 Keim, L., Schlager, W. (1999): Automicrite facies on steep slopes (Triassic, Dolomites, Italy). – Facies, 41, 15-26 Lampitt, R.S. (1996): Snow fall in the open ocean. – In: Summerhayes, C.P., Thorpe, S.A. (eds.): Oceanography. An illustrated guide. – 96-112, New York (Wiley) Lasemi, Z., Sandberg, P. (1993): Microfabric and compositional clues to dominant mud mineralogy of micrite precursors. – In: Rezak, R., Lavoie, D.L. (eds.): Carbonate microfabrics. – 173-185, New York (Springer) Loreau, J.-P. (1972): Pétrographie de calcaires fins au microscope électronique à balayage: introduction à une classification des ‘micrites’. – Comptes Rendúes, Acad. Sci. Paris, 274, 810-813 Munnecke, A. (1997): Bildung mikritischer Kalke im Silur von Gotland. – Courier Forschungsinstitut Senckenberg, 198, 1-131 Munnecke, A., Samtleben, C. (1996): The formation of micritic limestones and the development of limestone-marl alternations in the Silurian of Gotland, Sweden. – Facies, 34, 159-176 Neuweiler, F., Reitner, J., Monty, C. (eds., 1997): Biosedimentology of microbial buildups. IGCP Project No. 380, Proceedings of the 2nd Meeting Göttingen/Germany 1996. – Facies 36, 195-284 Pratt, B.R. (1995): The origin, biota and evolution of deepwater mud-mounds. – In: Monty, C.L., Bosence, D.W.J., Bridges, P.H., Pratt, B.R. (eds.): Carbonate mud-mounds. Their origin and evolution. – Int. Ass. Sedimentologists Spec. Publ., 23, 49-123 Pratt, B.R. (2001): Calcification of cyanobacterial filaments: Girvanella and the origin of lower Paleozoic lime mud. – Geology, 29, 763-766
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Significance of Micrites
Box 4.5. Significance of micritic limestones - a first glance. Paleoenvironmental proxies Water energy levels: Microcrystalline carbonates are tradionally thought to represent mud accumulated in lowenergy quiet waters on the sea floor. However, fine-grained calcite also forms as an internal precipitate or sediment in cavities below the sediment/water interface, especially in reef carbonates. This internal micrite can invalidate interpretations relating the amount of micrite matrix to the degree of water energy. Another pitfall is that finegrained carbonates can be deposited and preserved in high-energy environments, if organisms (e.g. microbes) bind the sediment (Sect. 12.1.1). Changes in sediment input: The formation of accretionary organomicrites requires low or moderate rates of sediment input. Bottom conditions: Differences in organisms living on and within the sediment, and favoring soft, solid and hard bottoms are indicators of substrate consistency Sect. 12.1.3). Depositional settings of fine-grained carbonates Lagoons: Limestones formed in lagoons can be differentiated from open-marine micritic carbonates by dominating microfacies types and specific associations of fossils (Sect. 10.3). Reefs: Fine-grained carbonates originate in reefs inside framework and solution cavities. Different parts and growth stages of framework reefs are characterized by the specific microfabrics of these internal sediments, which correspond partially to allomicrites, but also include biologically induced automicrites. Mud mounds exhibit diagnostic microfacies types (Sect. 16.2.5.1). Periplatforms: The transitional character of fine-grained periplatform carbonates accumulating on slopes or in basins is reflected by specific microfacies types. Diagnostic criteria are platform-derived grains associated with pelagic biota as well as diagenetic fingerprints (Sect. 15.7.5). Slopes: Steep ancient and modern slopes were stabilized by biologically induced automicrites that exhibit specific microfacies patterns (Sect. 15.7.4). Pelagic micrites: Microfacies as well as diagenetic criteria and geochemical signals characterize depositional settings and assist in the understanding of the ooze/chalk/limestone transition (Sect. 7.1.4 and 15.8). Global secular variations Temporal changes in the dominating mineralogy: The dominant mineralogy of the mud precursors of micrites has a strong influence on the resulting microfabric. A distinction between the aragonite-dominated-precursor (ADP) and calcite-dominated precursor (CDP) micrites can be made by recognizing aragonite relicts, differences in crystal surfaces and crystal size, differences in microfabrics, as well as on the basis of geochemical data. The study of the Phanerozoic record of micrites and other carbonate constituents points to the existence of secular variations in carbonate mineralogy and global changes in the chemistry of the oceans over time. Technological properties Reservoir rocks: The microporosity of micrites is one of the reasons that giant hydrocarbon fields are related to finegrained limestones, not only to genuine chalks in the North Sea but also to chalky micritic limestones, e.g. in the Middle East (Sect. 17.1.2). Limestone resources: Many micrite-dominated reefs, particularly mud and reef mounds, are characterized by the exceptional purity of the limestones, which makes the carbonates very suitable for industrial use. Micritic limestones with high intergranular porosity have a high aptitude for mechanical disintegration and grainy destruction, which are useful properties for the production of high-quality technical chalk. Many technological properties of finegrained carbonates are controlled by the sedimentary fabric and the course of diagenesis (see Chap. 18).
Reid, R.P., Macintyre, I.G., James, N.P. (1990): Internal precipitation of microcrystalline carbonate: a fundamental problem for sedimentologists. – Sed. Geol., 68, 163-170 Reitner, J. (1993): Modern cryptic microbialite/Metazoan facies from Lizard Island (Great Barrier Reef, Australia): formation and concepts. – Facies, 29, 3-39 Reitner, J., Gautret, P., Marin, F., Neuweiler, F. (1995): Automicrites in a modern marine microbialite. Formation model via organic matrices (Lizard Island, Great Barrier Reef, Australia). – Bulletin de l’Institut océanographique, Monaco, no. spécial, 14, 237-263 Reitner, J., Neuweiler, F. (1995): Mud mounds: a polygenetic spectrum of fine-grained carbonate buildups. – Facies, 32, 170 Scoffin, T.P. (1993): Microfabrics of carbonate muds in reefs.
– In: Rezak, R., Lavoie, D.L. (eds.): Carbonate microfabrics. – 65-74, New York (Springer) Riding, R.E., Awramik, S.M. (eds., 2000): Microbial sediments. – 331 pp., Berlin (Springer) Sorby, H.C. (1879): The structure and origin of limestones. – Proceedings of the Geological Society London, 35, 59-95 Winter, A., Siesser, W.G. (1994): Coccolithophorids. – 344 pp., Cambridge (Cambridge University Press) Wise, S.W., Stieglitz, R.D., Hay, W.W. (1970): Scanning electron microscopy study of fine grain size biogenic carbonate particles. – Transact. Gulf Coast Ass. Geol. Soc., 20, 287-302 Wolf, K.H. (1965): Gradational sedimentary products of calcareous algae. – Sedimentology, 5, 1-37 Further reading: K033, K078
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Carbonate Grains
4.2 Carbonate Grains Various names are in use for organic and non-organic particles of limestones that are larger than the groundmass: Grains, particles, constituents, and ‘allochems’. The last term is a collective term for mechanically deposited grains that have undergone transportation in most cases (Folk 1962). Grosso modo skeletal grains and non-skeletal grains can be separated. Skeletal grains comprise complete or fragmented fossils. Common ‘non-skeletal’ grains are peloids, various coated grains (Peryt 1983), grain aggregates and clasts. The strict differentiation of ‘skeletal’ and ‘non-skeletal’ grains, however, has been discarded, because the formation of most coated grains and grain aggregates as well as some ooids and peloids is substantially controlled by biotic processes.
COATED GRAINS
PELOIDS
Terminology of carbonate grains: Fig. 4.7 summarizes the definitions between the major grain type categories commonly used to describe carbonate rocks. Excellent colored thin-section photographs of grain types can be found in Adams and MacKenzie (1998). Subtypes and the controls on the formation of grains are discussed in the following paragraphs of this chapter. The nomenclatoral differentiation of carbonate grains developed from the study of modern carbonate environments in the 1960s and was rapidly augmented by facies studies of ancient carbonate rocks. Today, more than thirty useful and less useful terms are known, having been proposed for specific grain types that differ in real or inferred origin, morphology, internal structures and size. Microfacies analysis needs well-defined grain categories, but it should be kept in mind that a careful description is more important than a desperate search for an instant label for a grain.
Small micritic grains, commonly without internal structure. Subrounded, spherical, ovoid or irregular in shape. Size between <0.02 and about 1 mm, commonly 0.10 to 0.50 mm.
CORTOIDS
Rounded skleletal grains and other grains covered by a thin micrite envelope. Boundary between the central grain and the envelope indistinct. Size between <1 mm to a few centimeters.
ONCOIDS
Large and small grains consisting of a more or less distinct nucleus (e.g. a fossil) and a thick cortex formed by irregular, non-concentric, partially overlapping micritic laminae. Laminae may exhibit biogenic structures. No tendency to increase sphericity during growth. Size from <1 mm to a few decimeters.
OOIDS
Spherical or ovoid grains, consisting of smooth and regular laminae formed as successive concentric coatings around a nucleus. Laminae may exhibit tangential and radial microfabrics. Size between 0.20 and about 2 mm, commonly between 0.5 and 1 mm.
PISOIDS
Large subspherical and irregularly shaped grains, consisting of a mostly nonbiogenic nucleus and a thick cortex formed by conspicuously, often densely spaced laminae exhibiting tangential and radial microfabrics. Pisoids occur as isolated grains or are incorporated in crusts. Size generally >2 mm, up to >1 cm.
GRAIN AGGREGATES
Compound grains consisting of two or more originally separated particles (e.g. ooids, skeletal grains) that have been bound and cemented together, forming grape-like or rounded lumps. Intergrain spaces filled with micrite or spar. Outline irregular lobular or rounded. Size 0.5 to more than 2 mm.
CLASTS
Synsedimentary or postsedimentary lime clasts, reworked partly consolidated carbonate sediment or already lithified material. Shape and size are highly variable: angular to rounded. Size ranges between <0.2 mm and several decameters. Very small clasts are hardly distinguishable from peloids.
SKELETAL GRAINS
Fragmented or complete skeletons of organisms. Size from 0.05 mm to many centimeters.
Fig. 4.7. Descriptive terminology of the major categories of carbonate grains. Some of these grain types represent gradational stages: Cortoids may develop into oncoids; spherical isolated pisoids may resemble ooids structurally, and very small micritic clasts are often included within the peloid group.
Bioclasts (Skeletal Grains)
Significance of carbonate grains: Grain types are paleoenvironmental proxies both for non-marine and marine carbonates (e.g. water energy levels, sedimentation rates); characterize and differentiate specific depositional settings and sea-level fluctuations, and provide significant insights into global secular variations of carbonate mineralogy in Phanerozoic oceans. Grain association patterns are central to the reconstruction of paleoclimate and paleolatitudinal zones (Sect. 12.2.2.1). The ‘compositional maturity’ of limestones (the extent to which a sediment approaches the constituent endmembers: synsedimentary clasts, ooids, fossils, peloids, micrite, and terrigenous minerals) augment the complexity of the processes that operate on carbonate sediments (Smosna 1987). Changing grain composition reflects cyclic sedimentation and assists in evaluating of sequence stratigraphic models (Sect. 16.1.2.2). Grain type, mineralogy and spatial variations in the distribution of grains are major controls on the porosity development of reservoir rocks, as exemplified by Fig. 4.8.
4.2.1 Bioclasts (Skeletal Grains) Terminology Originally, the term bioclast was used for fossils which were transported, broken and abraded and which became part of the organic debris (Grabau 1920). Today, the term has a rather vague meaning in the context of microfacies studies, and is commonly used for fossils seen in thin sections, regardless of whether the fossils are fragmented (e.g. crinoid ossicles or broken shells) or still completely preserved (e.g. double-valved bivalves). The latter have been assigned to biomorpha. The term skeletal grain is synonymous with the term bioclast. Organoclasts are organic particles of various origin with or without structure. They include phytoclasts (Bostick 1971: clay- to fine sand-sized, land-derived plant remains), palynomorphs (HCl and HF resistant organic microfossils comprising cyanobacterial and algal remains, e.g. dinoflagellates and acritarchs; spores and pollen), degraded organic matter as well as zooclasts of organic microfossils (e.g. organic relicts of benthic foraminifera; scolecodonts: polychaete jaws). Phytoclasts and other organoclasts commonly are studied in palynofacies slides after extracting the fossils from sediments by chemical methods, but can rarely also be recognized in thin sections. The analysis of organoclasts found in Mesozoic and Cenozoic carbonate rocks is of increasing importance for understanding the relations between carbonate and terrestrial environments and provide very useful information on
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high-resolution sequence stratigraphy and paleoenvironmental interpretation of shallow-marine and deepmarine carbonates (Courtinat 1989; Steffen and Gorin 1993; Tyson 1995; Pittet and Gorin 1997; Rameil et al. 2000). Reworking of whole or fragmented skeletal grains (e.g. foraminifera, mollusks) from unconsolidated or semiconsolidated sediments is common in shallowmarine environments. Read (1974) termed those grains lithoskels and described modern examples from Shark Bay, Australia, where banks of skeletal grainstone have been formed by accumulation of grains eroded from Pleistocene sediments. Ancient lithoskels are common in lag deposits (Pl. 9/7). Sizes of bioclasts The average size of skeletal grains seen in thin sections varies widely, and for larger fossils depends on the size of thin sections used. The sizes of skeletal grains and fossils in the thin sections vary within a range of <1 mm to a few millimeters to several centimeters. The ordinary size ranges of skeletal grains are reported in the context of the discussion of the fossil groups in Chap. 10. The size of skeletal grains reflects growth size (e.g. tests of planktonic foraminifera), and size patterns; these are caused by transport processes leading to the diminution of fossils or fragments of fossils, and depend on hydraulic modifications. The shape and bulk density of bioclastic carbonate grains as well as the skeletal structure are important factors affecting the composition and grain size distribution of carbonate rocks (Maiklem 1968). The effect of the hydraulic modification of the initial grain size varies with the shape of the grain, whether blocks (corals, coralline algae), rods (larger foraminifera, turreted gastropods), plates (Halimeda segments, pelecypod valves), or spheres (rounded foraminifera). In experiments, plate-shaped grains settle significantly more slowly than any of the other grain types. Spheres also settle slowly. Block-shaped grains finer than about 2 mm settle slowly, grains coarser than 2 mm exhibit higher and irregular settling velocities. Rods of all sizes tend to settle quickly and smoothly with their long axis horizontal, if the interior is not partly filled with sediment. Constraints on the preservation of skeletal grains The record of carbonate bioclasts in limestones is strongly controlled by (1) the primary skeleton mineralogy, (2) taphonomic and diagenetic processes and (3) methodological biases. (1) Primary skeleton mineralogy and biomineralization: Skeletal grains consist of mineral phases and
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Preservation of Skeletal Grains
Fig. 4.8. Implication of carbonate grain studies. A statisticallybased microfacies analysis of core slabs of gas-producing carbonate reservoir rocks demonstrates the relationships between grain types, depositional subfacies, and porosity distribution. The subtypes can be recognized in cuttings as well. A: Presence and composition of grains are indicative of subfacies types that reflect specific environmental conditions and settings on a tidal to subtidal carbonate platform. Peloids and coated grains are the most reliable indicators of subtypes. Peloids are differentiated according to shape and size, ooids according to size, number of laminae and shape, and aggregate grains (including grapestones) according to internal structure and size. B: Intergranular and moldic porosity increases distinctly with increasing ooid size and abundance. The amount of algal content is negatively correlated with porosity. Late Permian (Zechstein Ca2 platform): Northwest Germany. Modified from Steinhoff and Strohmenger (1996).
organic phases (Mann 2001; Cuif and Sorauf 2001). Both phases are involved in biomineralization and affected by destructive processes. Today more than 60 biominerals occurring in bacteria, plants, protists and metazoans are known. About 50% of the more common skeleton-building biominerals are represented by carbonate phases. Other more common skeletonbuilding minerals are phosphates and silica (Fig. 4.9). Biomineralization follows two basic lines, characterized by biologically controlled or biologically induced processes. Bio-controlled deposition of minerals takes place in isolated spaces and in close association with an organic matrix (Degens 1979, 1989; Lowenstam and Weiner 1989; Simkiss and Wilbur 1989; Addadi and Weiner 1992). According to this model, a matrix of acidic macromolecules rich in aspartic and glutamic aminoacids acts as a binding sur-
A B
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Skeletal Mineralogy
Cyanobacteria Pyrrhophyta: Chrysophyta: Chlorophyta:
Rhodophyta:
Radiolaria Foraminifera Ciliata: Sponges:
Scyphozoa: Hydrozoa Corals:
Bryozoa Brachiopoda: Mollusca:
Tentaculitida Annelida: Arthropoda:
Echinodermata Tunicata Vertebrata Conodonts
Calciodinoflagellata Diatoms Coccolithophorida Dasycladaceae Udoteaceae Gymnocodiaceae Charophyceae Solenoporaceae Squamariaceae Corallinaceae
Aragonite
Low-Mg Calcite
High-Mg Calcite
o
● ●
o
Aragonite + Calcite
CaPhosphates
Silica
● ● ● ● ● ●
●
● ● ● ● o
Calpionellida Demospongea Calcarea Sphinctozoa Stromatoporoidea Chaetetida Archaeocyathida Hexactinellida Conulata Octocorallia Rugosa Heterocorallia Tabulata Scleractinia Articulata Inarticulata Monoplacophora Polyplacophora Scaphopoda Bivalvia Gastropoda Nautiloidea Ammonoidea Belemnoidea Serpulida Trilobita Ostracoda Cirripedia Decapoda
● ● o ● ● ● ● ●
● o ●
● ●
●
● ● ● o
o ● o
o o ● ● ●
o ● o
● ●
o o
o
o ●
o ●
● ● ● ● ● ● ●
●
● ● o ● Aptychus
● o o ●
● ●
o
● o (otoliths)
● o ● ● ●
● o
o
o ●
o ● ● ● ● ●
Fig. 4.9. Primary skeletal mineralogy of organisms relevant for microfacies studies. The dominating mineralogy is indicated by black circles, less common mineralogy by open circles. Note that bimineralic skeletons occur in several groups (e.g. bryozoans, arthropods). Adapted from Leadbeater and Riding (1986), Lowenstam and Weiner (1989), Mann et al. (1989), Carter (1990), and Van de Poel and Schlager (1994).
face for Ca2+, thus controlling the precipitation of CaCO3. The controlled, organic matrix-mediated mineralization is responsible for the bulk of the Phanerozoic fossil record. This biomineralization type appears in the Late Precambrian-Early Cambrian interval. Bioinduced formation of minerals results from the activi-
ties of organisms on their external environment with no necessary specific interaction with organic matrices. Primary examples are the precipitation of calcium carbonate induced by metabolic processes of bacteria (Sect. 4.1.1), the formation of iron sulfides as products of sulfate-reducing bacteria, or the precipitation of ara-
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gonite crystals in spaces between cells due to the removal of carbon dioxide in the process of photosynthesis of algae. Bacterially-induced carbonate formation was already abundant early in the Proterozoic at least 2.4 billion years ago. There is a growing body of data on biomineralization processes that act in the formation of carbonate hardparts of invertebrates and algae. The current state of the art is documented in the books listed at the end of this subchapter. Conflicting interpretations of carbonate biominerals relate to the functional forms, the interactions between matrix and crystals and between macromolecules and crystals (Mann 2001). A useful glossary of skeletal biomineralization was made by Bandel et al. (1990). (2) Taphonomic and diagenetic biases: Becoming a fossil is a rather difficult task. The processes that lead to fossil destruction and fossil preservation are summarized with the term ‘taphonomy’. Process-oriented taphonomic studies are a major tool of modern geobiological analyses. Microfacies interpretations should rely strongly on the results of taphonomic studies. Some basic references are listed at the end of this subchapter. In marine environments, carbonate hardparts are degraded through physical abrasion, fragmentation and comminution, infestation of microendoliths, bioerosion by grazing and predation, micritization, syndepositional recrystallization related to the decomposition of organic tissue, and biologically and/or chemically-induced dissolution (Bathurst 1966; Freiwald 1995; Perry 1998; Cuif et al. 1999; Peterhänsel and Pratt 2001). Breakdown is most rapid at the sediment-water interface and in the surficial sediments, where dissolution is fastest and microborers most abundant. Dissolution and bioerosion rates decrease rapidly with depth in the sediment. Bioturbation may speed up or slow down the degradation of carbonate grains (Bradshaw and Scoffin 2001). The breakdown of carbonate hardparts and skeletal grains is controlled by primary mineralogy, ultrastucture, crystal size, shape and size, the porosity and permeability of the skeletons, the resistance of organic coatings against decomposition as well as shell strength (Zuschin and Stanton 2001) and the morphology of the skeleton. The dissolution rate of biogenic grains is inversely related to grain diameter. These factors and the rates and modes of burial influence the physical and chemical behavior of hardparts and control the record of skeletal grains in shallow and deep marine environments (Force 1969; Walter and Morse 1984: Henrich and Wefer 1986).
Taphonomy
Several morphological types of skeletons or tests can be distinguished on the basis of their resistance to fragmentation and the resulting skeleton grains (Wilson 1975): 1 Needles embedded in organic tissues yield fine scattered particles. Examples are sponges: (Pl. 78/3), octocorals, holothurians, tunicates and green algae, e.g. Penicillus. 2 Segmented robust skeletons linked together by organic tissue will suffer intact transport of parts of the skeleton, but in the end yield sand-sized and larger particles (green algae, e.g. Halimeda: Pl. 149/ 3; articulated red algae, bryozoans, echinoderms: Fig. 10.55; trilobites: Pl. 94/1). 3 Branched skeletons composed of blade-like or cylindrical elements yield non-equant fragment sizes (some corals, bryozoans: Pl. 106/3, red algae and phylloid algae) 4 Chambered hollow skeletons are relatively resistant to disintegration (coralline sponges: Pl. 79/5, gastropods, serpulids: Pl. 92/2; brachiopods, some bivalves and crustaceans, foraminifera, dasyclad green algae: Pl. 61/1). 5 Encrusted skeletons are ususally rather resistant to mechanical breakup (sessile foraminifera, stromatoporoids, corals, byozoans, serpulids, red algae). 6 Massive skeletons are most resistant to fragmentation mainly because of their large size and their moundlike shape (e.g. colonial corals: Pl. 84/1 stromatoporoids: Pl. 80/7). Calcareous skeletal microstructures can be grouped into ten major categories (Box 4.6) whose criteria are better revealed by SEM magnification than by lightoptical studies. The definition of microstructure types differs considerably due to different techniques and magnifications used to study the basic microstructural elements. For instance, diagenetic alterations of primary structures frequently are overlooked in thin section. Thin-section microphotographs of common microstructures can be found in Horrowitz and Potter (1971) and Majewske (1974), SEM photographs of skeleton ultrastructures in the biomineralization atlas of Carter (1990). Taphonomic loss resulting in a loss of information, is caused by the transport of hardparts and skeletal grains and by time-averaging (Allison and Briggs 1991; Kidwell and Behrensmeyer 1993; Martin 1999; Behrensmeyer et al. 2000). During transport processes, skeletal grains can loose their original shape and become increasingly worn and rounded. Univalved shells
Microstructures of Invertebrate and Algal Skeletons
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Box 4.6. Major microstructure categories occurring in invertebrate and algal skeletons (after Wendt 1990 and Majewske 1974). The organic matrix within the hardparts is an important control on the preservation of microstructures and the post-mortem destruction (Glover and Kidwell 1993; Cuif et al. 1999). Organic-rich microstructures (OR) of mollusks are in fact commonly preserved only as molds; organic-poor microstructures (OP) are often fairly well-preserved. 0-dimensional Microgranular: Irregularly interlocked anhedral grains without preferred orientations. Aragonite or calcite. Common in cyanobacteria (Pl. 7/6) and calcareous algae, foraminifera (Pl. 68), calpionellids, serpulids; uncommon in mollusks; rare in corals and coralline sponges (Pl. 79/1, 6). 1-dimensional: Thin, elongate crystals with length/width ratios between 1:5 and 1:25. Irregular fibrous: Randomly oriented individual fibers. Aragonite or calcite. Foraminifera, coralline sponges, ostracods, trilobites (Pl. 94/3), mollusks. Orthogonal (fibro-normal): Crystals oriented parallel to each other and perpendicular to the surface of the skeletal elements. Aragonite or calcite. Calcareous algae, foraminifera, coralline sponges, rugose and tabulate corals, ostracods, trilobites, mollusks. Clinogonal (trabecular): Crystals diverging at low angles from a central axis or several aligned centers. Usually aragonite, rarely calcite. Coralline sponges, hydrozoans, and rugose, tabulate and scleractinian corals. Spherulitic: Spherical or hemispherical aggregates of crystals, centrically or excentrically arranged. Usually aragonite, very rarely calcite. Coralline sponges, hydrozoans, octocorallia; very rare in bivalves. 2-dimensional: Planar or curved blades which may be composed of subcrystals with the same optical orientation. Lamellar (foliated): Very thin lamellae regular or more or less parallel to surface of skeletal elements, or arranged in stacks of slightly different optical orientation. Mostly calcite, rarely aragonite. Hydrozoans, rugose and tabulate corals, serpulids (Pl. 92/6, 8), brachiopod shells and spines (Pl. 86/1, 6), tentaculitids (Pl. 91/7), bryozoans (Pl. 85/4), gastropods, bivalves (OP). Nacreous: Thin, organic-rich tablets composed of subcrystals of the same optical orientation, oriented parallel to the surface of skeletal elements or growth lamellae. Aragonite. Found only in mollusks (inner layer of certain bivalve and gastropod shells; major parts of cephalopod shells; OR). Crossed lamellar: Thin interfingering or branching first order lamellae composed of parallel-arranged second order lamellae. The optical orientation of the first-order lamellae alternates. Commonly aragonite, rarely calcite. Found only in mollusks (OP; bivalves, gastropods: Pl. 149/5; hyolithids). 3-dimensional: Large crystals (about 0.1 mm to several cm in size). Normal prismatic: Polygonal prisms arranged parallel to each other and perpendicular or oblique to the surface of skeletal elements. Each prism is an individual crystal with unit extinction. Usually calcite, rarely aragonite. Inner carbonate layer of most punctate and impunctate brachiopods, outer carbonate layer of bivalves (Pl. 87/2, 5, 8; OR) and gastropods, rostra of belemnites (Pl. 90/7), ostracods (Pl. 93/2). Optical uniform: Morphologically well-defined skeletal elements or whole shells appear to be single calcite crystals in the light microscope. Calcite cleavage visible under ordinary light. Typical of echinoderms (Pl. 95/3) and spicules of some calcareous sponges. Extremely rare in foraminifera and bivalves.
(gastropods, cephalopods) and thin bivalved shells (brachiopods, bivalves, ostracods) are highly subject to intact transport of the skeleton, whereas multi-element, tightly sutured (echinoids, some crinoids) and loosely articulated skeletons (most crinoids, trilobites) are subject to disarticulation and fragmentation (Brett and Baird 1986). Out-of-habitat transport of fossils is commonly associated with settings affected by high-energy currents or steep depositional slopes. Time-averaging is the mixing of skeletal elements of non-contemporaneous organisms (Walker and Bambach 1971). Mollusks and benthic foraminifera from near-shore marine settings are time-averaged over decades to hundreds of years, and organisms living on carbonate platforms and in reefs over a time scale of
decades to ten thousands of years (Kidwell and Behrensmeyer 1993). A model showing the distribution of the relative abundance of the types of time-averaging in different marine environments was developed by Fürsich and Aberhan (1990). However, despite all these constraints, death assemblages of resistant skeletal elements are excellent records of environmental distribution and community composition for many major groups (Kidwell and Flessa 1996). Many skeletal grain types have specific preservation potentials within a given depositional environment, depending on the susceptibility of grains to mechanical destruction, duration of deposition/exhumation/re-
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distribution processes, and early dissolution of grains preceding cementation. Brachert et al. (1996) refer to this particular type of microfacies as microtaphofacies. These microfacies types reflects the taphonomic history of bioclast rather than original distribution patterns. The microtaphofacies concept provides important insights into the controls on microfacies of carbonate ramps and platforms by sealevel fluctuations (Sect. 16.1.2.2). Major preservation biases are caused by dissolution processes involved in skeletal grain diagenesis (Chave 1964): Skeletal diagenesis is a matter of multiple choices in space and time, as exemplified by the diagenetic sequence patterns of corals (Schroeder 1984; see Pl. 35). Because skeletal durability relates strongly to mineralogy and ultrastructure, relative diagenetic preservational orders of skeletal structures are common, and can be used to estimate diagenetic changes. The relative order of preservation in meteoric environments is calcite > aragonite > High-Mg calcite. High-Mg calcite is unstable in freshwater and can easily transform to Low-Mg calcite by incongruent leaching or only little change in microstructures and microfabrics (Bathurst 1971). The process of magnesium loss during stabilization of Mg-rich calcites and the effect of that loss on the resulting calcite microfabric are not well understood (Towe and Hemleben 1976; Dreifuss 1977; Oti and Müller 1985). Some of the altered HCM parts show textural alterations, others do not. The characteristically finer size distribution and the relic-free microfabrics resulting from HCM stabilization are distinctly different from the coarse, pitted, relic-bearing calcites which replace originally aragonitic constituents. The interpretation of diagenetic alterations in fine-grained calcitic skeletal grains (porcelaneous foraminifera: Budd and Hiatt 1993; echinoderms: Richter 1979, 1984) is also complicated by the level of observation since the resulting ‘grain diminution’ (Sect. 4.1.2) is difficult to substantiate in thin sections. Aragonite skeletons (mollusks, corals, green algae) are replaced by calcite (Bathurst 1964; Talbot 1972), and have a consistent alteration behavior (Sandberg 1984). The replacement crystals are one or more orders of latitude larger than the crystals they replace. The original microstructures are commonly not preserved. The transformation of aragonite skeletons into calcite bioclasts takes place in meteoric and phreatic environments by cementation after dissolution of aragonite via a mold stage (Pl. 149/5), by in situ calcification and micro- to macrorecrystallization connected with a solution film between aragonite and replacing
Utility of Different Methods
calcite (Pingitore 1976), and by biogenic micritization. Aragonitic bioclasts were likely prone to sea floor dissolution, but preservation of Quaternary aragonitic pteropods is reported from the deep continental slope. Generally, preservation of aragonitic fossils in carbonate rocks requires very specific depositional and diagenetic conditions (Füchtbauer and Goldschmidt 1964; Scherer 1977). The selective diagenetic changes in invertebrate skeletal grains were investigated by experimental dissolution of recent skeletal carbonates, particularly of mollusks, bryozoans, echinoderms, corals and calcareous algae (Walter 1985). These studies and the study of the microarchitecture of hardparts show the dominating role of the ultrastructure for the relative reactivity of skeletal carbonates during dissolution and preservation (Dullo 1987). The impact of the diagenetic loss of bioclasts on the composition of carbonate rocks has been thoroughly studied by comparison of modern and Pleistocene carbonates affected by meteoric diagenesis (Land 1967), and is exemplified by the ‘diagenetic sieve’ (Dullo 1984; Fig. 4.10). Initial diagenetic alterations of the hardparts of marine organisms may occur within a short time (Flessa and Brown 1983). The diagenetic thickening of skeletal elements of recent corals takes place within a time range of twenty years, prior to the alteration of aragonite to calcite (Potthast 1992). Selective dissolution and replacement of aragonitic bioclasts continues in shallow-marine burial environments (Brachert and Dullo 2000). In addition to all the processes described above, dolomitization (Murray 1964), silicification, burial diagenesis and compaction can contribute strongly to the obliteration of skeletal grains and to their complete loss in the geological record (see Pl. 36/1, 3, 4; Pl. 37/7, Pl. 47/4, Pl. 109/8). (3) Different methods can lead to different results: Thin sections of limestones do not document the total record of skeletal grains. Some very small skeletal grains may be missed. Agglutinated or silicified microfossils may be underrepresented. Toomey (1983) while studying a Late Paleozoic shelf limestone in thin sections and acid residue samples, demonstrated that only about 80% of agglutinated foraminifera could be recorded in the acid residue samples. Some organism groups are almost always underrepresented or missing in thin sections when compared with acid residues of limestones (e.g. agglutinated foraminifera, ostracods, echinoderms, scolecodonts, conodonts, fish remains). One example are holothurians
Influence of Diagenesis
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Fig. 4.10. The ‘diagenetic sieve’ (after Dullo 1990). The study of modern submergent reefs and Pleistocene emergent reefs allows the stages and timing of the diagenetic alterations of skeletal grains to be recognized (e.g. James 1974; Scherer 1975; Gvirtzman and Friedman 1977; Vollbrecht 1990). Comparison of the biotic composition of recent reefs off the Red Sea Coast with a strongly altered reef on the Pleistocene terraces, occurring in different levels of elevation above the present sea, exhibits progressive diagenetic alterations of bioclasts. These alterations result in the partial to complete leaching of primary aragonitic biota, a gradual decrease in Mg-calcitic hardparts, concommitent processes of replacement and recrystallization, and the precipitation of carbonate cements. Aragonitic bioclasts (scleractinians, hydrozoans, aragonitic mollusks) and Mg-calcitic bioclasts (echinoderms, foraminifera, coralline calcareous algae) are drastically reduced in the geological record by meteoric diagenesis. The Pleistocene reef limestone yields a higher amount of unidentified bioclasts, peloids, carbonate cement and pores than modern reefs. The examples demonstrate the significant modification of biotic composition and the frequency of skeletal grains in limestones by selective fossil diagenesis. The degree of progressive diagenesis within the reef terraces is proportional to the age of the Pleistocene reefs, which varies between 86 000 and about 340 000 years.
which occur in high abundance in acid residues of basinal limestones, but are extremely rarely observed in thin sections. A combination of thin section studies and the investigation of acid residues is strongly recommended. These differences are important for the evaluation of ‘fossil diversity in thin sections’ (Sect. 6.2.1.4). The advantages and disadvantages of two-dimensional versus three-dimensional aspects of skeletal grains are illustrated in Pl. 9/1-6. Significance of bioclasts Skeletal grains are the most valuable grain types in the context of determining the age of limestone samples. The type and composition of skeletal grains is highly sensitive to the depositional environment and offer significant proxies for paleoenvironmental controls on carbonate depositionand paleoclimatic conditions. Under favorable conditions, the diagenesis of skeletal grains can document sealevel fluctuations as exemplified by the study of Pleistocene beach sediments in Bermuda Island (Vollbrecht and Meischner 1993). The diagenesis of sediments is reflected in the diagenesis of fossils. Because the original mineralogy of most calcareous fossils is known or can be inferred, fossil diagenesis is a perfect tool for reconstructing diagenetic pathways.
Basics: Preservation of skeletal grains (bioclasts) Adams, A.E., MacKenzie, W.S. (1998): A colour atlas of carbonate sediments under the microscope. – 180 pp., London (Manson) Majewske, O.P. (1974): Recognition of invertebrate fossil fragments in rocks and thin sections. – International Sedimentary Petrographical Series, 13, 101 pp., Leiden (Brill) Peryt, T.M. (ed., 1983): Coated grains.– 655 pp., Berlin (Springer) Richter, D. (1983): Classification of coated grains: discussion. – In: Peryt, T. (ed.): Coated grains. – 7-8, Berlin (Springer) Steinhoff, I., Strohmenger, C. (1996): Zechstein 2 carbonate platform subfacies and grain-type distribution (Upper Permian, Northwest Germany). – Facies, 35, 105-132 Biomineralization Carter, J.G. (ed., 1990): Skeletal biomineralization patterns, processes and evolutionary trends. Vol. I. – 832 pp., New York (Van Nostrand) Degens, E.T. (1979): Why do organisms calcify? – Chem. Geol., 25, 257-269 Degens, E.T. (1989): Perspectives in biogeochemistry. – 423 pp., Berlin (Springer) Leadbeater, B.S.C, Riding, R. (eds., 1986): Biomineralization in lower plants and animals. – Syst. Ass. Spec. Vol., 30, 308 pp., Oxford (Clarendon Press) Kempe, S., Kazmierczak, J. (1994): The role of alkalinity in the evolution of ocean chemistry, organization of living systems and biocalcification prozesses. – In: Doumenge, F. (ed.): Past and present biomineralization processes. –
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Skeletal Grains
Plate 9 Skeletal Grains (Bioclasts): ‘Thin-Section Fossils’ and Taphonomic Pitfalls Skeletal grains are the prime source of carbonate sands and muds deposited in marine environments. The value of bioclasts in facies analysis depends largely on the level of systematic assignment. Determination of ‘thinsection fossils’ at a genus or species level depends on the paleontological background of the student and her or his ability to reconstruct a three-dimensional picture from random cuts. Reconstruction and taxonomic attribution is facilitated by comparison of ‘thin-section fossils’ with ‘acid-residue’ fossils (–> 1–6). The joint use of thin-section and acid-residue techniques is strongly recommended – last but not least because thin-sections reflect only parts of the total fossil record of the limestone sample. The poor preservation of taxonomically important criteria caused by taphonomic and diagenetic processes (–> 7) is responsible for generalized taxonomic assignments (e.g. ‘limestone with gastropods and foraminifera’). Two- and three-dimensional aspects of fossils. Selective silicification of a Late Pennsylvanian shelf limestone (Carnic Alps, Austria) allows of thin-section and acid-residue data to be compared (Fohrer 1991). 1 Thin section: Median section of the microgranular calcareous foraminifera Deckerella Cushman and Waters (Palaeotextulariacea, Fusulinina). Genus criteria: Elongate test, short and wide chambers, biserially and later uniserially arranged. Microgranular wall dark in transmitted light; exterior part of the wall agglutinated. Aperture of later stages with parallel slitlike openings. The thin section provides all the information necessary for generic determination. The genus is restricted to the Carboniferous. 2 Acid residue of a specimen of Deckerella. The morphology of the test and the arrangement of chambers are shown more distinctly than in the thin section. 3 Thin section: Oblique axial section of the dasyclad green alga Anthracoporella Pia (compare Pl. 59/6). ‘Stem cell’ filled with sediment; the thick wall exhibits tangential, longitudinal and oblique sections of pores (branches). The arrow points to encrusting foraminifera. 4 Acid residue of Anthracoporella. Shape and arrangement of pores are more obvious than in the thin section. Quantitative species-diagnostic data (e.g. pore diameter) are different from thin-section data. 5 Thin section: Fenestrate bryozoan exhibiting the obverse side of the net-like colony, fenestrules (F) and cross sections of zooecia apertures (ZA); compare Pl. 85/10. These criteria allow them to be designated only as ‘fenestrate bryozoa’. 6 Acid residue.The taxonomically important arrangement, size and shape of the zooecia are more obvious than in thin sections and allow them to be assigned to the genus Fenestella. Measurements can be done more exactly than in thin sections. Taphonomic and diagenetic controls on the preservation of skeletal grains. Essential criteria for describing the ‘microtaphofacies’ of this sample are abrasion, fragmentation, bioerosion and encrustation. Microtaphofacies refers to the taphonomic biasing caused by displacement, redeposition and mixing of grains as well as selective dissolution which in turn is controlled by the primary skeletal mineralogy. 7 The limestone conglomerate (Early Tertiary, Eocene: Croatia) consists of fossils (gastropods - G, bivalves - P, foraminifera, dasyclad algae), micritic limestone pebbles (MLP), peloids and small reworked and rounded micrite clasts (MC). Gastropods and alveolinid foraminifera exhibit different preservation due to different primary mineralogy: The aragonitic gastropod shells are dissolved, but the shape of the test is still preserved. Molds are infilled with sediment and microfossils. In contrast, the porcellaneous calcitic tests of the miliolid foraminifers Alveolina (A) and Orbitolites (O) exhibit no diagenetic changes. Low-Mg calcite is less susceptible to dissolution than aragonite or High-Mg calcite. The example exhibits a common relative order of bioclast preservation within an environment affected by oscillating marine, phreatic and vadose diagenetic processes: Low resistance of primary aragonitic skeletal grains (e.g. gastropods, scleractinian corals, dasyclad algae) to dissolution, and higher resistance of primary calcitic grains represented by miliolid foraminifera, coralline algae, and fibrous lamellar bivalve shells (Gindy 1987; Bathurst 1994). Taphonomic history of gastropods: Reworking and redeposition of the ‘steinkerns’ (lithoskels) took place subsequent to dissolution of aragonitic shells and mold formation together with the deposition of well-rounded micritic limestone pebbles (MLP). Interstices between larger skeletal grains and pebbles were filled with alveolinid and other foraminifera, dasyclads and small bioclast debris. Gastropod molds were encrusted by foraminifera (F) and spirorbid worms (W). Other organisms settling on hard substrates are barely recognizable barnacles (B), compare Pl. 93/6, 7. Borings of shells are evident. Meteoric dissolution is recorded by irregular solution voids (SV) filled with silt. Temporal sequence of events: Reworking of back-reef biota (indicated by Alveolina and Orbitolites) and deposition of molds of aragonitic shells, pebbles and micrite clasts in a high-energy environment –> encrustation and boring of larger fossils –> cementation and meteoric dissolution –> deposition of foraminiferal sand. The mixed deposition of worn fossils, allochthonous lithoclasts and blackened peloids as well as strong encrustation and boring characterize a ‘lag deposit’, formed in an environment with slow accumulation of coarse material within a winnowing zone (Standard Microfacies Type: SMF 14). This is the case in coastal zones which are prone to rapid changes in sedimentation rates, destruction, dissolution and renewed marine deposition. –> 1–6: Fohrer 1991
Plate 9: Skeletal Grains
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Bulletin Institute de Océanographie Monocao, Special Issue, 13, 61-117 Lowenstam, H.A., Weiner, S. (1989): On biomineralization. – 324 pp., New York (Oxford University Press) Mann, S. (2001): Biomineralization - principles and concepts in bioinorganic materials chemistry. – 210 pp., Oxford (Oxford University Press) Mann, S., Webb, J., Williams, R.J.P. (eds., 1989): Biomineralization. Chemical and biochemical perspectives. – 541 pp., Weinheim (VCH Verlag) Simkiss, K., Wilbur, K.M. (1989): Biomineralization. Cell biology and mineral deposition. – 337 pp., San Diego (Academic Press) Towe, K.M. (1972): Invertebrate shell structure and the organic matrix concept. – Biomineralization Research Reports, 4, 2-14 Westbroek, P., de Jong, E.W. (eds., 1983): Biomineralization and biological metal accumulation. – 533 pp., Dordrecht (Reidel) Further reading: K043 Taphonomy Allison, P.A., Briggs, D.E.G. (1991): Taphonomy: releasing the data locked in the fossil record. – 560 pp., New York (Plenum) Behrensmeyer, A.K., Kidwell, S.M., Gastaldo, R.A. (2000): Taphonomy and paleobiology. – In: Erwin, D.H., Wings, S.L. (eds.): Deep time. Paleobiology’s perspective. – Paleobiology Supplement, 26/4, 103-147 Brett, C.E., Baird, G.C. (1986): Comparative taphonomy: a key to paleoenvironmental interpretation based on fossil preservation. – Palaios, 1, 207-227 Donavan, S.K. (ed., 1991): The process of fossilization. – 303 pp., New York (Columbia University Press) Fürsich, F.T., Aberhan, M. (1990): Significance of time-averaging for paleocommunity analysis. – Lethaia, 23, 143152 Kidwell, S.M., Behrensmeyer, A.K. (eds., 1993): Taphonomic approaches to time resolution in fossil assemblages. – Short Courses in Paleontology, 6, 302 pp., Knoxville (Paleontological Society) Kidwell, S.M., Flessa, K.W. (1996): The quality of the fossil record: populations, species and communities. – Annual Review of Earth and Planetary Sciences, 24, 433-464 Maiklem, W.R. (1968): Some hydraulic properties of bioclastic carbonate grains. – Sedimentology, 10, 101-109 Martin, R.E. (1999): Taphonomy. A process approach. – 524 pp., Cambridge (Cambridge University Press) Murray, R.C. (1964): Preservation of primary structures and fabrics in dolomite. – In: Imbrie, J., Newell, N.D. (eds.): Approaches to paleoecology. – 388-403, New York (Wiley) Perry, C.T. (1998): Grain susceptibility for the effects of microboring: implication for the preservation of skeletal carbonates. – Sedimentology, 45, 39-51 Further reading: K080 Fossil diagenesis Bathurst, R.C.G. (1964): The replacement of aragonite by calcite in the molluscan shell wall. – In: Imbrie, J., Newell, N. D. (eds.): Approaches to paleoecology. – 357-387, New York (Wiley) Brachert, T.C., Dullo, W.-Chr. (2000): Shallow burial diagenesis of skeletal carbonates: selective loss of aragonite shell material (Miocene to Recent), Queensland Plateau and
References on Skeletal Grains
Queensland Trough, NE Australia) - implications for shallow cool-water carbonates. – Sedimentary Geology, 136, 169-187 Chave, K.E. (1964): Skeletal durability and preservation. – In: Imbrie, J., Newell, N.D. (eds.): Approaches to paleoecology. – 377-387, New York (Wiley) Cherns, L., Wright, V.P. (2000): Missing mollusks as evidence of large-scale, early skeletal aragonite dissolution in a Silurian sea. – Geology, 28, 791-794 Dullo, W.-Chr. (1983): Fossildiagenese im miozänen Leithakalk der Paratethys von Österreich: Ein Beispiel für Faunenverschiebungen durch Diageneseunterschiede. – Facies, 8, 1-112 Dullo, W.-Chr. (1990): Facies, fossil record, and age of Pleistocene reefs from the Red Sea (Saudi Arabia). – Facies, 22, 1-46 Flessa, K.W., Brown, T.J. (1983): Selective solution of macroinvertebrate calcareous hard parts: a laboratory study. – Lethaia, 16, 193-205 Glover, C.P., Kidwell, S.M. (1993): Influence of organic matrix on post-mortem destruction of molluscan shells. – J. Geol., 101, 729-747 Gvirtzman, G., Friedman, G.M. (1977): Sequence of progressive diagenesis in coral reefs. – Studies in Geology, 4, 357380 Henrich, R., Wefer, G. (1986): Dissolution of biogenic carbonates: effects of skeletal structure. – Marine Geology, 71, 341-362 Pingitore, N.E. (1976): Vadose and phreatic diagenesis: Progress products and their recognition in corals. – Journal of Sedimentary Petrology, 46, 985-1006 Richter, D.K. (1979): Die Stufen der meteorisch-vadosen Umwandlung von Mg-Calcit in Calcit in rezenten bis pliozänen Biogenen Griechenlands. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 158, 277-333 Further reading: K042, K044
4.2.2 Peloids: Just a Term of Ignorance? Many marine and non-marine limestones contain micron- and millimeter-sized, more or less structureless subrounded micritic grains, that are commonly called peloids or pellets (Pl. 10/1-3). Although the term peloid is often used as a nongenetic term of ignorance, the genetic differentiation of these grains offers important informations on the depositional setting, paleoenvironmental conditions and diagenetic features. Terminology Peloid (McKee and Gutschick 1969) is a comprehensive descriptive term for polygenetic grains composed of micro- and cryptocrystalline carbonate. Peloids are commonly devoid of internal structures but may contain fine-grained skeletal debris and other grains. The term was proposed in order to replace the widely used term ‘pellet’ which, for many authors had become a synonym for ‘fecal pellet’.
Modern Peloids
The name pellet should be restricted to grains of fecal origin; for all the other micrite grains the term ‘peloid’ is recommended. Many peloids are subrounded or rounded, but ovoid and rodlike shapes may also occur. The dimensions of the silt- to fine-sand-sized particles vary within a range of a few µm to a few mm, but most calcareous peloids are rarely larger than 500 µm, and commonly exhibit a diameter of < 200 µm, often 30 to 100 µm. Comparison with other grain types Peloids differ from ooids and oncoids by the absence of concentric or radial internal structures, and from small rounded intraclasts by their uniform shape, good sorting and small size (<200 µm). Peloids are generally smaller than ooids, pisoids and oncoids. Occurrence Peloids occur in rock-building abundance forming peloid limestones (Pl. 121/1-3) or as only minor constituents of carbonate rocks together with other grain types (Pl. 122/1). Peloids constitute grain- or mud-supported fabrics (contributing to wackestone, packstone, grainstone and bindstone textures), are found as isolated or amalgamated grains (Pl. 50/2), and occur within layered (Pl. 50/2, 5) and clotted textures (Pl. 10/1). Peloids are common in shallow-marine tidal and subtidal shelf carbonates and in reef and mud mounds, but are also abundant in deep-water carbonates. By contrast to the abundance of peloids in tropical shallow-marine carbonate, peloids are rare or absent in non-tropical coolwater carbonates. Modern peloids Marine peloids are conspicuous constituents of lowenergy shallow tidal and subtidal platform carbonates (Pl. 3/1) but are also found in sheltered reef cavities, and slope and basinal settings. The Great Bahama Bank is a good example of the importance of peloids for the makeup of shallow-marine mud-rich carbonates (Imbrie and Purdy 1962; Purdy 1963; Neumann and Land 1975; Harris 1979). The ‘pellet-mud facies’ and ‘mud facies’ of the restricted interior platform, originating in very shallow waters with increased and fluctuating salinities and covering an area of about 10 000 km2, comprises abundant peloids forming up to 75% of the total sand fraction. Most pellets are lithified organic excrements (fecal pellets) produced by carbonate ingesting organisms, particularly by polychaete worms, gastropods and some crustaceans. Open interior parts of the platforms are floored by peloid and aggregate grain sands. These peloids result from the micritization of various carbonate grains (‘bahamite peloids’). Other
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sites with abundant peloids are the Persian Gulf and the sea off Egypt (Stoffers et al. 1980). Peloids are also generated within algal/cyanobacterial mats; examples are known from sea-marginal hypersaline pools along the shores of the Gulf of Aqaba, the Red Sea, and Canary Island (Friedman et al. 1972; Gerdes and Krumbein 1987). The ‘peloidal question’: What are the genetic subtypes of ancient peloids? The origin of ancient peloids is often in doubt and the attempt to distinguish them is limited to recognizing of peloids of fecal, detrital or in situ precipitated origin (MacIntyre 1985). There are, however, some clues for better genetic specification of peloids. Several terms have been suggested for characterizing genetic subcategories. Shape, size and sorting as well as the composition of the grain associations are diagnostic criteria for interpreting the origin of peloids. The following peloid subcategories are used in microfacies studies (Fig. 4.11). Peloids of biotic origin: (1) Fecal pellets (Fig. 4.12; Pl. 136/8): Rounded, usually fine-grained particles caused by organisms that eat mud, digest organic matter from the mud and excrete the non-digested lime-mud. It should be kept in mind that many authors still use the term ‘pellet’ and not ‘fecal pellets’ for peloids produced by carbonate mud-ingesting organisms. Producers of fecal pellets in shelf environments are deposit- and detritus-feeders, herbivorous animals and plankton-feeders, including arthropods, mollusks (intertidal gastropods and polyplacophorid chitons), worms, echinoderms, and fishes. Deposit-feeding crustaceans, gastropods, holothurians, and worms are quantitatively important. Callianassa, a shallow-marine thalassinid decapod crab, deposits as much as 4.5 mm/ year of fecal pellet mud. The assignment of fossil fecal pellets to specific pellet-producing animals is difficult because fecal pellets of most invertebrates lack specific morphological criteria. Exceptions are gastropod and crustacean fecal pellets, which are characterized by specific dimensions and morphological criteria. Decapod fecal pellets are pierced by longitudinal canals whose number, size and outline in transverse sections are used for taxonomic classification (Favreina, Pl. 93/ 5, Sect.10.2.4.7). Fecal pellets accumulate not only in modern tropical and temperate intertidal and subtidal settings, but also in deep-marine environments where planktonic shells (e.g. coccolithophorids) are transported to the sea bottom via fecal pellets of zooplanktonic and nektonic animals (see Fig. 4.4).
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Peloids: Diagnostic Criteria
In-situ Formation
Alteration of Grains
Reworking of Mud and Grains
Biotic Origin
Origin
Types
Diagnostic criteria
Lithified organic excrements
1 Fecal pellets
Rounded elongated, rod-shaped or ovoid dark-colored micritic grains, rarely spherical. Commonly homogeneous or with silt-sized inclusions; rarely with defined internal structures. Sizes <100 µm to several millimeters. Sometimes associated with bioturbation structures . . . . . . . . . . . . . . . . . . . . Pl. 136/8
Abrasional products of algae and calcimicrobes
2 Algal peloids
Irregularly shaped, rounded micritic grains, exhibiting gradations from grains with relicts of algal structures to homogeneous grains. Size <20 µm to ~2 mm . . . . . . . . . . . Pl. 136/4
Grains resulting from hardpart-boring and rasping activity of organisms
3 Bioerosional peloids
Scoop-shaped subrounded and angular grains. Sizes from 20 µm to 100 µm
Synsedimentary and postsedimentary reworking of carbonate mud and micrite
4
Mud peloids (Lithic peloids)
Variously shaped micritic grains, commonly without internal structures. Wide size ranges, poor sorting. Frequent occurrence within distinct beds or laminae . . Pl. 10/2, Pl. 121/2
Internal micritic molds of bivalved shells
5 Mold peloids
Ovoid micritic grains, sometimes with relicts of still undissolved shells (ostracods, small bivalves). . . . . . . Pl. 132/8
Ooids and rounded skeletal 6 grains whose microstructures have been lost through micritization
Bahamite peloids
Round micritic grains, some of which with relicts of the primary microstructures. Association of peloids, aggregate grains and ooids. Transition of micritized bioclasts to peloids of the same size. Larger than algal peloids . . . . . Pl. 10/3, Pl. 43/1
Ooids and skeletal grains; microstructures destroyed by recrystallization
7
Pelletoids
Microcrystalline grains, in places exhibiting vague residual internal structures. Diffuse outlines due to amalgamation and compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pl. 38/6
Biochemical precipitation triggered by microbes and organic substances
8
Microbial peloids
Rounded micritic grains associated with laminated and clotted fabrics. Sizes from <80 µm to >600 µm . . . . . . Pl. 8/6, Pl. 10/1
Chemical precipitation of carbonate cements with or without organic controls
9
Precipitated peloids
Tiny peloids within carbonate cements; consisting of a cloudy micritic center surrounded by clear exterior rims of crystals. Occurrence in cavity fill precipitates (e.g. in reefs) . . . . Pl. 8/5
Fig. 4.11. Origin and diagnostic criteria of carbonate peloids. The term ‘peloid’ is only descriptive until thin-section studies reveal information providing a genetic connotation.
Criteria: Fecal pellets are commonly elongate, rodshaped or ovoid grains. Most fecal pellets are homogeneous, but some may contain silt-sized inclusions (quartz or skeletal debris). The dark color of the grains or the peripheral rim in transmitted light is caused by the high content of organic matter or iron sulfides. Fecal pellets occur in isolated nests and are common in burrowed limestones. Parautochthonous occurrences are often characterized by small, moderate to wellsorted pellets. The sorting of fecal pellets provides a first approximation of the hydrodynamic behavior of these grains (Wanless et al. 1981). Environment: Carbonate fecal pellets are produced in tropical marine and in non-marine environments, but are more commonly preserved in subtidal and lower intertidal zones of inner platform or ramp settings, with low water energy and reduced sedimentation rates. The fossilization of the originally soft particles requires bacterial decomposition of the organic mucus and intragranular cementation by Mg-calcite or aragonite. Lithi-
fication takes place preferentially in warm shallow waters that are supersaturated with respect to calcium carbonate, i.e. the interior part of the Bahama platform (Kornicker and Purdy 1957; Land and Moore 1980). Non-marine fecal pellets occur in salt lakes and freshwater lakes, and are known from pedogenic carbonates. Fecal pellets of filter-feeding zooplanktonic crustaceans (copepods), reported from ancient lacustrine limestones, are elongated grains with long diameters of 250 to 500 µm. Pedogenic and caliche peloids vary greatly in shape and sizes (10 µm to more than 1 mm). Ovoid grains are fecal pellets produced by worms, ants or gastropods (Klappa 1978; Jones and Squair 1988). Sometimes vegetal microstructures are preserved within the grains. Other caliche peloids originate from carbonate precipitation around roots and root hairs or in association with the decomposition of vegetal and fungal matter (types 8 and 9 in Fig. 4.11) or by micritization of skeletal grains (type 6 in Fig. 4.11; Calvet and Julia 1983).
Peloid Types
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Fig. 4.12. Three of nine peloid types (see Pl. 10 for other types!). Left: Early Cretaceous lagoonal limestone with fecal pellets, characterized by rounded, elongated micrite grains of different size. Center: Middle Devonian shelf limestones with mold peloids, representing steinkerns of ostracods whose valves have been dissolved. Right: Strongly recrystallized Carboniferous limestone exhibiting round micritic pelletoids and spar-filled intergranular pores. Scale is 500 µm.
(2) Algal peloids (Pl. 136/4) originate from the disintegration of calcified algae and calcified microbes. The term was introduced by Wolf (1965b). The peloids are the final product of the gradational abrasion of algal elements. Limestones with algal peloids are characterized by the association of structureless peloids, irregularly shaped micritic grains with inherited algal structures, and skeletal grains exhibiting preserved algal structures. Algal peloids can be derived from the mechanical destruction of upward-growing ramose, and nodular and laminar growth forms of green and red algae as well as calcimicrobes. Another process leading to the formation of algal peloids is the destruction of exterior carbonate precipitates of filamentous algae (Schroeder 1972). The peloids are commonly round to angular detrital particles. Environment: Algal pellets resulting from the comminution of algae and cyanobacteria (e.g. Girvanella tubules) are particularly common in Early Paleozoic shelf carbonates (Wolf and Conolly 1965; Coniglio and James 1985). (3) Bioerosional peloids: Rock- and biogenic hardpartboring demosponges (mainly members of the family Clionidae), occurring in tropical and temperate oceans, penetrate the substrate by means of cellular etching (Hatch 1980). They release scoop-shaped chips which contribute significantly to the fine-grained mud and peloid fraction in modern coral reefs and to carbonate coastal sediments (Fütterer 1974; Torunski 1979; Acker and Risk 1985; Edinger and Risk 1996). Endolithic sponges already appeared as early as the Cambrian and became more common during the Mesozoic (Perry and Bertling 2000). Rasping by gastropods and sea urchins results in the erosion of small limestone clasts corresponding in size and subangular shape to peloids.
Peloids resulting from reworking of mud and grains: (4) Mud peloids (Pl. 10/2, Pl. 17/1, Pl. 121/2): This category includes variously shaped micritic grains caused by the reworking of lithified carbonate mud and micrite clasts. These grains were called ‘mud grains’ (Fabricius 1966), ‘small intraclasts’ (Wilson 1967), ‘small mud aggregates’ (Purdy 1963), ‘pseudopellets’ (Fahraeus et al. 1974: reworked and current-transported lithified carbonate mud) or ‘lithic peloids’. The latter term is widely used. In fact, many mud grain peloids are genetically intraclasts formed by erosion and redeposition within the same basin. The nomenclatorial distinction between small intra- and lithoclasts and peloids is based on an arbitrary size limit (150 µm: Wilson 1967; 150 µm: Folk 1962; 500 µm: Coniglio and James 1987). A size limit of 200 µm is recommended. Criteria: Mud peloids are characterized by the similar composition of the micrite forming the peloids and the micritic matrix; irregular contact between areas with and without peloids (interpreted as bottom reworking); the preferential occurrence of peloids in lenses, laminae or layers; the hydraulic equivalence of peloids and quartz grains found together; and evidence absence of burrowing (i.e., no fecal pellets). In addition, poor sorting, a broad spectrum of peloid shapes and sizes as well as repeated occurrence within a stratigraphic profile provide evidence for interpreting the grains as lithic peloids. The last-mentioned criterion indicates fluctuations in water energy and erosion. This assists in recognizing cyclic shallowing patterns in shelf carbonates (Carss and Carozzi 1965) and allochthonous sedimentation on deeper shelves and within basins (Wilson 1969). Mud peloids form mm- to cm-thick grainstone or packstone layers within micritic limestone.
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Peloids and Cortoids
Plate 10 Peloids and Cortoids Peloids are rounded micro- and cryptocrystalline carbonate grains appearing dark in transmitted and white in direct light. Most peloids except some of fecal origin show no differentiated internal structures. The term ‘pellet’ should be reserved for peloids of fecal origin. Peloids are generally smaller than other carbonate grains. The term peloid includes genetically different grains (see Fig. 4.11) originating from biotic activities (fecal pellets, Fig. 4.12; algal peloids; bioerosional peloids), reworking of mud and grains (mud peloids or lithic peloids, –> 2; mold peloids, Fig. 4.12), early and late diagenetic alterations of grains (bahamite peloids, –> 3; pelletoids, Fig. 4.12), as well grains formed in situ by biochemical and/or chemical processes (microbial peloids, –> 1; precipitated peloids). Peloids are abundant in tropical and subtropical shallow-marine environments (platforms, ramps, reefs), but can be also found on slopes and in basins, and contribute to the formation of non-marine carbonates in pedogenic, freshwater and lacustrine settings as well. Cortoids are carbonate grains (frequently bioclasts) exhibiting thin micrite envelopes. The micritic coatings are commonly thinner then those of oncoids which may develop from the overgrowth of cortoids. In the case of mineralogical differences between the coated bioclast and the micrite rim, cortoids can preserve shape and gross structure even of diagenetically highly altered grains (–> 4; Pl. 130/3). Destructive and constructive micritization are probably the most common modes of cortoid formation. Destructive micrite envelopes are formed by the interplay of repeated microboring and infilling of the tiny little voids by microcrystalline Mg-calcite or aragonite cements. Constructive micrite envelopes are formed by an intergrowth of externally calcified filamentous algae on the surface of carbonate grains. Both modes are connected with the activity of microbes, algae and fungi. The most intensive carbonate grain micritization takes place in warm shallow tropical seas, and is carried out almost exclusively by cyanobacteria. Micritization of carbonate grains caused by microbes, however, also occurs in the deep oceans. In non-tropical settings the formation of cortoids predominantly results from chemical dissolution and recrystallization of carbonate grains in cold, carbonate-undersaturated waters. Cortoids are common in shallowmarine shelf and platform carbonates, but also originate in non-marine environments. Considering the strong impact of light on the vertical distribution of microborers, many authors have used marine cortoids as paleobathymetrical indicators. More cortoids are shown in Pl. 33/8, Pl. 35/2, Pl. 118/3, Pl. 127/1, and Pl. 130/3. 1 Microbial peloids. The clotted structure of the peloids and the occurrence as constituents of bindstones point to a microbially induced origin of the grains formed in situ within mats (‘benthic peloids’: Gerdes et al. 1994; Kazmierczak et al. 1996). Bedded lagoonal limestone. Late Permian (Zechstein): Harz, Germany. 2 Mud peloids. Small, subrounded and subangular micritic grains. Poor sorting and distinct differences in size and shape indicate that these grains originated from the reworking of weakly lithified carbonate mud. Genetically, these peloids are small intraclasts, but are conventionally assigned to the peloid group with an arbitary size limit of about 200 µm. SMF 16NON-LAMINATED. Early Carboniferous (Viséan): Nötsch, southern Austria. 3 Bahamite peloids. Peloids originating from the micritization of ooids (‘pseudopeloids’: Lakschewitz et al. 1991). The microstructure of small ooids is completely or highly altered by crystal diminution. Some ooid grains (arrows) still have preserved relics of the original structure, but most grains exhibit the characteristic criteria of peloids (structureless round micrite grains). Pseudopeloids are the dominant grains within the grainstone texture (GT) as well within the packstone texture (PT) of synsedimentary clasts. The ooids and peloids were formed on submarine pelagic swells (as indicated by pelagic faunal elements such as protoglobigerinids as nuclei of the ooids) and redeposited subsequent to transport in deeper parts of the slopes. The pseudopeloid grainflows are overlain by micritic limestones with pelagic skeletal grains. The succession reflects sedimentation on a subsiding swell accompanied by syntectonically caused downslope sediment movements. Late Jurassic (Red Nodular Limestone): Chiemgau Alps, Bavaria, Germany. 4 Cortoids. The micrite envelope of the outer part of the shell is caused by the attacks of microborers associated with micrite precipitation. The boundary between the micrite envelope and the shell exhibits the characteristic irregular outline. Note the distinct microborings on the surface (black arrows). The infestations have already destroyed the interior part of the shell fragment (white arrow). Early Carboniferous: Cracow area, Poland. 5 Cortoids, rounded fossils and redeposited oncoids. Micrite envelopes (ME) caused by microbial microboring and infilling of the holes by micrite is evident in echinoderm fragments (E). Echinoderm fragments show a characteristic cleavage structure. The micrite envelopes of this sample are stained by hematite impregnations. This points to contributions of Febacteria (see Mamet and Boulvain 1990). Note the imbricate structure represented by the parallel and inclined arrangement of the bioclasts, and indicating current transport. Dasyclad algal fragments (DA), fusulinid tests (F) and oncoids (O) are conspicuously rounded. In total, these criteria reflect the following steps: (1) Reworking and rounding of the grains formed within a subtidal environment, (2) formation of microborings and bacterially-induced staining during a quietwater phase, followed by (3) final deposition within a current-influenced ramp environment. Early Permian: Carnia, Southern Alps.
Plate 10: Peloids and Cortoids
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Reworked and redeposited mud peloids are common in protected shallow-marine environments. Larger lime mud intraclasts and smaller mud peloids can originate if lime mud dries out and small chips become comminuted and rounded, or if storms and currents disturb the bottom sediments. A never ending story – A comment on the term ‘pseudo-ooid’: The original definition of this term (Bornemann 1886) referred to transported and worn fragments of fine-grained limestone clasts observed in thin sections of the Triassic Muschelkalk. Later uses of the term connoted well-rounded, structureless carbonate grains of very different origin (Flügel and Kirchmayer 1962). But this connotation coincides with the definition of pseudopellets as well as with that of small well-rounded intra- and lithoclasts. Some authors still use pseudo-ooid to describe ooids with poorly preserved structures or as a name for strongly micritized grains. However, the term pseudo-ooid should no longer be used now. (5) Peloids resulting from the internal molds of fossils (Fig. 4.12; Pl. 132/8): Internal molds of thin-shelled ostracods, mollusks and foraminifera can be preserved as small peloids when the shells are dissolved.
Peloids resulting from syn- or postdepositional alterations of grains: (6) Bahamite peloids: Micritized grains (Pl. 10/3, Pl. 43/1, Pl. 121/4): This category comprises isolated and composite peloids whose formation is in some way comparable to that of morphologically similar grains in the modern Bahama Banks. Intensive micritization of bioclasts and of ooids, leading to the complete loss of the microstructure of the grains, can produce round micrite grains. These grains occur in association with aggregate grains, ooids and micritic intraclasts. The peloids are generally larger than algal peloids. Originally the name bahamite was proposed for ancient shallow-marine limestones with grains that closely resemble grain assemblages found in the interior of the Bahama Banks (Beales 1958), but also known from other modern shallow-marine environments (Kendall and Skipwith 1969; Logan 1974; Pusey 1975). The term refers to carbonates composed of accretionary and composite grains including aggregate grains and peloids. Gygi (1969), studying Jurassic platform carbonates, introduced the name bahamite grain for structureless microcrystalline peloids and aggregates of these peloids. Genetically these peloids are altered fecal pellets and
Peloids: In-situ Formation
micritized and recrystallized ooids or bioclasts. Reid et al. (1991, 1992) and Reid and MacIntyre (1998) emphasize the widespread occurrence of micritization/recrystallization processes affecting foraminiferal and algal grains, and leading to the formation of micro- and cryptocrystalline grains. Peloids originating from the micritization of ooids have been called pseudopeloids (Lakschewitz et al. 1991; Pl. 10/3) aiming to distinguish these peloids from peloids that represent micritized bioclasts. (7) Pelletoids (Fig. 4.12; Pl. 38/6): The term was introduced by Blatt et al. (1972) for microcrystalline grains that have been generated by the recrystallization of various carbonate particles, such as skeletal grains (e.g. foraminifera) and ooids. Pelletoids are often amalgamated and have diffuse outlines due to compaction and aggrading crystal growth. Some pelletoids may show vague residual internal structures of the original grain. Distinguishing between pelletoids and bahamite peloids can be difficult. –> Note that some authors use the term ‘pelletoid’ for textures exhibiting pellet-like grains which are believed to have been formed by incipient or incomplete recrystallization of micrite. Other authors use the adjective ‘pelletoid’ to describe limestones with (non-altered) peloids.
In-situ formation: (8) Microbial peloids (Pl. 8/6, Pl. 10/1): There are increasingly more reports about the in-situ formation of carbonate particles in modern microbial mats, on ancient micrite crusts, and within the peloidal textures of mud mounds (Box 4.7; Friedman 1978; Javor and Castenholz 1981; Pickard 1992; see Facies, volume 29, 1993). These particles correspond morphologically to peloids, ooids, oncoids and aggregate grains. In-situ growth of modern oncoids has been reported from lakes, lake margins and sabkhas (Jones and Wilkinson 1978), and rivers or streams (Freytet and Plaziat 1965; Hartkopf-Fröder et al. 1989). Ancient in situ growth has been claimed for Tertiary lacustrine oncoids (Leinfelder and Hartkopf-Fröder 1990). Algal and cyanobacterial pellets formed in situ (benthic peloids) occur in microbial mats of salinas, saline lakes and intertidal settings (Gerdes et al. 1994; Kazmierczak et al. 1996). The rounded micritic grains as well as autochthonous micrite (Sect. 4.1.2) are the result of biochemical precipitation triggered by microbial activity leading to the decomposition of organic
Fabrics of Peloidal Limestones
tissue (e.g. sponges) and organic substance. Smallscaled peloids occurring in microcavities off Lizard Island, Great Barrier Reef, grow in organic mucilages enriched by Ca-binding proteins (Reitner and Neuweiler 1995). Tuberoids are considered a special kind of peloid. The term was suggested for rounded and nodular, darkcolored and variously sized micritic particles (Fritz 1958) occurring in abundance in bedded limestones adjacent to sponge reefs. These grains are common in Jurassic and Cretaceous limestones with remains of siliceous sponges and micritic crusts (Pl. 116/1; Kott 1989). The grains are interpreted as precipitates of carbonate induced by decaying organic matter. (9) Precipitated peloids (Pl. 8/5): Inorganic and organically induced chemical in-situ precipitation has been proved for the origin of very small (20–60 µm, usually 40 µm) peloids within calcite cements and occurring within inter- and intraskeletal cavities of reefs as well as in non-reef cavities (Macintyre 1985; Chafetz 1991; Reitner and Neuweiler 1995). These peloids are a common characteristic of submarine Mg-calcite cements (Macintyre et al. 1968; Shinn 1969; Land and Goreau 1970; Alexandersson 1972). The peloids consist of well-developed dentate rims of microcrystalline rhombic calcite forming small rosettes (5–10 µm in length) around a nondescript cloudy anhedral nucleus. This bimodal texture has been explained as the result of (a) initial rapid precipitation of anhedral magnesium calcite, followed by a second stage of slower precipitation of larger euhedral Mg-calcite, forming the dentate rims (Macintyre 1977, 1978; Lighty 1985); (b) intermediate stages of the development of advancing crystal growth from microcrystalline Mg-calcite crystals (Alexandersson 1972); (c) crystal growth around a nucleus of suspended silt-sized clastic particles that were flushed into cavities by wave pumping (James et al. 1976); and (d) growth of cement crystals induced by bacteria (Buczynski and Chafetz 1991; Sect. 4.1.2). Peloidal carbonate cements occur as peloidal fill (e.g. in small reef cavities), in the form of rim cement and coatings on grains, and as thin dense crusts on the surface of coral skeletons and reef rubble. Common peloid microfabrics Carbonate peloids contribute to the formation of four widely distributed microfabrics (Box 4.7) important in the context of facies interpretations.
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Diagenesis of peloids Modern marine peloids consist of High-Mg calcite and/or aragonite, freshwater peloids are composed of Low-Mg calcite (Brand and Veizer 1983; Rao et al. 1994). Peloid limestones appear to have suffered remarkably little compaction (Beales 1965). The preservation of carbonate peloids of fecal origin is only possible if lithification takes place rapidly. The precipitation of Mg-calcite peloid cement plays an important role in the submarine cementation of upper slope sediments. Gradations from amorphous microcrystalline calcite to clotted peloidal micrite and peloidal micrite, followed by the in-situ precipitation of polypeloidal aggregates have been observed on Bahamian slopes (Wilber and Neumann 1993).Selective dolomitization of peloids is common in fecal pellets as well as algal peloids and mud grain peloids. Significance of peloids Ancient peloid limestones account for thousands of cubic kilometers of lagoonal, shelf, bank, reef-core and mud mound deposits, but they are relatively rarer than other limestone types. Paleoenvironmental proxies: Many shallow-marine peloidal grainstones have a scarcity of biotic constituents or contain only specific biota (ostracods, benthic foraminifera, green algae) indicating conditions were inadequate for many animals with regard to salinity and substrate. By contrast, peloidal packstones and wackestones with abundant fecal pellets and burrows indicate adequate nutrients and oxygen. Depositional proxies: Many ancient peloid limestones appear to have formed in shallow warm-water environments (banks, shelf- and reef lagoons). Finegrained peloidal limestones are generally regarded as typical of shallow, low-energy, restricted marine environments. This generalization is supported by many case studies, but the examples should be critically evaluated for the following reasons: 1 Accumulation of peloids may record specific preservation sites rather than specific depositional environments. 2 Mass occurrences of fecal pellets can be allochthonous or autochthonous, and the size and sorting of these peloids, therefore, may not be a measure of current energy. 3 Peloids may be deposited near the place of their origin but can also be transported to shallower or deeper environments. Understanding of the hydrodynamic behavior of carbonate fecal pellets is important for interpretating of the origin of peloidal carbonates (Wanless et al. 1981).
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4 Peloid pack- and grainstones are not restricted to shallow-marine environments but are also important constituents of slope and basin sequences formed by turbidites and debris flows. 4.2.3 Cortoids – Carbonate Grains Characterized by Micrite Envelope A specific category of carbonate grains are bioclasts, ooids or lithoclasts and peloids, whose periphery ex-
Cortoids
hibits a circumgranular non-laminated micritic rim commonly called a ‘micrite envelope’. Because of the significance of these grains as facies indicators, a separate grain type category called ‘cortoid’, was proposed (from ‘cortex’, Latin for bark; Flügel 1982). Micritization and micrite envelope Micritization is a process whereby the margins of carbonate grains or the total volume of grains are replaced by crypto- or microcrystalline carbonate crys-
Box 4.7. Common fabrics in peloidal limestones usually interpreted as products of microbially or biologically induced carbonate production. Micritic clotted fabric (Pl. 10/1) Texture: Densely spaced, variably sized globular and irregular peloids forming amalgamated clots; commonly within a microspar, micrite or spar matrix. Origin: (a) Recrystallization of carbonate mud and peloidal micrite (Cayeux 1935; Bathurst 1970; Leeder 1982), (b) diagenetic alteration of soft-pellet grainstones to wackestones (Bathurst 1970) (c) diagenetic amalgamation of precipitated peloids (Reid 1987), (c) diagenetic alteration of soft-bodied organisms (Bourque 1984), (d) diagenetically modified algal debris (Coniglio and James 1985), (e) products of in situ calcified mats of benthic coccoid cyanobacteria (Kazmierczak et al. 1996), and (f) grazing and decay of algal mats (Pratt 1982). Occurrence: Common in tidal laminated carbonates, in mud mounds, and reefs (Sun and Wright 1983; Neuweiler 1993) as exemplified by Triassic reefs, where about 75% of the reef core framework can consist of precipitated peloids occurring in cavities, intraskeletal voids and organic crusts (Reid 1987). Sparry peloidal fabric (Pl. 8/5) Texture: Abundant, very small peloids formed in situ and consisting of a microcrystalline core surrounded by a dentate crystal rim (type 9: Fig. 4/11). The peloids are part of submarine carbonate cements. Origin: Chemical or biochemical precipitation (type 9, Fig. 4/11). Occurrence: Common in intra- and interskeletal voids of reef rocks similar to clotted fabrics, associated with internal sediments and carbonate cements and within crusts associated with corals and sponges. Thrombolititic peloidal fabrics (Pl. 50/5) Texture: Non-laminated structure. Abundant peloids of irregular shape and size, some retaining filamentous structures. Common calcite ‘spherulites’ (peloids) consisting of radially-oriented, non-ferroan calcite, typically 12–20 µm in diameter with a cloudy micrite core, 1–10 µm in diameter. Origin: (a) In-situ precipitation. The formation of in-situ spherulites within the sediment is controlled by supersaturation of CaCO3 in solution and a site for nucleation. Supersaturation of CaCO3 is caused by (a) the release of CO2 and NH4 on the decay of organic matter and (b) bacterial sulphate reduction. Bacteria can form spherulites in aerobic, anaerobic, agitated and non-agitated laboratory conditions, but experimental precipitation of spherulites without the presence of bacteria points to the possibility that spherulitic textures also can form as a by-product of decayed organic matter. Spherulites in stromatolites have been interpreted as calcified cyanobacteria. (b) Allochthonous grains. Deposition of small allochthonous fecal pellets and algal peloids within cavities (James et al. 1976) or formation of pellets by cryptic filter feeders (e.g. clionids: Land and Moore 1980). These interpretations are supported by the fact that peloids can settle at the base of cavities. Occurrence: Thrombolitic fabrics occur in lacustrine limestones (Pl. 131/5) and beachrocks, and are common in laminated tidal and subtidal carbonates (stromatolites) as well as in subtidal shelf carbonates, mud mounds and non-skeletal microbialite reef frameworks (particularly in the Carboniferous and Permian, Jurassic and Cretaceous). Biocementstones (Tsien 1981) - Sect. 8.2 Texture: Fine-grained micritic limestones consisting of peloids, delicate small and/or poorly calcified organisms contained within abundant, localized carbonate cement. Origin: Fabric, formed by free-living microorganisms (cyanobacteria, algae, sulfur-reducing and sulfate-reducing bacteria) and small encrusting benthic invertebrates, which trap and build carbonate mud from bottom waters, resulting in the formation of small micritic ‘protopeloids’ (Tsien 1985). Amalgamation of the peloids leads to the formation of peloidal automicrite. A large part of the limestones is formed by biologically induced carbonate cements. Occurrence: Reef frameworks, specifically in the Permian and Triassic (Webb 1996).
Origin of Cortoids
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tals. Incomplete micritization leads to the formation of cortoids, whereas complete micritization can bring about a gradual to total alteration of the original grain (Pl. 134/7) and the formation of bahamite peloids (Fig. 4.11). The term micritization was worked out in the context of studies of modern lagoonal carbonates (Bathurst 1966). Later, the term was expanded to include all processes resulting in obliteration of original carbonate microstructures through gradual alteration to cryptocrystalline textures. Micritization processes are controlled by biological and chemical factors and take place in shallow- and deep-marine as well as in terrestrial and lacustrine environments. Micritization of aragonitic skeletal grains has been described from the Bahamas and the Persian Gulf, micritization of Mg-calcite skeletal grains from Florida and Belize (Reid et al. 1992). Micritization processes are fundamental in loosening the surface, abrading and rounding carbonate grains, completely destroying the original structures of skeletal and other grains. Persistent micritization results in the formation of carbonate muds (Sect. 4.1.1). The term micrite envelope was first used by Bathurst (1964) in the study of syngenetic and early diagenetic changes affecting modern skeletal grains. The term refers to a thin, non-laminated coating of very fine micrite around carbonate grains, particularly skeletal grains or ooids. Modern envelopes consist of aragonite or High-Mg calcite. In limestones, the calcitic envelopes of cortoids appear white in reflected light and dark in transmitted light. The thickness of the envelopes varies between a few µm and about 500 µm. Origin of cortoids Micrite envelopes originate from destructive and constructive processes that take place at or near the sediment-water surface. Both photosynthetic and non-photosynthetic organisms play a major role in the formation of cortoids. Common modes are: (1) Destructive micritization related to microboring organisms: The classical model is a three-step process (Bathurst 1966; Fig. 4.13: A1 and A2). (a) Microbolic products of microendoliths lead to biochemical dissolution of skeletons (Ehrlich 1999) leaving microborings (Golubic et al. 2000). Endolithic microborers attacking the surface of bioclasts and ooids, particularly in shallow-marine environments, produce tiny little voids (size < 1 µm to about 50 µm) which are colonized by filamentous cyanobacteria, green and red algae as well as fungi (Sect. 9.1). Boring organisms can colonize a grain within
Fig. 4.13. Origin of cortoids. A – Three-step destructive micritization related to microborings (adapted from Bathurst 1966). A1: Endolithic cyanobacteria, algae or fungi produce (1) tube-like microborings (black) on the surface of bioclasts. (2) Vacant tubes (white) originating after the death of the endoliths are (3) infilled with microcrystalline carbonate cement (stippled in A2). Multiple boring and filling result in the formation of micrite envelopes. B – Destructive micritization by concurrent cyanobacterial microboring and filling of boreholes (modified after Reid and Macintyre 2000). Microboring (black) just beneath the grain surface produces tunnels which are filled by radial-fibrous cement as the organism advances. C – Constructive micritization caused by precipitation of microcrystalline calcite around exposed filaments of endo- and epilithic algae and cyanobacteria (adapted from Kobluk and Risk 1977a). C1: Initial microborings; C2: Colonization of the grain surface by endo-epilithic filamentous algae; C3 and C4: Precipitation of calcite cement on the surface of filaments; C5: Resulting micrite envelope. Most micritic rims are thinner than 500 µm.
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a matter of days and infest a grain in a few weeks; (b) Death of the microborers and vacation of tubes; (c) Emplacement of micritic aragonite or High-Mg calcite cements within vacant tubes. The formation of cements can be triggered by organic acids produced during the bacterial decomposition of organic compounds. Multiple repetition of boring and filling destroys the peripheral zone of the grains and finally results in the formation of a circumgranular micritic rim recognized in thin section. The surface of the rim exhibits shallow pits; the boundary between the envelope and the skeletal core is irregular and may preserve microtube structures penetrating into the skeletal grain (Pl. 10/4). Destructive cortoids are abundant in shallow-marine environments (Lloyd 1971; Margolis and Rex 1971; Harris et al. 1979), but also occur in transitional and non-marine (e.g. caliche: James 1972; lacustrine environments: Schneider et al. 1983; Pl. 130/3) and vadose settings (Pl. 33/4; Pl. 35/2). A special mode of destructive micritization has been proposed by Reid and Macintyre (2000). This model focuses on concurrent microboring and filling of bore holes (Fig. 4.13B). In contrast to the Bathurst model, micritization occurs in microborings of cyanobacteria penetrating just beneath the grain surface and producing tunnels. The tunnels are filled by radial-fibrous aragonite cements that are precipitated as the organism advances. Extensive multiple repetition of the process results in the micritization of just the peripheries of the grains or the whole grain. As opposed to micritized by processes corresponding to the Bathurst model, the grain margins here are almost completely preserved. (2) Constructive micrite envelopes related to epilithic organisms (Kobluk and Risk 1977a, 1977b): These cortoids result from the precipitation of microcrystalline calcite around and between dense populations of exposed filaments of endo- and epilithic algae and cyanobacteria (Fig. 4.13C). The precipitation occurs predominantly upon dead filaments. The outer boundary of the micrite rim appears irregular. Contrary to mode 1, the micrite envelope is formed without destruction or alteration of the grain periphery. The process involves addition of carbonate to the exterior of the grain and requires grain stabilization for the duration of the envelope formation. This occurs in low-energy environments or shallow burial. Another possibility for the formation of constructive micrite envelopes is the growth of nannocrystals within biofilms coating carbonate grains (Loreau 1970) or the dissolution and reprecipitation within a muci-
Significance of Cortoids
laginous coating of the grains (Kendall and Skipwith 1969). Modified versions of these interpretations are gaining more and more acceptance. Constructive micrite envelopes are formed both in marine and in continental and lacustrine environments (e.g. karst: Jones and Kahle 1995; eolianites: Calvet 1982; lake carbonates: Schneider et al. 1983). Terrestrial vadose cortoids are caused by calcified filaments of fungi and actinomycetes that collapse and coalesce forming an intertwined organic meshwork that in turn triggers the precipitation of microcrystalline cements. Recognizing different modes of origin: Destructive and constructive micrite envelopes are difficult to distinguish in thin sections. Constructive cortoids may show irregularities on the outer surface of the grains and the micritic rim may be of variable thickness, whereas destructive cortoids exhibit irregularities towards the grain interior and often reveal relicts of microborings. Because primary mineralogy controls the susceptibility to the effects of microborings in carbonate grains (Perry 1998), aragonitic bioclasts often are more frequently bored than primary calcitic bioclasts. Micritization is selective, affecting some shells more than others. Shells with homogeneous prismatic microstructures exhibit micritic envelopes more often than shells with fibrous or foliated microstructures. (3) Partial dissolution and recrystallization: Partial dissolution of skeletal grains associated with recrystallization and crystal diminution are other processes known to lead to the formation of micrite envelopes. Mollusk shells composed both of aragonite and Mg-calcite are affected differently by submarine early diagenetic selective dissolution. This may result in a peripheral rim consisting of residual micrite (Winland 1968; Alexandersson 1972; Pl. 52/6). The micritic preservation of echinoderms in the chalk of Texas is an outcome of partial dissolution called ‘inorganic micritization’ (Neugebauer and Ruhrmann 1978). High-Mg calcite echinoderm fragments may lose Mg and convert to Low-Mg calcite without a dissolution stage. Inorganic micritization and ‘shell-residue micrite’ envelopes are widespread at the bottom of modern seas in temperate regions. Micritization also can be the result of decomposition of indigenous organic matter within shells, producing the simultaneous growth of microcrystalline carbonate (Purdy 1968; Burgess 1979).
Significance of cortoids Preservation potential: The susceptibility of skeletal grains to micritization differs with regard to the
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Oncoids and Rhodoids
fossil groups affected. This is demonstrated, for example, by fossils in Silurian shelf carbonates showing decreasing susceptibility to micritization from trilobites to ostracods, brachiopods, bryozoans, corals, and tentaculites (Mohamad and Tucker 1992). Micrite envelopes are relatively resistant to dissolution, protect the grains from being destroyed, and often contribute to the preservation of casts of primary aragonitic skeletal grains during freshwater diagenesis (Pl. 33/4). The stability of a micrite envelope depends on its stable mineralogical composition and on a high content of organic matter within the micritic rim. The envelopes are prone to selective dolomitization; this may imply that the mineralogy of the envelopes was once High-Mg calcite (Buchbinder and Friedman 1970). Paleoenvironmental proxies: In accordance with the original explanation that micrite envelopes originate from light-dependent boring organisms (Bathurst 1966), coated grains are often regarded as indicators of ‘shallow-marine warm-water’ environments (e.g. Leonard et al. 1981). Having observed differences in the abundance of microboring frequency and intensity at different water depths (Vogel et al. 1995), some authors use cortoids and micritized grains as indicators of specific water depths, thus considering the abundance of bored grains to be reliable indicators of shallow waters down to a few tens of meters (Swinchatt 1969; BernardDumanois and Delance 1983). Although there are many modern and ancient examples supporting this interpretation, caution is necessary because a few records of micrite envelopes being formed via non-light dependent microborings and infillings of calcite cements are also known from deepsea settings (Friedman et al. 1971; Hook et al. 1984). In addition, strong boring and leaching is also a common feature of colder carbonate-undersaturated marine waters (Zeff and Perkins 1979; Campbell 1983). But because carbonate precipitation within the boreholes is rare or absent, typical cortoids are rarely preserved in these settings. Depositional setting: Bedded grainstones with a high percentage of cortoids and other coated grains, together with rounded bioclasts characteristically form in areas of constant water action, at or above wave base. These conditions occur in the ‘winnowed platform edge sands’ of carbonate platforms (standard microfacies type 11, see Sect. 14.3.1) as well as in current-washed sand shoals of inner ramp settings (Pl. 18/1, Pl. 33/ 8). Grainy cortoid sediments are often associated with patch reefs scattered across the shelves or occurring in a leeward position behind reef belts (Pl. 118/3, Pl. 127/1).
4.2.4 Oncoids and Rhodoids Nodular coated grains formed by microbes, algae and other encrusting organisms are common constituents of platform, reef and slope carbonates. These grains represent two major categories: 1) oncoids (formed predominantly by calcibionts and algae) and 2) rhodoids (formed by free-living calcareous red algae). Oncoids and rhodoids are excellent indicators of paleoenvironmental and climatic conditions, and sea-level fluctuations and depositional settings. As opposed to other carbonate grain types, the prevailing settings of oncoid formation changed over the Phanerozoic and rhodoids started to dominate from Late Mesozoic to Cenozoic. Oncoids (onchos, Greek for nodule) are unattached, rounded, mm- to cm-sized, calcareous or non-calcareous nodules that commonly exhibit a micritic cortex consisting of more or less concentric and partially overlapping laminae around a bio- or lithoclastic nucleus. The term oncoid is not a genetic term and should be regarded as purely descriptive. The name is commonly used in Europe, but not in America where the term ‘pisoid’ (or pisolith) was and is still used both for coated grains, called oncoids here, but also for grains similar in internal structure but larger than ooids (Sect. 4.2.6). Other not very precise names applied to oncoids are ‘algal balls’ and ‘algally-coated grains’, ‘mummies’ or ‘osagid grains’. Oncoids and rhodoids are common in limestones and marly sediments. The latter allow individual grains to be investigated with regard to their composition, size and shape. and grain assemblages. The combined study of thin sections and acid residues augments information on the organisms involved in the formation of oncoids. SEM observations and geochemical studies are necessary to evaluate the contribution of clay material and organic matter to carbonate and non-carbonate oncoids (e.g. Carbone and Civitelli 1974; Andrews 1986; Scudeler Bacelle and Marusso 1983; Palmer and Wilson 1990). Comparison with other grain types: Oncoids differ from cortoids by the greater thickness of the micritic coatings and the frequency of preserved cyanobacterial, algal or foraminiferal structures. Ooids and pisoids are commonly smaller und have regular concentric laminae. As a rule but with a few exceptions, many oncoids and rhodoids are considerably greater than all other carbonate grains. Terminology Various names have been proposed for organically coated rounded grains consisting of a core and a cortex formed by encrusting by organisms.
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Oncoids and Macroids
Spongiostromate vs. porostromate oncoids: The name ‘oncoid’ was suggested by Heim (1916) for micritic grains found in Jurassic shelf limestones in Switzerland. The author discussed a bacterial origin. Apparently unaware of Heim’s definition, Pia (1927) considered oncoids to be unattached stromatolites. Classifying blue-green algae (cyanobacteria) of unknown affinities he distinguished two informal groups, Spongiostromata and Porostromata. Spongiostromata are nodules or heads with defined growth forms but without (or very rarely with) visible organic microstructures. Concentric laminae originate from fine grains of sediment adhering to the mucilagineous surface of cyanobacteria and algae and also from the precipitation of calcium carbonate in response to the withdrawal of carbon dioxide during photosynthesis. Porostromata comprise forms exhibiting preserved tubular microstructures. To subdivide the Spongiostromata, Pia used the terms ‘Stromatolithi’ and ‘Oncolithi’. The former include forms which grow attached to the substrate; the latter term refers to unattached free nodules. Today, the descriptive term spongiostromate oncoid (Pl. 11/3) is used for micrite oncoids possessing a laminated dense micritic or spongy fabric without visible filaments; porostromate oncoid (Pl. 12/3) applies to grains exhibiting fine micritic-walled tubes, generally less than 100 µm in size (Monty 1981). Since these tubes are prone to obliteration by recrystallization, the apparent lack of tubular structures is not an unequivocal sign of spongiostromate oncoids. The co-existence of oncoids with well-preserved porostromate microstructures and laminated oncoids of the same size and shape, but without tubular microstructures may point to diagenetically altered porostromate oncoids (‘paraporostromate oncoids’: Catalov 1983).
Skeletal and non-skeletal - a basic subdivision: Because the Porostromata include fossils of different systematic position, Riding (1977) introduced the term skeletal stromatolites for “stromatolites in which the organisms responsible for their formation are commonly preserved as calcified fossils”. The term non-skeletal is identical with ‘spongiostromate oncoid’. Skeletal stromatolites and porostromate oncoids are typically constructed by calcified cyanobacteria, but green algae and microproblematica also take part in the formation of non-attached or attached structures. Recent counterparts of Girvanella, a common microfossil in porostromate oncoids, consist of Low-Mg calcite. The same mineralogy is inferred for ancient Girvanella oncoids, but ferroan calcite in some Girvanella skeletons indicates a High-Mg calcite mineralogy, too (Richter and Füchtbauer 1978).
Oncoids and stromatolites: A particular type of modern and ancient oncoids is associated with intertidal and shallow subtidal algal mats and stromatolites. These oncoids are microbialites and sometimes formed in situ within algal mats (Dahanayake et al. 1985). A geometrical classification of modern shallow-marine stromatolites and oncoids (including ‘algal biscuits’: Gebelein 1969) was proposed by Logan et al. (1964) and promptly adapted to ancient oncoids (e.g. Radwanski and Szulczewski 1966). However, the premise that recent marginal-marine oncoids are useful paleoenvironmental analogues of ancient lithified calcareous oncoids was questioned by Monty (1972), who stated that modern freshwater algal nodules are the much better analogues for pre-Cenozoic marine oncoids.
Non-algal oncoids and macroids: The original definition of the term oncoid has been broadly expanded. It presently covers carbonate nodules formed as biogenic accretions not only by microbes, algae and micro-organisms, but also by bryozoans or serpulids for example. Bryozoan nodules, also called ‘bryooncoids’ or ‘bryooids’, are built by free-living bryozoan colonies associated with microbes (Hillmer et al. 1996, Scholz 2000). The laminar growth of bryozoan colonies circumcrust litho- or bioclastic nuclei. Modern cm-sized bryooncoids occur in temperate regions (Rider and Enrico 1979). Ancient counterparts have been described from Neogene and Cretaceous sediments (Balson and Taylor 1982).
Rhodoids and cyanoids - more sensitive categories: Since it is desirable to separate coated nodular grains of different origin and formed in different environments, specific names have been coined, starting with the term ‘rhodolite’ for “nodules and detached branch growth with a nodular form composed principally of coralline algae” (Bosellini and Ginsburg 1971). Because rhodolite has priority for a variety of the mineral garnet, the name was later replaced by ‘rhodolith’. Peryt (1983) suggested the name rhodoid for nodular grains formed mainly by coralline but also by other red algae. Finally, Riding (1979, 1983) proposed the name cyanolith or cyanoid for porostromate oncoids formed predominantly by calcified filamentous and coccoid cyanobacteria that are commonly preserved as tubes. Corresponding stromatolites were called ‘skeletal stromatolites’. Oncoids with tiny tubular microfossils that can not be identified as being of cyanobacterial origin should simply be called ‘porostromate oncoids’.
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Classification of Oncoids, Rhodoids and Macroids
Compositional Types
Examples
Dominating Biota
Micrite oncoid
Pl. 12/2
Microbes
Rhodoids and Macroids
Oncoids
Spongiostromate oncoids Pl. 11/3-4; Pl. 12/7; (non-skeletal) Pl. 89/3; Pl. 131/5
Microbes and algae
Note: Micrite oncoids and spongiostromate oncoids also originate from micritization of porostromate oncoids
Porostromate oncoids (skeletal)
Pl. 12/1,3,4; Pl. 142/11; Pl. 136/4
Calcified cyanobacteria, e.g. Cyanoids (formed by calcialgae microbes); Girvanella oncoids
Foraminiferal oncoids
Pl. 11/5-6; Pl. 51/2 Pl. 111/18
Encrusting foraminifera
e.g. Nubecularia oncoids
Composite oncoids
Pl. 51/9; Pl. 119/3
Association of cyanobacteria, algae, foraminifera or other encrusting organisms
e.g. Girvanella-foraminiferal oncoids; Girvanella-Spirorbis oncoids; foraminiferal-algal oncoids
Rhodoids
Pl. 56/8
Calcareous red algae
e.g. Coralline algal rhodoids; Archaeolithophyllum rhodoids
Composite rhodoids
Pl. 12/6; Pl. 149/6
Associations of red algae and other encrusting organisms
e.g. Foraminifera-red algal rhodoids
Macroids
--
Foraminifera or bryozoans, or other encrusting organisms
e.g. Acervulinid foraminiferal macroids; bryozoan macroids (bryoids)
Fig. 4.14. Practical classification of oncoids, rhodoids and macroids.
The name macroid was suggested by Peryt (1983) for organically coated grains with sizes >10 mm (see Fig. 4.14). Presently the term is used for cm-sized coated grains formed predominantly by encrusting metazoans or protozoans, sometimes associated with algae (e.g. balanid macroids, serpulid macroids). Foraminiferal macroids constructed by Acervulina or Gypsina are common constituents of Cenozoic warm- and coolwater carbonates (Perrin 1992). Nodular growth forms produced by free-living corals were called coralliths (Gill and Coates 1977). Composite growth types: Many oncoids and rhodoids consist of two or more encrusting organisms that occur in abundance and form repetitive or randomly distributed structures. Examples are oncoids formed by calcimicrobes and encrusting foraminifera, or oncoids built by foraminifera and coralline algae (Pl. 149/6). The latter have been called ‘for-algaliths’ (Prager and Ginsburg 1989), but a simple descriptive designation as foraminiferal-coralline algal oncoids of these composite oncoids is sufficient, too (Fig. 4.14). Highly diverse oncoid communities indicate rather stable conditions during the growth of the grains over time. Size catagories: Kutek and Radwanski (1965) and Peryt (1983) proposed names describing size grades of
oncoids: micro-oncoid or microids (<2 mm), piso-oncoid (2 to 10 mm), and macro-oncoid or macroids (>10 or >20 mm). Most modern and Tertiary macroids which are not formed predominantly by coralline algae (e.g. foraminiferal macroids) exhibit sizes ranging from about 10 to 100 mm. Since the abundance of very small and of exceptionally large oncoids may indicate specific hydrodynamic controls, the differentiation of micro-oncoids and macro-oncoids is meaningful. The term piso-oncoid should be replaced by meso-oncoid for oncoids with sizes from about 2 mm to about 10 mm, but usually only up to 5 mm, because the suffix piso- is too close to ‘pisoid’ which refers to a grain category distinctly different from oncoids. Semantic problems: Oncoid designates the individual grain, oncolite a limestone consisting of more than 50 % oncoids. This definition parallels the strongly recommended use of ooid for the grain and oolite for a sedimentary rock composed of ooids (Teichert 1958). Analogously, the terms cyanoid and rhodoid should be used for the the individual grains, and cyanolite and rhodolite for the carbonate rocks. But many authors still prefer the suffix ‘-lith’ and not ‘-oid’ in order to name the individual grains (rhodoliths: Bosence 1983). The terms ‘oncoid-bearing’ or ‘rhodoid-bearing’ characterize limestones with less than 50% oncoids.
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Practical classification Fig. 4.14 outlines a practicable classification of common oncoid and rhodoid types, based on the dominating biota involved in the formation of the grains: (A) Oncoids: The main types are differentiated by (a) the existence/absence of non-skeletal or skeletal microstructures, the biotic composition and the occurrence of non-algal/cyanobacterial (zoogenic) biota, and (b) the diversity of co-existing biota that contribute to the cortices. Micrite oncoids consist of micrite, exhibit no distinct lamination, but are comparable in shape to other oncoid types. These oncoids result from trapping of fine-grained sediment on the surface of vanished microbial/algal felts and/or microbially induced precipitation of carbonate. The oncoids correspond genetically to spongiostromate oncoids. Many micrite oncoids, however, are caused by obliteration of porostromate and other oncoids due to pervasive microboring and recrystallization. Spongiostromate oncoids are poorly or distinctly laminated grains, consisting of micritic layers around a bio- or lithoclastic nucleus. The laminae are the result of sediment trapping and the calcification of microbial or algal filaments. Owing to the rapid decomposition of organic filaments, commonly only the lamination but not the filaments are visible. Spongiostromate microfabrics are variable. They embrace micritic, spongy, fenestral, peloidal and other fabrics. These fabrics are produced by differences in environmental factors, biotic composition and diagenetic overprint. Comparable textures are known from recent as well as ancient stromatolites (Logan et al. 1974; Monty 1976, 1976). Organisms which can be incorporated within spongiostromate oncoids are foraminifera, bryozoans and various microproblematica (Kutek and Radwanski 1965; Elliott 1966; Dahanayake et al. 1976). Porostromate oncoids are layered or non-layered, usually micritic grains exhibiting remains of skeletal microbes and calcareous algae. Oncoids formed by recognizable cyanobacteria are called cyanoids. Porostromate microfabrics are defined by loose or tangled, variously oriented, straight or sinuous calcified filaments. Typical cyanoids are Girvanella oncoids, known from the Cambrian to the Jurassic. A characteristic feature of these oncoids is the alternation of dark Girvanellabearing and light barren laminae. Girvanella oncoids are more rapidly lithified than spongiostromate oncoids as evidenced by more abundant macroborings in porostromate oncoids. Foraminiferal oncoids are formed by encrustations of calcareous or agglutinated tests, starting with settlement on a hard surface, e.g. a bioclast. Some foraminiferal oncoids are distinctly layered,
Modern Oncoids
others are not. Composite oncoids consist of several taxa that form alternating layers. Common types are (1) grains built by various cyanobacterial and algal taxa; (2) cyanobacteria and/or algae associated with encrusting foraminifera (e.g. ‘osagid grains’: Henbest 1963; Toomey et al. 1989); and (3) cyanobacteria and algae, associated with taxonomically different encrusters (e.g. foraminifera, bryozoans, serpulids (Girvanella-spirorbid oncoids: Poncet and Lapparent 1975). Rare groups which contribute to the formation of oncoids, particularly in non-marine and marginal-marine environments, are spirorbid worms and vermetid gastropods (Poncet and de Lapparent 1975; Burchette and Riding 1977). (B) Rhodoids: They are easily recognized by the very fine micritic network of the microstructures (Pl. 54, Pl. 56). Opposed to oncoids, the internal structure of rhodoids is strongly controlled by the internal growth pattern of the red algal thalli. Rhodoids, therefore, are not necessarily laminated. Nodular thalli of very finegrained calcareous red algae may be confused with micrite oncoids, e.g. Marinella Pfender, a common constituent of Late Jurassic and Cretaceous platform carbonates (Pl.136/3). Composite rhodoids originate from the growth of red algae associated with encrusting foraminifera and other sessile organisms. The diversity within the cortex may be low, moderate and high, and is a sensible measure for short-term changes of environmental conditions. In-depth classifications: Oncoids have been subdivided according to biotic composition, internal structures (lamination), shape and growth form as well as size (Maslov 1960; Dragastan 1964; Zhuravleva 1964; Logan et al. 1964; Radwanski and Szulczewski 1966; Catalov 1970; Dahanayake 1977, 1978). Similar criteria have been used for categorizing rhodoids, but the types of internal structures, systematic composition of the red algae as well as the shapes of the oncoids are more relevant (Bosellini and Ginsburg 1971; Bosence 1983). Catalov (1983) differentiated ‘phytogenic’ oncoids (corresponding to porostromate oncoids), ‘zoogenic’ oncoids (foraminiferal oncoids), and ‘zoophytogenic oncoids’ (composite oncoids).
4.2.4.1 Oncoids Modern oncoids Present-day oncoids form in a wide variety of environments, ranging from freshwater lakes, streams and marshes to marine inter- and subtidal conditions. It is
Recent Oncoids
generally believed that many oncoids and concentric laminae are formed by adhesion of fine grains of sediment to the mucilaginous surface of algal mats (Pl. 11/ 1, 2) and also by precipitation of calcium carbonate in response to the withdrawal of carbon dioxide by algae during photosynthesis. Nodule turning is facilitated by waves and tidal currents, a sloping surface, or bioturbation of the underlying surface by grazing herbivorous fish. The contribution of cyanobacteria and other bacteria to the formation of cyanoids and spongiostromate oncoids has generated a large number of publications in the last few decades that have clearly established the fundamental role played by cyanobacteria (Pentecost and Riding 1986; Gerdes et al. 1994). These observations were reinforced by experiments that succeeded in creating calcite and aragonite biocrystals under laboratory conditions approximating natural conditions (Chafetz and Buczynski 1992; Knorre and Krumbein 1992; Castanier et al. 1999). The eminent function of microbes for the formation of freshwater oncoids and ooids was exemplified by Davaud and Girardclos (2001) in Lake Geneva, Switzerland: In surficial sediments nucleation of Low-Mg calcite occurs in close association with organic films formed by the mucus of filamentous cyanobacteria and diatoms. In subsurface sediments the intercrystalline porosity is reduced and the micritic porous and organic-rich coatings are transformed into a dense mosaic with scattered molds of filaments, creating a microfabric similar to ancient oncoids. Recent non-marine oncoids: Free-lying, rounded irregularly laminated oncoids occur in fluviatile and lacustrine habitats; they form in association with calcareous tufa, and as microbial coatings on tropical karst cavities (‘terrestrial oncoids’: Jones 1991). Lacustrine and fluviatile oncoids are similar with regard to their morphology and internal criteria. Oncoids from freshwater lakes have been studied in Europe and North America (e.g. Pia 1933; Jones and Wilkinson 1978; Schäfer and Stapf 1978; Murphy and Wilkinson 1980; Schneider et al. 1983; Obenlüneschloss 1991). The ellipsoidal or subspherical, centimeter to tens of centimeter-sized oncoids are produced by cyanobacteria in association with green and red algae as well as diatoms. The oncoids typically possess a laminated, porous microstructure, with layers or tufts of micrite tubes representing the encrusted sheets of cyanobacteria or green algal filaments. The taxonomically differentiated filaments form a fine intertwined network or a conspicuous radial microfabric. Common oncoid nuclei are mollusks, plant debris or rock pebbles. The basic types correspond to porostromate as well as spongiostromate
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oncoids. Smaller and discoidal oncoids (e.g. Schnegglisteine in Lake Constance) exhibit a distinct lamination with alternating porous and dense carbonate layers reflecting seasonal growth (Pl. 12/5). Freshwater lake oncoids are characteristically concentrated in shallow, near-shore zones with only slight wave- or current-agitation. Their distribution allows climatically controlled lake level fluctuations to be reconstructed (Magny 1992). As shown by oncoids in the Lake Constance, changing textural and diagenetic features of the oncoids are related to differences in the composition and growth pattern of the microbes and algae that participate in the formation of the nodules. Modern lacustrine oncoids demonstrate that the formation of spongiostromate and porostromate oncoids does not necessarily need strong or repetitive turning of the nodule and that in-situ growth is a common feature. An interesting aspect related to the nutrient controls of oncoid-forming algae are freshwater oncoids created by industrial pollution (Dean and Eggleston 1984). Fluvial oncoids, formed by often highly diverse algae and cyanobacteria, are extremely variable in shape (ellipsoidal, discoidal, spherical, subcylindrical, irregular), size and internal structures (Glazek 1965, Golubic and Fischer 1975). Many oncoids exhibit excentric growth forms due to the contacts of the grains. The layers are often asymmetrically arranged, interrupted or obscured. Oncoid sizes range from smaller than 5 cm to larger than 30 cm; large oncoids are rather common. Many oncoids exhibit distinct concentric lamination and radial fibrous structures. Recent marine oncoids: In marginal-marine settings cyanoids originate in intertidal to shallow subtidal environments with normal, low or high salinity (Golubic and Campbell 1981; Richter et al. 1979; Neuser et al. 1982; Richter and Sedat 1983). Some spongiostromate oncoids are formed by trapping and binding of sediment on the surface of finely filamentous mats that disintegrate and roll on the substrate (Gebelein 1976; Jones and Goodbody 1985). Modern open-marine oncoids formed in intertidal to subtidal environments (e.g. Bermuda, Bahamas, Netherlands Antilles) often do not display distinct laminations, because algal growth patterns are continuously destroyed. The nuclei of these oncoids are provided by algal chips and skeletal grains from adjacent organisms living on seagrass. Oncoids with concentrically stacked laminae (spheroidal structures, type-SS structures of Logan et al. 1964) originate in the low intertidal zone exposed to waves, and in the agitated shallow-water zone below the low-water mark. The prime prerequisite for the growth of these oncoids is believed to be the overturning of the laminated bod-
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How to Describe Oncoids
Box 4.8. How to describe oncoids? Check your samples and use information from the field! Field criteria Depositional regime: Predominantly carbonate, mixed siliciclastic-carbonate or siliciclastic environments? Position of oncoid beds in the section: Occurrence in defined horizons (–> event beds, part of shallowing or deepening upward sequences)? Repetitive occurrence (–> cyclic patterns)? Occurrence in mud-dominated or grain-dominated limestones (–> different settings and environmental controls)? Abundance and distribution of oncoids: Either more or less than 50% of rock volume (–> oncolite or oncoid-bearing limestone). Densely packed (grain-to-grain contacts) or scattered oncoids (–> autochthonous and allochthonous oncoids)? Color of oncoids and matrix (–> may reflect differences in composition and/or weathering)? Sedimentary structures: Bedding, cross-bedding, bioturbation, grading, parallel orientation, imbrication, geopetal structures (–> clues to movement of oncoids and water energy)? Co-occurrence with stromatolites cryptalgal or fenestral structures, intergrowth of oncoids and stromatolites (–> intertidal and shallow subtidal environments)? Association with ooids and/or lithoclasts (–> reworking of oncoids)? Associated fossils (type, preservation, abundance; –> paleoenvironment)? Field and laboratory criteria Limestone texture: Wackestone/floatstone, packstone, grainstone/rudstone, bindstone? Matrix (homogeneous micrite, fine-bioclastic micrite, sparry calcite)? Oncoid types: Micrite oncoids, spongiostromate or/and porostromate oncoids (cyanoids), foraminiferal oncoids, composite oncoids? Single or multiple oncoids? Orientation of oncoids and other grains (–> current transport?) Size ranges, average, frequency distribution, sorting (–> occurrence and frequency of micro-, meso- or macro-oncoids)? Variations in shape, surface and growth forms (symmetrical, asymmetrical) of oncoids); ( –> water energy; in-situ growth)? Associated grains and fossils? Laboratory criteria Geopetal fabrics in oncoids (–> turning during growth, several growth phases)? Size: Ranges, average, dominating size fractions? Shape: Spherical, subspherical, ellipsoidal, discoidal, flattened (–> environmental and biological controls)? Contorted oncoids (–> soft oncoids, compacted oncoids )? Surface: Smooth, grooved, pustular, lobate; conspicuously round or rounded; bored; encrusted; preserved or eroded growth patterns (–> reworking and redeposition)? Individual growth patterns of oncoids: Changes in basic microfabrics and biotic contribution, multi-stage patterns (–> reflecting changes in environmental factors)? Borings: Totally bored or partially bored oncoids (–> alternations of hard and soft stages; early lithification during the growth of oncoids)? Cracks in the oncoids: Type, frequency, infilling (–> syn- or post-depositional origin, desiccation)? Cortex: Basic microfabric of the cortex: Lamination, meshwork (fine or coarse arrangement of sparry fenestrae, biotic contribution) or predominantly micritic? Layers and lamination: Layers and/or laminae; texture (distinct or vague, densely and loosely laminated; compound laminae and laminae sets, S-, I,- and R-type; see Fig. 4.15); configuration (concentric or non-concentric, continuous or non-discontinuous, tapering out of laminae); composition of the laminae composition (micritic, granular, biotic, sparry); surface of laminae (smooth, wavy, crenulated; disruption and micro-unconformities; impregnations, e.g. by glauconite, limonite); laminae thickness; inclusions of sediment between laminae; alternation of different laminae types (dark and light laminae, biotic and micritic laminae; organically-rich and organically-poor laminae; clay-rich and clay-poor laminae)? Biotic composition of the cortex: Taxonomic inventory, morphotypes (shape and size, abundance of different groups, diversity, successions), comparison with organisms outside the oncoids (–> biological controls, continuous/discontinuous growth, selective oncoid biota)? Nucleus: Present or absent, single or several nuclei, size, composition (bioclast, lithoclast, other grains), preservation, comparison of nuclei and non-oncoid grains with regard to grain types and biotic diversity (identical or different –> autochthonous/allochtonous oncoids)? The nucleus may be a complete or broken skeleton, an inorganic grain, or a soft particle which has subsequently decayed. Cortex and nucleus: Size and ratio; cortex larger or smaller than nucleus; different thickness of the cortex on the upper and lower side of the oncoid (–> in situ growth)? Is there a significant correlation between the size of the oncoids and the thickness of the cortex?
How to Describe Onocids
ies, generally accomplished by wave and current action; this allowing phototrophic growth to occur on all sides of the oncoids. Varying degrees of agitation and movement of the grains result in several geometric types (Fig. 4.15): SS-I structures (biconvex lenticular oncoids), SS-R oncoids (oncoids consisting of randomly arranged hemispheroids that make up the structure), and SS-C structures (characterized by concentrically stacked spheroids). How to describe oncoids and oncolites? The analysis of oncoids requires field data and careful studies of thin sections. Box 4.8 summarizes the criteria which should be examined. Dimensions and shape: Oncoid size is controlled both by environmental conditions and biological factors. The former include water energy and sedimentation rate, the latter metabolic processes of the biota involved in the oncoid formation, competition with grazing organisms and the size of the bioclastic or lithoclastic nuclei used. Marked variations in oncoid size within limestone horizons are important indicators of changes in deposition patterns, and are used to evaluate sequence boundaries as well as sea-level fluctuations. The size of modern freshwater oncoids ranges from smaller 2 mm to larger 100 mm. Ancient spongiostromate oncoids commonly vary from smaller 2 mm up to 50 mm. The average size of ancient porostromate oncoids is approximately 50 mm, but the range may extend up to more than 100 mm. Shape modifications both of oncoids and rhodoids can be best described by means of the Sneed and Folk sphericity-form diagram for particle shapes. The triangle diagram displays long, intermediate, and short axes and defines the basic form types (discoidal, spheroidal, and ellipsoidal). Many oncoids are elongate or subspherical in shape and may be flattened on one side. Laminated versus non-laminated micofabrics: Modern stromatolites and oncoids offer the possibility of understanding the controls of microstructural elements of oncoids. The basic lamination in ancient oncoids consists of couplets seen as dark, organic-rich micritic laminae, alternating with microspar or sparry as well as lighter colored laminae poor in organic matter, and often thicker than the dark laminae. The generation of individual laminae is intimately related to life activities of microbial assemblages. The generation of couplets with two distinct laminae members implies iterative changes in environmental variables that control the life activities of microbes, e.g. sedimentation rates, periodic deposition of sediment onto the growth surface -
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induced by tides or storms, day to night variations or seasonal changes in water chemistry resulting in changes in saturation and nutrients. Seasonal long-term environmental changes may affect water temperature, salinity, water level and turbulence, nutrients and sediment influx. The thickness of the laminae depends on the phototactic response of photosynthesizing cyanobacteria, the makeup of the microbiota, seasonal changes which may affect lithification and carbonate precipitation, and the modifications of initial laminae thickness by diagenesis. The microrelief along a single lamina, e.g. small-scale crenulations, tufts or domeshaped structures, is produced by particular microbes, but may also reflect diagenetic modifications. Descriptive subdivision of oncoids: Several attempts have been made to classify oncoids with laminated fabrics (e.g. Dahanayake 1977; Malchus and Kuss 1988; Kuss 1990). These classifications consider the composition of the laminae and the biota incorporated within the cortex. Because different lamina types reflect changes in water energy levels and water depths (e.g. Bowman 1983), a simple descriptive subdivision is proposed here (Fig. 4.15): The subdivision of the geometric types follows Logan et al. (1964), who classified oncoids (called ‘spheroidal stromatolites’) according to the arrangement of the laminae, reflected in turn by the shape of the nodules. Type C is characterized by concentrically arranged laminae, Type R by randomly arranged, non-continuous and overlapping laminae, and Type I by inverted laminae. In addition, a fourth Type L for randomly growing lobate forms with dome-shaped laminae is proposed here. The five compositional laminae types may occur in isolation, but many oncoids show multi-stage sequences consisting of several laminae types. These sequences are used to estimate hydrodynamic changes during oncoid growth, e.g. from quiet-water conditions (reflected by types 1 and 2 in Fig. 4.15) to moderate to high-energy conditions (type 5). Note that a given oncoid type can be transformed to another type by diagenetic alteration. The micrite laminae of Type 1 are microbial in origin or were formed by trapping of fine-grained sedimentary material (Pl. 11/2). This type is the most common microfabric in spongiostromate oncoids and is regarded as indicative of an environment with constant low-energy conditions, e.g. protected shallow subtidal or basinal oncoids. Some of these oncoids may be completely recrystallized porostromate oncoids. Type 1 oncoids are commonly sphaeroidal or ellipsoidal. The sparry spots of type 2 (corresponding to ‘grumous’ lami-
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nae, Dahanayake 1978) most commonly represent molds of calcimicrobes or algae, but in some cases may be caused by diagenetic alteration of peloidal fabrics. Type 2 oncoids are common in low-energy environments. Type 3 is of particular interest because it reflects the response of microbiota to short-term environmental changes. Type 4 is often characterized by sessile foraminifera or porostromate calcimicrobes (e.g. Girvanella). Girvanella tubes are usually visible in the exterior parts of the oncoids, and rapidly loose their shape to the inner oncoid parts.The encrusting foraminifera depend on hard substrates, and therefore testify to
Oncoids
the existence of sporadic lithification during oncoid growth. The diversity of type 4 can be both low and also very high, thus reflecting the appropriate conditions for the growth of epiphytes and epizoans in shallow subtidal environments with low sedimentation rates. The fabric of Type 5 has been called fenestral. Both Type 4 and Type 5 may lack a well-defined lamination or the meshwork characterizing Type 5 may form non-laminated oncoids. Non-laminated oncoids are rounded micritic grains commonly containing abundant relicts of calcified biota. Shapes are often irregular, lobate and not strictly
Plate 11 Oncoids: Calcareous Nodules Formed by Microbes, Algae and Other Encrusting Organisms Oncoids are mm- to cm-sized, rounded or irregularly formed grains consisting of a layered micritic cortex and a bio- or lithoclastic nucleus. The layers originate from crusts of sessile organisms (predominantly calcimicrobes and algae, but also foraminifera, bryozoans, serpulids and others). These organisms enhance and initiate the precipitation of calcium carbonate, act as encrusters and binders, trap fine-grained sediment or enhance synsedimentary lithification. Oncoids are common in marine and non-marine environments. Basic oncoid types are distinguished by the existence of skeletal or non-skeletal microfabrics (porostromate and spongiostromate oncoids) and by the prevailing biotic composition (e.g. microbial oncoids, foraminiferal oncoids). The shape, internal structures and biotic composition denote paleoenvironmental factors (e.g. water energy, sedimentation rates) and specific settings (see Box 4.11). The main settings of Phanerozoic oncoids have changed over time: Paleozoic spongiostromate oncoids grew predominantly in lacustrine and transitional marine environments; Cambrian to Jurassic porostromate oncoids flourished in marine subtidal environments. Starting in the midJurassic, porostromate oncoids were increasingly substituted for by spongiostromate oncoids, which became the only type in marginal-marine and particularly in lacustrine settings. In normal marine settings oncoids were replaced by rhodoids starting with the Late Cretaceous. Oncoids are commonly larger than ooids and vadoids (pisoids). They exhibit more irregular and often overlapping layers which are considerably thicker than the coatings of cortoid grains. 1 Modern oncoid formed by snowball-like sediment-trapping on the slimy surface of microbial films. A poorly developed laminar fabric is indicated by micritic areas (arrows). The oncoids with a maximum size of about 5 cm occur in quartzbearing calcareous sands on the coast of the Red Sea. Ras Muhammed, Sinai, Egypt. 2 Detail of –> 1. The oncoid consists of strongly micritized foraminifera (F), cortoids, rounded intraclasts and small angular quartz grains (Q). Note the open interparticle porosity (P). 3 Spheroidal spongiostromate/porostromate oncoids. The micritic layers of the cortex are accentuated by differences in color and the incorporation of spherical microfossils. Intergranular pores are occluded by sparitized calcitic rhombs indicating dedolomitization processes similar to those described from Middle Jurassic reservoir rocks in the Paris Basin (Purser 1985). The complex diagenetic history is characterized by partial dolomitization of oncoid floatstones associated with a replacment of matrix and peripheral parts of some grains, dissolution of dolomite crystals and filling of the rhombic molds by calcite. The limestones represent the upper part of a shallowing-upward cycle. Late Middle Jurassic: VrhnikaLogatec near Ljubljana, Slovenia. 4 Detail of –> 3. Vague remains of very thin, vertically arranged, branched filaments are recognizable within the gray layers (A). In contrast, horizontally growing Girvanella-like tubes occur in layer B. The dark-colored zone at the top of A point to cessations in the growth of the oncoid. 5 Foraminiferal oncoids. Irregularly shaped and partly contorted oncoids built by the encrusting foraminifera Nubecularia reicheli Rat. Bioclastic oncoid grainstone. Bioclasts used as nuclei are identical to the bioclasts occurring outside the oncoids. Reworking of oncoids is indicated by variations in shape and size (< 2 to 5 mm). These oncoids formed on highenergetic shoals at the boundary of the restricted and open-marine parts of a carbonate platform. Middle Jurassic (Bajocian): Bourgogne, northwestern France. 6 Nubecularia oncoid. The core (an echinoderm fragment, E) is surrounded by layers of densely spaced foraminiferal tests. The chambers exhibit thick walls that touch each other (arrows). Same sample as –> 5.
Plate 11: Oncoids
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130
Ancient Oncoids
Fig. 4.15. Descriptive subdivision of laminated oncoids in regard to growth patterns and the composition of the laminae. A - Basic geometric types of oncoids found in recent shallow-marine settings and related to varying degrees of agitation and movement of the grains (modified from Logan et al. 1964): Type C (characterized by Concentrically stacked spheroid layers, SS-C); Type R (oncoids consisting of Randomly arranged hemispheroidal layers); and Type I (biconvex lenticular oncoids exhibiting Inverted hemispheres, SS-I). Irregular random growth also can produce Lobate growth forms (Type L). Type I describes a particular case characterized by the interruption of oncoid growth, partial destruction and renewed growth after reworking and transport. B - Five types of laminae are commonly composed of: 1 - Laminated fabric consisting predominantly of micritic laminae (Pl. 12/7), 2 - Micritic laminae exhibiting tiny sparry spots whose distribution accentuates the laminar structure of the oncoid (Pl. 11/3), 3 - Couplets of micritic and sparry laminae (Pl. 12/3), 4 - Micritic laminae including or alternating with layers of recognizable encrusting microfossils (Pl. 119/3), 5 - Laminae exhibiting an open meshwork with spar-filled voids (e.g. Pl. 12/4, Pl. 118/3). Nucleus (black), micrite (stippled), sparry calcite (white).
spheroidal. The biota are represented by calcimicrobes, foraminifera and various skeletal metazoans with specific growth forms and succession patterns. Organisms contributing to Late Jurassic non-laminated oncoids include foraminifera, serpulids, bryozoans, chaetetid coralline sponges, Bacinella, Marinella, and porostromate calcimicrobes. Oncoids devoid of a continuous lamination have been called pseudoncoids (Dahanayake et al. 1976). Biotic composition of the cortex: The diversity of the microfossils in oncoids reflects small-scale environmental conditions that assist in interpretating longterm environmental trends (Scherze 2001). Autochthonous ‘foraminiferal-algal’ oncoids composed of intimately associated calcimicrobes and encrusting tubular foraminifera may indicate deeper-water settings (Peryt and Peryt 1975). Nucleus: The shape of the nuclei often determines the shape of the oncoid grain. A comparison of bioclasts acting as nuclei with the skeletal grains of the sample indicates whether the oncoids have been brought
to the depositional environment from other areas or whether the oncoids were formed within the depositional environment. Carozzi et al. (1983) uses the ‘index of biotic diversity of benthics’; the IBDB relates the number of groups or taxa forming the nuclei to the bioclasts occurring together with the oncoids. Ancient oncoids Most ancient oncoids consist of calcite, sometimes with minor admixtures of clay minerals, quartz and organic matter. The composition of these admixtures usually corresponds to that of the minerals occurring in the vicinity of the oncoids. Non-carbonate oncoids include phosphatic, ferruginous and manganese nodules as well as, less commonly, oncoids consisting of pyrite (Schieber 1999) or chlorite (Misik and Sucha 1997). The first three types are often associated with pelagic carbonates. The oncoids are found together with stromatolitic crusts on discontinuity surfaces, often corresponding to major stratigraphic gaps and related to eustatic highstand phases. The formation of these oncoids appears to be strongly
Ancient Non-Marine Oncoids
controlled by microbes which induce mineralization and act in trapping and binding sediment (Monty 1973; Champetier and Joussemet 1979). Different sources contribute to mineralization. Manganese nodules result from enrichment of deep oceanic bottom waters in manganese from submarine volcanisms, basalts and riftzone hydrothermal fluids. Manganese and ferromanganese oncoids, crusts and pavements are widely known throughout the Mesozoic of the Tethyan region in condensed pelagic limestones (Jenkyns 1970). Nodular phosphorites containing phosphatic oncoids originate under conditions of marine upwelling, but input of phosphorus from river drainage must also be taken into account (see Pl. 110). Common features of non-carbonate oncoids are variations in the contribution of encrusting organisms (particularly foraminifera: Greenslate 1974), the geometry and arrangement of oncoid laminae and the shape of the nodules. The size ranges from <1 mm to 2–10 cm (manganese micro- and macronodules: Pattan 1993). Examples of ancient ferruginous oncoids are limonitic Middle Jurassic nodules known as ‘snuff-boxes’ and believed to have been formed in an oxic environment with periodical current influence (Gatrall et al. 1972; Palmer and Wilson 1990). These oncoids are associated with episodes of reduced sedimentation. The oncoids exhibit asymmetric cortices. The original ferric composition of the laminae was probably under the control of non-photosynthetic microorganisms. Phosphatic oncoids exhibit dense laminated and microstromatolitic fabrics (Krajewski 1983; Martin-Algarra and Vera 1994). A genetic interpretation of phosphatic oncoids and stromatolites formed at discontinuity surfaces of drowned carbonate platforms is shown in Fig. 4.16. Ancient non-marine oncoids: Oncoids formed in ancient lake systems have been known since the Precambrian (Buck 1980), but have been studied more thoroughly only from the Tertiary (Box 4.9). The oncoids provide information about lake depositional systems and provide insights into the regional climate and lake level fluctuations (Awramik et al. 2001). Lacustrine oncoids are commonly associated with high-energy, often mixed siliciclastic-carbonate near-shore environments. They can mark the initiation of climatically induced transgressive phases. Conditions favorable for oncoid formation are high carbonate input associated with the increase in precipitation and the contribution of supersaturated groundwaters. Reworked oncoids indicate falls in lake level. Oncoids formed in soda lakes are known from Tertiary lake deposits (Dixit 1984). Freshwater and saline oncoids can be separated by dif-
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Fig. 4.16. Formation of phosphatic oncoids on a Mesozoic pelagic swell. The ellipsoidal and subspherical oncoids are several centimeter in size. The microstromatolitic laminae were formed by microbes, phosphate minerals, trapped pelagic micrite, and encrusting foraminifera. The nuclei of the microstromatolitic oncoids are mud clasts as shown in Fig. A: 1 growth of microbes (m) on a hardground (hg); 2 deposition of pelagic sediment and burial of microbial mats; 3 formation of clasts by burrowing and/or currents; 4 encrustation of isolated clasts and and rolling under moderate currents; 5 disymmetric upwards development of microbial lamination; 6 renewed growth of microbes (m) on the oncoid. Fig. B demonstrates the origin of endostromatolites in voids within the oncoids: 1 oncoid grown on a micritic clast with a glauconitic rim; 2 burrowing and boring (b) affecting the oncoid and the nucleus, growth of microbes in the cavity (m); 3 infilling by pelagic sediment (stippled). Modified from Martin-Algarra and Vera (1994).
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Box 4.9. Selected case studies of ancient lacustrine-fluvial spongiostromate and porostromate oncoids and oncolites. Pleistocene: Awramik et al. 2001; Brochier et al. 1986; Casanova 1985; Dixit 1984. Tertiary: Anadon and Zamarreno 1981; Bertrand-Sarfati et al. 1994; Bignot 1981; Bradley 1929; Colom 1993; Crouzel et al. 1972; Dixit 1984; Freytet and Plaziat 1982; Herbig 1991, 1994; Klähn 1926; Leinfelder and Hartkopf-Fröder 1990; Link et al. 1978; Llombart and Krauss 1982; Mäcker 1997; Moik 1990; Nickel 1983; Obrhel 1984; Ordonez et al. 1983; Rutte 1953; Speck 1949; Surdam and Stanley 1979; Truc 1978; Weiss 1969. Cretaceous: Carozzi 1989; Monty and Mas 1981. Jurassic: Leinfelder 1985. Triassic: Clemmensen 1978. Permian: Clausing 1990; Schäfer and Stapf 1978; Stapf 1989. Carboniferous: Bertrand-Sarfati and Fabre 1972; Obrhel 1977. Devonian: Donovan 1975; Fannin 1969; Obrhel 1976.
ferent types of bioclastic nuclei (e.g. unionid bivalves, gastropods, decapod crabs). Mollusk shells are more commonly used as nuclei for lithoclasts or gravels. ‘Hollow’ central parts of non-marine and marine oncoids or nuclei consisting of silt or sandy material can be caused by the disintegration and subsequent infilling of plant stems acting as nuclei (Jones and Goodbody 1985); those oncoids might be indicators of the existence of former macrophyte meadows. A common feature, particularly of larger lacustrine oncoids, are fine radial structures (Pl. 12/7) sometimes emphasized by vertically arranged micro-cracks. Some of these radial structures are inaugurated by vertical growth patterns of algae. The composition of non-marine or extremely marginal marine oncoids depends on the growth patterns of algae, cyanobacteria and sometimes also of foraminifera and other organisms. The abundance and shape of oncoids is only sometimes hydrodynamically controlled. Many lacustrine oncoids originate from in-situ growth. Lacustrine oncoids have ellipsoidal and spheroidal, sometimes cylindrical shapes. Common sizes vary between 15 and about 200 mm. Sphaeroidal oncoids are generally attributed to fluvial channels and bottom lags of ponds; more irregular forms occur preferrentially in shallow parts of lakes (e.g. Nickel 1983, Casanova 1985), but these relations should not be transferred to all ancient examples (Leinfelder and Hartkopf-Fröder 1990).
Ancient Marine Oncoids
As opposed to marine oncoids, lacustrine and fluvial oncoids often exhibit distinct sets composed of morphologically identical laminae, e.g. couplets of dark and light laminae. Alternating thin micritic laminae and thicker laminae with vertical structures reflect the different compositions of the original algal associations (Monty and Mas 1981; Freytet and Plaziat 1982; Moik 1990; see Pl. 8/3). Laminae thickness and micromorphology depend on the interplay between carbonate precipitation and sediment binding on algal filaments and the decay of the benthic microbial communities (Golubic 1976; Burne and Moore 1987). Binding and inorganic carbonate precipitation appears to be quantitatively more effective than biologically induced precipitation (Pentecost 1978). As opposed to marine oncoids, studies of non-marine oncoids allow a higher resolution of the development and time involved in environmental changes. Growth types of freshwater and brackish oncoids and stromatolitic structures are powerful aids in reconstructing the depositional history of lacustrine, fluvial and deltaic sediments (e.g. Leinfelder 1985). The form, size, and microstructures are used to distinguish depositional zones of fluvial systems, e.g. channel base, abandoned channels, or the mouthes of rivers (Anadon and Zamarreno 1981; Nickel 1983; Mäcker 1997). Distributional patterns of oncoids assist in differenting specific lacustrine and fluvial depositional settings in coastal playa lakes and hypersaline lakes (Herbig 1994). Common microfacies criteria of non-marine oncoid-bearing limestones are micrite, peloids, black intraclasts, small radial ooids, stromatolites, gastropod and other mollusk shells, and ostracods (Link and Awramik 1978). Diagnostic criteria for ancient freshwater oncoids are nuclei formed by freshwater mollusks, lamination patterns analogous to modern freshwater oncoids, well-preserved microstructures, occurrence together with terrestrial sediments, and often large sizes. Ancient marine oncoids: Phanerozoic marine oncoids were formed in at least eight different settings including (1) tidally influenced marginal-marine environments, (2) open-marine shelf lagoons, (3) the environs of platform patch reefs, (4) back-reef areas, (5) associated with platform margin reefs, (6) on the upper slope, (7) basins, and (8) on pelagic platforms. These settings were best studied for Late Jurassic carbonates, particularly in Switzerland and France. Some of these settings can be differentiated by the types and composition of the oncoids, and in the associated sedimentary facies. Fig. 4.17 is a simplified summary of the main criteria of Late Jurassic oncoites and oncoid-bearing limestones. Oncoids are frequent constituents of
Depositonal Environment of Late Jurassic Oncoids
Location
1
2
3
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4
5
6
7
8
Depositional texture
Wackestone Packstone Bindstone
Wackestone Packstone Floatstone
Wackestone Packstone Floatstone
Grainstone Rudstone Wackestone
Packstone Rudstone
Wackestone Floatstone
Floatstone
Grainstone Packstone
Associated grains
Peloids Cortoids Stromatolites Birdseyes
Bioclasts
Bioclasts Peloids
Cortoids Bioclasts Intraclasts
Bioclasts Lithoclasts
Bioclasts
Bioclasts
Peloids
Associated fossils
Foraminifera Algae
Mollusks Foraminifera Algae
Foraminifera Reef biota
Foraminifera Calcar. algae Mollusks
Reef biota
Mollusks Echinoderms
Ammonites
Planktonic foraminifera
Type of oncoid
Simple Multiple
Simple
Simple Multiple
Simple
Simple
Simple
Simple
Simple
Prevailing size
10 - 50 mm
< 2 mm 5 -60 mm
2 - >10 mm
2 - 10 mm
2 - 20 mm
10 - 20 mm
30 - 300 mm
1 - 2 mm
Growth type
Concentric Lobate Radial
Lobate Concentric
Concentric Radial Lobate
Concentric
Concentric Inverse
Elliptical
Concentric Radial Inverse
Concentric
Basic types of oncoids
Cortex: Laminated Non-laminated Lamination type Biotic composition Borings
M Por > Sp Comp
L NL 1, 2, 5 Calcimicrobes Algae +
Nucleus
Lithoclasts Algae
Remarks
Mixed types Eroded and broken oncoids
Por > Sp M Comp Foraminifera L 1, 2, 3 Calcimicrobes Serpulids Foraminifera + Bivalves Corals Echinoderms ––
Por > Sp Comp
L NL 1, 2, 4, 5 Calcimicrobes Highly diverse metazoans + Corals Sponges Mollusks Echinoderms ––
Sp M
L NL 1, 3, 4, 5 Calcimicrobes Low-diverse metazoans –– Corals Mollusks Algae ––
Por Comp
M
L
Sp M
M
L
L NL
1, 3
1
NL 4, 3 Calcimicrobes
–– Corals Echinoderms Mollusks ––
-Calcimicrobes Serpulids Mollusks +
Microbes
––
Microbes
––
Cephalopods Echinoderms
Cephalopods Lithoclasts
Peloids
Glauconite
Ferruginous coatings
Ferruginous + glauconite coatings
Fig. 4.17. Depositional environments and characteristics of Late Jurassic (Oxfordian-Kimmeridgian) oncoids. The criteria are ordered according to their frequency in about 15 case studies. Although there are many deviations from this scheme, differences between shallow- and deeper-water oncoids are evident. Associated fossils, shapes and growth types, the lamination patterns and the type of the bioclasts forming the nuclei are of particular interest. Oncoid sizes reflect changing environmental conditions rather than specific depositional sites. Information on oncoids is provided by studies of the Late Jurassic ‘Hauptmumienbank’, which can be followed in northern Switzerland and France for at least 100 km. These studies indicate paleowater depths of a few meters for lagoonal oncoids, about 50 m for oncoids formed on the upper slope, and > 100 m for basinal oncoids. Based mainly on Dahanayake (1983), Fezer (1988), Gasche (1956), Gasiewicz (1983), Gygi and Persoz (1987), Gygi (1992), Kutek and Radwanski (1965), Leinfelder and Schmid (2000), Massari (1983), Pümpin (1965), Scherze (2001) and Schmid (1996). Comp = composite oncoid, M = micritic oncoid, Por = porostromate oncoid, Sp = Spongiostromate oncoid.
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Jurassic shelf carbonates. Most of these settings are also documented by non-Jurassic examples (see Box 4.9). Oncoids formed in upper slope and basinal settings are not known prior to the Permian. Reef oncoids: Concentrically structured porostromate and micritic oncoids are common constituents of Paleozoic reef limestones (e.g. Devonian stromatoporoid-coral reefs in western Canada: Machielse 1972; late Permian coralline sponge-Archaeolithoporella reefs of the Tethys region: Sano et al. 1990). Oncoids are common constituents of Devonian back-reef sands and fore-reef sediments. Locally porostromate oncoids grow in place on the slope, sometimes together with stromatolites. Very good examples are the capped oncoids formed by Rothpletzella (Playford et al. 1976). Differences in the biotic composition of the oncoids give hints of changes in sea level and water depths. Pelagic oncoids: These spongiostromate grains, formed on the tops of seamounts and drowned platforms, became more frequent during the Jurassic. Ex-
Precambrian and Phanerozoic Oncoids
amples are known from the Liassic and Late Jurassic of the Tethyan region (originally described as ‘pelagic oolites’: Jenkyns 1972; Massari 1983). The oncoids are characterized by polyphase coatings consisting of micritic and granular laminae, the latter composed of finegrained skeletal debris. Encrusting organisms (e.g. Nubecularia) are common. Nuclei are bioclasts (e.g. fragments of ammonites, sponges) and lithoclasts. The sizes vary between <2mm and 100 mm. The oncoids occur in association with microstromatolites and originated through sediment-binding and/or mineral-precipitating activity of microbes in regions with low sediment input, probably at storm-wave depth. Box 4.10 lists important case studies of Cambrian to Cretaceous marine oncoids. Precambrian and Phanerozoic oncoids Oncoids occur already in Archean carbonates, together with stromatolites and ooids (Hofmann et al. 1999; Grotzinger and James 2000). The grain type is a common constituent of Proterozoic platform carbon-
Box 4.10. Selected case studies of ancient marine spongiostromate and porostromate oncolites. In contrast to modern oncoids, these oncoids were common in Paleozoic and Mesozoic subtidal and intertidal environments as well as in slope and basinal settings. The last marine cyanoids are known from the Cretaceous. Cyanoids were replaced by coralline algal rhodoids during the Late Cretaceous and the Tertiary. Cretaceous: Carozzi 1989 (tidal setting; oncoid bars on shallow carbonate platform); Lewy 1972; Masse 1979; Weiss 1969. Jurassic: Bollinger and Burri 1970 (patch reefs); Colacicchi et al. 1975 (tidal setting); Dahanayake 1977, 1978, 1983 (lagoonal carbonates); Dangeard 1935; Dragastan 1969; Elliott 1966; Fezer 1988 (platform margins and openmarine lagoonal settings, lagoonal carbonates); Gasche 1956; Gasiewicz 1983 (peritidal); Gygi 1969, 1992 (openmarine lagoon, slope, basin); Heim 1916; Hug 1999; Jenkyns 1972 (pelagic platform); Kutek and Radwanski 1965 (peritidal and basinal settings); Kuss 1990; Leinfelder et al. 1993; Leinfelder and Schmid 2000; Massari 1983; Massari and Dieni 1983 (pelagic platform); Misik 1966; Pümpin 1965 (reef, back-reef sands); Purser 1975 (peritidal); Scherze 2001 (patch reef);. Schmid 1996 (lagoonal and reef settings) Triassic: Biddle 1983 (reef-derived material); Catalov 1970, 1983 (platforms); Fischer 1964 (tidal setting); Goldhammer et al. 1990 (tidal setting); Hagemeister 1988; Jerz 1966 (siliciclastic-carbonate inner shelf); Kisten et al. 1990; Peryt 1977, 1980; Radwanski 1968; Schuler 1968; Tichy 1983. Permian: Flügel 1979 (lagoonal and reef settings); Füchtbauer 1968 (near-reef); Kerkmann 1966 (reefs, back-reef sands); Mazzullo and Cys 1983; Peryt 1981; Peryt and Wagner 1981 (basinal setting); Rothe 1969; Termier and Termier 1955 (reef), Toomey 1974, 1983, Toomey et al. 1988, 1989 (subtidal carbonate shelves, open-marine and restricted shelf lagoons). Carboniferous: Bowman 1983 (subtidal platform); Burchette and Riding 1977; Carozzi 1989 (tidal setting; shoals in inner ramp position); Gerhard 1985; Henbest 1963; Johnson 1946; Neal 1969; Obrhel 1977; Perret 1971; Toomey 1975; Wright 1983 (peritidal setting) Devonian: Kaufmann 1964; Machielse 1972 (back-reef sands and forereef setting); Pia 1932; Playford and Cockbain 1966 (forereef setting at platform margin); Poncet and Lapparent 1975 (siliciclastic-carbonate inner shelf); Wolf 1965 (reef). Silurian: Ratcliffe 1988 (shelf lagoon); Stel and De Coo 1977 (patch reef); Toomey and LeMone 1977. Ordovician: Benedict and Walker 1978; Blackwell et al. 1984 (lagoonal setting); Carozzi 1989 (tidal setting); Maslov 1960; Wood 1948. Cambrian: Gandin and Debrenne 1984 (associated with archaeocyathid patch reefs); Halley 1975 (peritidal setting); Hose 1961; Youngs 1977; Zhuravleva 1964.
Significance of Oncoids
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ates together with ‘katagraphs’ (non-laminated carbonate structures of different size, and a structure consisting of micritic particles and peloids: Zhuravleva 1964, 1979). Conspicuous criteria of many Precambrian oncoids are their large size (up to tens of centimeters) and radialfibrous microfabrics interpreted as bacterially induced structures. Until the Vendian-Cambrian transition, oncoids are represented by spongiostromate and micritic growth types. The first calcimicrobes forming porostromate oncoids appear in the Late Vendian. Peryt (1981) suggested a distinct change in the major oncoid types and their environments during the Phanerozoic. He assumes that Paleozoic spongiostromate oncoids were predominantly formed in lacustrine and transitional marine environments, whereas porostromate Girvanella oncoids flourished in marine subtidal environments from the Cambrian to the Jurassic. Starting in the mid-Jurassic, porostromate oncoids were replaced by spongiostromate oncoids, which became the only type in marine settings and continued to flourish in lacustrine settings. Monty (1979) made the point, that spongiostromate oncoids were substituted in marine settings by rhodoids starting in the Early Tertiary. This concept is supported by the predominance of cyanoids in Paleozoic shelf carbonates. However, porostromate and spongiostromate oncoids may occur in close association in the same beds, pointing to a much stronger biological control on oncoid formation than previously thought.
Significance of oncoids in microfacies analysis Oncoids provide important microfacies data. Box 4.11 summarizes some significant points. Two case studies Early Permian oncolites and rhodoliths of the Carnic Alps, Austria – different composition and different setting (Fig. 4.18A, Pl. 12/1): Oncolite beds consisting of porostromate meso-sized oncoids with ranges between 10 and >50 mm are intercalated in shallow-marine inner and outer platform carbonates. Four types can be recognized according to their biotic composition: (1) Girvanella-foraminiferal oncoids of near-coast settings with strong siliciclastic input (Fig. 4.18A); (2) large composite oncoids (size up to 70 mm) composed of Girvanella and Claracrusta of open-marine inner platforms with minor siliciclastic input; (3) large red algal-foraminiferal rhodoids made of Archaeolithophyllum and calcitornellid foraminifera, and associated with fusulinids, dasyclad algae and mollusks, occurring in open platform settings; and (4) centimeter-sized spherical Archaeolithoporella-oncoids, associated with
Fig. 4.18. Weathered surfaces of oncoid limestones. A: Oncolite (oncoid packstone) composed of spheroidal and ellipsoidal meso- and macro-sized oncoids embedded within argillaceous bioclastic micrite. Most oncoids are Girvanella oncoids (cyanoids) or composite oncoids made by cyanobacteria and encrusting foraminifera. Note the high amount of small broken oncoid debris between the larger well-preserved oncoids and aligned oncoids, indicating reworking by currents. Early Permian: Carnic Alps, Carinthia, Austria. B: Oncoid wackestone with variously sized and poorly sorted ellipsoidal oncoids. The argillaceous carbonate matrix contains fine-bioclastic debris. Note the large size of the nuclei and the thin micritic coatings. Late Triassic (Carnian): Obir, Karawanken Mountains, Carinthia, Austria.
cement-rich platform-margin reefs. Similar types occur in comparable settings in Late Carboniferous and Early Permian platform carbonates of the Tethys and in the United States (e.g. Toomey 1983; Toomey et al. 1988, 1989). Late Triassic oncoid-bearing beds of the Calcareous Alps, Austria – regional marker beds (Fig. 4.18B): The huge middle and upper Late Triassic carbonate platforms of the Alps are separated by Carnian siliciclastic-
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Significance of Oncoids
Box 4.11. Significance of oncoids and oncoid limestones Paleoenvironmental proxies Biotic composition and lamination types reflect environmental factors that control microbial and algal growth as well as encrusting communities (light, salinity, water energy, sediment input). The availability of light is not necessarily a limiting factor because cyanobacteria and various green algae also grow in areas of low light intensities (e.g. on the undersides of oncoids). Salinity: Oncoids originate in freshwater, brackish waters and marine waters. Ancient freshwater and brackish-water oncoids are differentiated from marine oncoids by nuclei formed by fresh water mollusks, specific lamination patterns analogous to modern freshwater oncoids, better preserved microstructures and the association with terrestrial sediments. Water-energy levels: Oncoids are often used as indicators of high energy, low intertidal to shallow subtidal environments, because it is thought that oncoids require frequent overturning and rolling to form, and that the degree of turbulence determines oncoid shapes. These generalizations do not hold in many freshwater and subtidal environments, where spongiostromate and porostromate oncoids originate without significant rolling but nonetheless exhibit concentric laminae. Stationary growth is shown by asymmetrical shapes and asymmetrical widths of laminations. Rolling appears to be indicated by subspherical shapes, well-laminated cortices, concentric symmetrical growth patterns, and abrasion of algal or cyanobacterial structures. Cessations in rolling are manifested by strongly encrusted laminae surfaces. Lack of rolling is revealed by branched and lobate oncoid shapes. Multistage oncoids exhibiting a change from laminar to lobate growth forms record a change from common rolling to less frequent rolling. Low-energy and highenergy oncoids are better differentiated by the combined approach of oncoid criteria and texture types: Low-energy, quiet-water conditions are indicated by large micritic single and multiple oncoids that exhibit discontinuous growth layers (oncoid type R, Fig. 4.15) and form floatstones and wackestones. Quiet-water conditions are also indicated nonlaminated or only weakly laminated meshwork oncoids (type 5, Fig. 4.15). Higher-energy environments are reflected by small oncoids, that exhibit regular laminations and form packstones and grainstones. Redeposition of oncoids is indicated by smoothing and breaking, co-occurrence of different oncoid types, and oncoids found in a matrix different than the micrite forming the oncoids. Paleo-water depths are reflected in biotic composition, nuclei and lamination types. Deeper-water oncoids are often un-laminated or have densely spaced micritic laminae (e.g. in the Permian Zechstein). This is a favorable setting for continuous oncoid growth. The slowly growing Girvanella oncoids are considered indicators of reduced sedimentation, preferrably in deeper settings at depths of some tens of meters. Spongiostromate oncoids develop over shorter periods in unstable ecological conditions. Oncolite beds, intercalated in platform sequences, can reflect breaks or hiatuses in sedimentation. Short-term alternations of stationary phases and oncoid-building phases are shown by changes in lamination and encrustation types, micro-unconformities of laminae surfaces, and the incorporation of sediment within the oncoid. Sea-level fluctuations Shallowing-upward sequences and cyclic depositional patterns: Shallowing is indicated by stromatolitic crusts overgrowing oncoids, marked changes in size, shape and composition of oncoids, and differences in depositional texture types. A good example was described by Ratcliffe (1988). Oncoids are characteristic constituents of shallow subtidal members of muddy and grainy platform cycles (e.g. Fischer 1964; Goldhammer et al. 1990; Purser 1975). Transgressive and regressive phases: Oncoids often occur at the base of transgressive sequences formed in shallow low-energy environments (e.g. Wright 1983) and mark maximum flooding surfaces. An oncolitic wackestone facies with whole fossils may indicate regressive conditions. Stratigraphic gaps and condensed sequences: Breaks in sedimentation are manifested by non-carbonate oncoids and by oncoids with ferruginous or glauconitic coatings, associated with microstromatolites, hardgrounds and lithoclastic beds. Regional correlations In ramp and platform settings, oncolite limestones can be excellent marker beds that can be followed laterally for tens and hundreds of kilometers. Examples are known from the Late Triassic (Fig. 4.18B) and the Jurassic. Depositional settings of marine oncolites Compositional and morphological features of oncoids, and their association with other grains and fossils are valuable criteria in recognizing depositional settings on platforms, slopes and basins (Fig. 4.17). Reservoir rocks Micritic oncolitic limestones are regarded as poor reservoir rocks, but exceptions are known, e.g. oncoid limestones associated with microbial reef carbonates (Late Jurassic Smackover Formation in the eastern Gulf Region: Becher and Moore 1979), and productive Cretaceous platform carbonates of Texas (Bebout et al. 1979) and offshore Brazil (Carozzi and Falkenhein 1983). The critical differentiation of oncoids and other coated grains is important; misleading environmental interpretations and exploration targets may result if oncolites are confused with subaerial fabrics (Gerhard 1985).
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Rhodoids and Macroids
carbonate intervals (Raibl Formation). Oncolite beds appear in several cycles, but are concentrated at the base and near the top of the shaly-sandy sequence. These beds have a thickness of some centimeters to several meters and alternate with coquina limestones and lime mudstones. Lower and upper oncolite beds differ in the size and growth forms of the oncoids, the material used as nuclei, and the biotic associations (bivalves, foraminifera, echinoderms, ammonites). The size of the oncoids found near the base of the sequence increases from the base to the top. Most oncoids are cyanoids and composite oncoids built by Girvanella and encrusting foraminifera. The laminae correspond to the lamination Types 1 and 4; calcareous laminae alternate with clay laminae. The upper horizon is a marly finely bioclastic oncoid packstone (Fig. 4.18B) with average oncoid sizes of rarely > 10 mm. Most oncoids 2-3 mm thick, between 3 and 15 mm long. The growth forms of the ellipsoidal porostromate and composite oncoids correspond to Types C, R, and L; the lamination pattern to Types 1, 2 and 4 (Fig. 4.15). Similar to the lower oncoid horizon, oncoid sizes increase towards the top. Oncoid growth was terminated by strong terrigenous
input. Both oncoid horizons can be laterally followed for more than 200 kilometers. The depositional environment is interpreted as deeper subtidal lagoons with reduced sedimentation: Quiet-water conditions are indicated by the association with marly sediments and the frequency of large lobate and asymmetrical oncoids. A deeper setting is indicated by the common association with crinoids and echinoids (which also provide most nuclei). In some regions, the oncoid originated in well-lit areas behind local ooid shoals. Microfacies and stratigraphic sequence suggest a general deepening-upward trend for these sections, whereby the oncoid beds mark regional sequence boundaries related to global sea-level rises during the Middle and Late Carnian. 4.2.4.2 Rhodoids and Macroids Rhodoids are unattached nodules composed predominantly of free-living encrusting calcareous red algae. Cenozoic rhodoids are commonly composed of coralline algae with High-Mg calcite skeletons (Pl. 54/1),
Fig. 4.19. Composite red algal-foraminiferal rhodoid (A). The nucleus is a rhodoid clast exhibiting slightly columnar growth forms. The outer part of the nodule is characterized by alternating concentric layers of encrusting acervulinid foraminifera (light laminae; C, D) and laminar coralline algae (dark laminae; E). The thickness of the foraminiferal laminae varies conspicuously. Macroborings (B) are restricted to the rhodoid clast, indicating lithification prior to the settlement of encrusting foraminifera. The borings are infilled with fine-grained and peloidal sediment. The lengths of these commonly ellipsoidal nodules range from 10 to 70 mm, and are frequently between 30 and 50 mm. The widths range from 5 to 30 mm (frequently 10-20 mm). Modern acervulinid foraminifera and acervulinid macroids preferentially occur in deeper waters of reef slopes down to about 70 m. Similar water depths are inferred for the rhodoid floatstones represented by the sample. These white- to cream-colored rhodoid limestones are widely used as decorative stones in the Mediterranean area. Tertiary: Northern Italy.
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but peyssoneliacean red algae with aragonitic skeletons (Pl. 64/10) can be of importance, too (Basso 1990). Late Paleozoic rhodoids were formed by ancestral red algae (Pl. 56). Centimeter-sized unattached nodules, predominantly consisting of encrusting organisms other than coralline algae (e.g. foraminifera), are called macroids. Distribution of modern rhodoids: Rhodoids are known from tropical and subtropical, temperate and polar seas. Living rhodoids occur in intertidal pools (Wehrmann et al. 1995) to subtidal shelves below normal wave base, down to depths of more than 200 m. Deeper-water settings are often dominated by macroids (e.g. outer Florida shelf: Prager and Ginsburg 1989). In tropical warm-water settings rhodoids (30 °S to 20 °N) occur in shallow-water environments (lagoons, tidal channels, reef flats, back-reef areas; seagrass beds), at shelf margins as well as in deeper forereef environments on terraces (Minoura and Nakamori 1982; Scoffin et al. 1985; Gischler and Pisera 1999). The common lower limit in the tropics is about 80 m, but living rhodoids also occur in depths between 160 and 180 m, too (pers. comm. W.-Chr. Dullo). Temperate rhodoids are known from the warm-temperate Mediterranean Sea in depths between about 35 and 150 m (e.g. Basso 1998), and from the transition between tropical and temperate waters off southern Queensland between 50 and 110 m (Lund et al. 2000). These rhodoids contribute to the widely distributed rhodalgal facies which interfingers with the bryomol facies (Sect. 12.2). On rocky coastal slopes of the cool-temperate Atlantic, living rhodoids grow in depths between about 20 and 40 m, in higher latitudes between 5 to 25 m (Tromsö/Norway: pers. comm. A. Freiwald). A characteristic sediment consisting of coralline algal branches, rhodoids and their detritus, is ‘maerl’ (Sect. 2.4.4.3). Wide and thick accumulations of rhodoids with only few detritus are called ‘rhodolith pavements’. In polar regions living corallinacean algae were reported from depths between 60 m to about 100 m (Svalbard shelf: Andruleit et al. 1996; northern Norwegian shelf). These rhodoids, however, exhibit only thin crusts. Some caution is necessary in a too rigorous use of ‘modern’ rhodoids for paleobathymetrical interpretations, as seen in originally Pleistocene ‘relic rhodoids’ which formed in shallow-marine settings and now are found in deep-water settings (e.g. South African shelf: Siesser 1972). Controls: The most important control on the growth of coralline algae and the formation of rhodoids is light. Light intensity is correlated with water depth and geographic latitude. A change in water depth combined with
Rhodoids and Macroids
sea-level fluctuations is an important factor for the biotic composition and internal growth patterns of rhodoids. Water temperature controls the composition of algal communities and their biogeographic distribution. Hydrodynamic energy is regarded as another very important ecological factor for rhodoid distribution. It controls external and internal growth forms and the success of taxonomic successions. Salinity control is of minor importance, because most calcareous coralline red algae are adapted to normal marine environments. The co-existence of corallines and cyanobacteria, however, indicates the possibility that brackish-water conditions can also be tolerated (Richter and Sedat 1983). Vertical accretion rates are high for tropical rhodoids (about 1-2 mm/year) and an order of magnitude lower for temperate and deeper-water rhodoids (Matsuda 1989). Growth rates of thin and thick crusts depend strongly on the effect of grazing herbivorous organisms (Steneck 1985). How to describe rhodoids and macroids? The methods for analyzing of modern and Tertiary rhodoids and macroids are firmly established (Bosellini and Ginsburg 1971; Bosence 1983a, 1983b, 1991). The main criteria are external shape, internal growth patterns, size, and the biotic composition of the cortices. In addition, nuclei types and associated fauna are of importance. Box 4.12 summarizes the features which should be checked. External shapes and internal growth patterns: The external shape depends on the internal growth patterns of the algae, and the frequency of turning (Braga and Martin 1988). Growth forms are apparently controlled by the frequency of overturning. Increasing overturning causes flattening of branches which join laterally. In tropical and subtropical climates branched rhodoids are formed in shallow and quiet waters; in sheltered areas they are spherical; in higher energy areas flattened ellipsoidal shapes dominate. Rhodoliths with consistent internal growth patterns indicate stable environmental conditions and/or fast growth. Ellipsoidal, spheroidal and laminar growth patterns are generally considered to be indicative of higher water energy. Discoidal and flat external forms, and branching and columnar internal growth forms are common in low-energy environments (Piller and Rasser 1996). Size: Modern rhodoids range in size from 10 to about 250 mm, more frequently from about 10 to 100 mm. Ancient coralline rhodoids have about the same size. The size of ancient rhodoids appears to be strongly controlled by the composition of the algal communities
Describing Rhodoids
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Box 4.12. How to describe rhodoids? A combination of field and laboratory criteria is recommended. Arrows indicate palecological indications and clues for interpreting paleo-environmental and depositional settings. Field criteria Depositional regime: Predominantly carbonate, mixed siliciclastic-carbonate or siliciclastic? Position of rhodoid beds in the section: Occurrence within defined horizons (–> event beds, part of shallowing or deepening upward sequences). Repetitive occurrence (–> cyclic patterns)? Occurrence in mud-dominated or grain-dominated limestones. (–> different settings, water depths and environmental controls)? Abundance and distribution of rhodoids: Either more than or less than 50% of rock volume (–> rhodolite or rhodoidbearing limestone). Densely packed (grain-to-grain contacts) or scattered? Color of rhodoids and matrix (–> differences in composition and/or weathering)? Sedimentary structures: Bedding, cross-bedding, bioturbation, grading, parallel orientation, imbrication, geopetal structures (–> movement of rhodoids)? Field and laboratory criteria Limestone texture: Wackestone/floatstone, packstone, grainstone/rudstone, bindstone? Microfabrics: Single or multiple rhodoids Composite rhodoids (e.g. algal-foraminiferal macroids)? Orientation of rhodoids and other grains (–> current transport?) Size ranges, average, frequency distribution, sorting (–> micro-, meso- or macroids)? Variations in shape, sphericity, surface and growth forms (symmetrical, asymmetrical) of rhodoids (–> hydrodynamic conditions)? Association with ooids and/or lithoclasts (–> reworking of rhodoids)? Associated fossils? Laboratory criteria Geopetal fabrics in rhodoids (–> overturning during growth)? Shape, external growth forms and sphericity: Spheroidal, ellipsoidal, discoidal (Folk and Sneath triangle diagram (–> hydrodynamic energy; environmental and biological controls). Symmetric or asymmetric growth? Contorted rhodoids? Surface: Smooth, bumpy, grooved, lobate; conspicuously round or rounded; bored; encrusted; surface preserved or eroded (–> reworking and redeposition)? Internal growth patterns: Laminar, branched, columnar, globular; densely or openly spaced branches (–> hydrodynamic energy). Mixed or multi-stage microfabrics? Layers and lamination: Configuration (concentric or non-concentric, continuous or discontinuous); thin or thick lamina crusts? Size: Dominating sizes, ranges and distribution; arithmetic mean diameter: (long + intermediate + short axis lengths):3. Biotic composition of coatings: Taxonomic inventory (–> paleo-water depths), mono- or multispecific morphotypes including shape and size, abundance of different groups, rhodoid diversity in relation to rhodoid size, type and frequency of encrusting organisms; comparison with organisms occurring outside the rhodoids (–> biological controls, continuous/discontinuous growth, selective biota) Ecologic successions and growth patterns of rhodoids: Changes in basic microfabrics and biotic contribution (–> changes in environmental factors) Borings: Totally or partially bored (–> hard and soft stages) Cracks within the oncoids: Type, frequency, infilling (–> desiccation, syn- or post-depositional origin) Nucleus: Present or absent; single or several nuclei; size; composition (bioclast, lithoclast, other grains; fragments of coralline algae); preservation of nuclei; comparison of nuclei and non-oncoid grains with regard to grain types and biotic diversity (identical, different –> autochthonous/allochthonous rhodoids) Cortex and nucleus: Size and ratio; different thickness of cortex on the top and bottom of the rhodoids (–> in-situ growth)? Associated fauna?
such as shown by different sizes of Late Paleozoic Archaeolithoporella and Archaeolithophyllum rhodoids. Biotic composition: Taxonomic successions may be the result of a change in rhodolith size: the larger the size, the higher the stability of communities over time.
But whether the succession of algal species represents changing environments or a change in rhodolith size is still an open question. Monospecific branching rhodoids occur both in temperate and arctic waters as well as in tropical environments.
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Ancient rhodoids and macroids Most fossil rhodoids are known from Tertiary shelf carbonates and reefs (Box 4.13). Coralline algae appear in the Late Jurassic. Only a few rhodoids have been reported from the Cretaceous (e.g. Early Cenomanian solenoporacean-corallinacean rhodoids, reef flat facies). Late Carboniferous and Permian rhodoids (Pl. 56/8) differ from Cenozoic rhodoids in their preference for near-coastal, very shallow settings.
Cyanoids and Rhodoids
Significance of rhodoids in microfacies analysis Rhodalgal limestones represented by shelf carbonates and reefs with large contributions of coralline algae are, volumetrically, the most important facies in the Miocene Mediterranean (Esteban 1996). The rhodalgal-type grain associations are indicative of platforms developing in temperate seas or in subtropical to tropic areas with anomalous (e.g. cooler, eutrophic, upwelling) conditions (Carannante and Simone 1996; Sect. 12.2).
Plate 12 Cyanoids and Rhodoids: Cyanobacterial and Red Algal Nodules Cyanoids (–> 1) are a specific type of porostromate oncoids, characterized by recognizable tubular microstructures which can be attributed to calcified cyanobacteria. Cyanoids are common constituents of Paleozoic shallow-marine limestones. Rhodoids (rhodoliths; –> 6) are nodules formed by encrusting (mainly rhodophycean) red algae. The freeliving encrusting calcareous algae exhibit characteristic finely meshwork structures (see Pl. 54). Other organisms contributing to the formation of nodules are encrusting foraminifera that form centimeter-sized macroids. Living rhodoids occur in tropical, temperate and polar settings and are found from tidal zones down to more than 200 m. They contribute to the formation of platforms, ramps and reefs. Ancient rhodalgal-type grain associations are often indicative of platforms formed in temperate seas or in subtropical to tropic areas, sometimes with anomalous (e.g. cool-water, upwelling and nutrient-rich) conditions. Diagnostic descriptive criteria are the external shape, internal growth patterns, size ranges, and the biotic composition of the cortices. In addition, nuclei types and the associated fauna are of importance. The growth form and shape are controlled by hydrodynamic factors, the biotic composition by light intensity and competition with herbivorous organisms. 1 Cyanoids, a group of skeletal oncoids formed by calcified porostromate cyanobacteria (see Pl. 53). Cores are platy phylloid calcareous algae (see Pl. 58). Note the irregular discoidal shape of the oncoids and breaks in growth of layers. The elliptical shape of the grains depends on the shape of the nuclei; the laminae form concentrically around the cores (Type C oncoid, Radwanski and Szulszewski 1966). Oncoid sizes in the oncolitic bed vary between about 5 mm and 1 cm and are within the range of piso- and macro-oncoids. Associated fossils are coated fusulinid foraminifera. SMF 13. Openmarine inner platform. Early Permian (Asselian): Carnic Alps, Austria/Italy. 2 Micrite oncoids produced by micritization of microbial oncoids, aggregate grains and peloids. Oncoid nuclei are shells and foraminifera. Note solution void filled with silt (arrow). Open-marine platform. Early Jurassic (Pantokrator Limestone): Peleponnesus, Greece. 3 Porostromate oncoids composed of Rothpletzella ( –> 3). These microfossils (attributed to green algae or to cyanobacteria) are common constitutents of Devonian open-marine and semi-restricted shelf limestones and reefs. They occur together with lamellar stromatoporoids and contribute to the formation of boundstones. Note the alternation of black (with tubular fossils) and white sparry laminae. Forereef slope. Late Devonian (Frasnian): Canning Basin, Australia. 4 Two-stage oncoid. The nucleus, a chaetetid coralline sponge is surrounded by an unlaminated meshwork corresponding to Bacinella (imterpreted as microbial in origin). Lamination is only developed at the periphery of the spherical oncoid. The grain is interlinked with other Bacinella oncoids, forming a framework texture together with corals and coralline sponges. High-energy subtidal environment of a carbonate platform. Late Jurassic (Plassen Limestone, Tithonian): Krahstein, Styria, Austria. 5 Modern soft lacustrine oncoid composed of cyanobacteria and green algae. The asymmetrical shape and the dome-like pattern of the laminae point to in-situ growth of the oncoid. The small micritic grains result from disintegration of the oncoids by waves and frosting during winter. The oncoids from Lake Constance are restricted to near-shore areas. There are two types, one smooth and hard, the other soft and exhibiting a spongy fabric. Soft oncoids occur embedded in a substrate of a sandy flat covered by 0.5 to 1.5 m water. Gnadensee, Lake Constance, southern Germany. 6 Sphaeroidal, laminar rhodoid formed by alternating crusts of corallinacean red algae (black) and encrusting foraminifera (white chambers). Note serpulid worms (S). The nucleus of the rhodoid is a poritid coral. Shallow subtidal back-reef environment. Early Tertiary: Lower Austria. 7 Brackish-water oncoids. The elliptical oncoids were formed by weakly calcified cyanobacteria and algae. Smaller particles are algal peloids and coated grains. Note discontinuous laminae and shrinkage structures within the oncoids (arrow). The recrystallized nucleus was initially coated by dense dark and light concentric laminae. The layers of the second growth stage follow the shape of the first growth stage. SMF 22. Coastal pond. Late Jurassic (Lusitanian): near Fatima, Portugal.
Plate 12: Cyanoids and Rhodoids
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Ooids: Terminology
4.2.5 Ooids Box 4.13. Selected case studies of ancient rhodoid limestones. A comprehensive list of references dealing with modern and ancient corallinacean red algae can be found in the Internet (http//paleoweb.net/algae).
As an introduction first let us show you some especially beautiful examples from the Early Triassic of Germany. Diameter of ooids: 0.4 mm.
Pleistocene: Hills and Jones 2000; Minoura and Nakamori 1982; Richter and Sedat 1983. Tertiary: Adey 1979; Adey and Macintyre 1973; Bassi 1995; Borrouilh-Le Jan and Hottinger 1988; Bosence and Pedley 1982; Boulanger and Poignant 1975; Braga and Martin 1988; Buchbinder 1977; Burgess and Anderson 1983; Dabrio et al. 1981; Dullo 1983; Esteban 1996; Fravega et al. 1993, 1998; Friebe 1988; Llombart 1982; Manker and Carter 1987; Martin et al. 1993; Maslov 1962; Moussavian and Höfling 1993; Nebelsick and Bassi 2000; Orszag-Sperber et al. 1977; Pedley 1996; Perrin 1992; Perrin et al. 1995; Pisera and Studencki 1989; Plaziat and Perrin 1992; Rasser 1994; Simone and Carannante 1985; Studencki 1988; Vanucci et al. 1993, 1994a, 1994b, 1996. Cretaceous: Reitner 1987. Carboniferous and Permian: Johnson 1946; Toomey 1975, 1983, 1985.
The environmental interpretation of rhodoid-rich limestones in the geological record must necessarily start with the appropriate choice of the present-day climatic model. The composition, growth forms and zonation of rhodoids and macroids are excellent facies indicators, but the risks of a too simple use of rhodoid data for paleoecological studies should be considered (Reid and Macintyre 1988). Rhodoids have been successfully used for estimating paleo-water depths (e.g. Perrin et al. 1995; Basso 1998), and evaluating sea-level fluctuations (Adey 1986; Bourrouilh-Le Jan and Hottinger 1988; Freiwald et al. 1991; Basso and Tomaselli 1994). In a geological section, variations in shape, size, algal communities and associated sedimentary features of rhodoids commonly reflect changes and temporal trends in depths and water energy (Braga and Martin 1988). The biotic composition of macroids reflect depth-controlled distributional patterns (Hottinger 1983; Piller and Pervesler 1989; Plaziat and Perrin 1992; Moussavian and Höfling 1993). Economic importance: Limestones with abundant rhodoids are popular decorative stones for exterior and interior use because of their pleasing appearance and resistance to abrasion. Rhodoid floatstones, formed in lagoonal and reef environments, are common in Oligocene and Miocene platform carbonates containing hydrocarbon rocks, for example in Indonesia, the Philippines and the South China Sea (e.g. Grötsch and Mercadier 1999).
Ooids have long been recognized as the most intriguing constituents of carbonate rocks (Sorby 1879; Kalkowsky 1908; Fig. 4.21). For a long time ooids were regarded as chiefly ‘inorganic’ grains, but strong biological controls on the formation of ooids are evident (Fig. 4.24). Microfabrics, the mineralogy, abundance and size of ooids reflect physical and chemical conditions of depositional environments in marine and nonmarine settings. Ooids, therefore, are valuable paleoenvironmental proxies for water energy, temperature, salinity and water depths. The paleoenvironmental interpretation of ancient ooids, however, should be done with some caution because (a) present-day marine Bahamian-type ooids formed in turbulent environments are not necessarily analogues of ancient ooids, and (b) the transport of ooids from the areas in which they were produced to the depositional areas is often underestimated. Changes in mineralogy and abundance of ooids over time have been related to changes in atmospheric carbon dioxide or to the rate of sea-level change and used as indicators of global secular variations in Phanerozoic climates. Valuable reviews on modern and ancient ooids were provided by Simone (1981) and Richter (1983). Oolites (calcareous rocks composed of ooids) formed on marine platforms and ramps comprise more than 50% of the world’s carbonate hydrocarbon reservoirs. Terminology Ooids are spherical and egg-shaped carbonate or non-carbonate coated grains exhibiting a nucleus sur-
Ooids: Shape and Size
rounded by an external cortex, the outer part of which is concentrically smoothly laminated. Most ooids are smaller than 2 mm in diameter; many ooids between 0.5 and about 1 mm in size. Modern calcareous ooids consist of aragonite and/or calcite. The bestknown noncarbonate ooids are iron ooids. The cortices of ooids exhibit a variety of microfabrics caused by different orientations of carbonate crystals in the laminae. Tangentially or randomly arranged crystals form tangential ooids (usually aragonite; Pl. 3/3, Pl. 120/1, Pl. 136/ 7), radially arranged crystals form radial-fibrous ooids (High-Mg calcite and/or aragonite; Pl. 13/2, Pl. 120/2, Pl. 130/3). Micrite ooids are spherical grains composed of a nucleus and a micritic cortex, in places exhibiting a vague concentric structure (Pl. 13/5, 6; Pl. 134/7). Spheroids are small grains composed predominantly of radially arranged fibrous crystals. Comparison with other grains Oncoids exhibit a more irregular, often non-concentric lamination; many oncoids are considerably larger than ooids. Pisoids are characterized by concentric laminae, but differ from ooids in their larger size (usually several millimeters) and often very irregular shape. Criteria Descriptive criteria providing clues to the environmental conditions and the depositional setting of ooid formation are the microfabric of the cortex, nuclei types, the shape and size of the ooid, associated fossils and sedimentary textures. Shape and size: Ooids occur as single or compound (complex) ooids (Pl. 136/7). The latter consist of two or more ooids bound together to form the nuclei for a new larger ooid (polyooid). ‘Multiple ooids’ result from the intergrowth of several ooids. Most ooids are spherical, but egg-shaped ooids are also common. The shape often results from the shape of the nuclei. Size and sorting provide clues to the hydrodynamic conditions. Ooid size is controlled by the supply of nuclei (Bathurst 1975), growth rate (Swett and Knoll 1989), mobilization and agitation (Carozzi 1989) and abrasion (Medwedeff and Wilkinson 1983). The maximum size of most ooids (about 1 mm) may reflect a balance between growth and abrasion. Experiments suggest that size differences between radial ooids and radial-concentric ooids may be explained by ooid transport processes (Davies et al. 1978). Heller et al. (1980) observed a critical size associated with a change in radial ooids to tangential ooids at 0.6 mm; this probably reflects a change in the mode of transport from suspension to bed load. Computer-based models demonstrate that a
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low supply of new nuclei and high cortex growth rate both promote the growth of large ooids. High average water velocity and velocity gradients also enhance ooid growth (Sumner and Grotzinger 1993). Conspicuously large ooids appear to be more abundant in Precambrian open-marine carbonates than in Phanerozoic carbonates (Pl. 16/8). Neoproterozoic giant ooids may have formed through a combination of lower nuclei supply (because of the lack of skeletal fragments during the Precambrian), higher growth rate, and higher average agitation level (which might be caused by a predominance of carbonate ramps over rimmed carbonate platforms during the Neoproterozoic; Sumner and Grotzinger 1993). A useful method for determinating the true grain size of ooids in thin sections was suggested by Fabricius (1967). Nucleus and cortex: Nuclei include lithoclasts, peloids, skeletal grains and minerals, e.g. quartz grains. Fecal pellets are common nuclei, particularly of ooids formed in low-energy environments (Jones and Goodbody 1984). Carozzi (1960) drew attention to the frequent association of ooids and rounded elliptical or spherical micritic grains representing rounded bio- or lithoclasts (Fig. 4.25). These grains (called ‘pseudooolites’) also occur as nuclei of ooids and indicate a long abrasion phase prior to overgrowth by ooid laminae. The cortex/nucleus ratio is an important descriptive criterion of ooids. Superficial ooids (Carozzi 1957) are ooids, in which the thickness of the cortex is distinctly less than half of the diameter of the entire ooids (Pl. 3/1, Pl. 29/6). ‘Normal’ ooids are characterized by a cortex whose thickness is equal or greater than half the diameter of the ooid. Superficial ooids often only reveal one or two laminae. Very thin oolitic films are a common feature of Bahamian ooids (Bathurst 1967). The average thickness of ooid laminae is 1–3 µm. Cortex thickness is related to environmental controls, particularly hydrodynamic factors. Small nuclei tend to have thick coatings, that probably form quickly, because smaller grains are more easily put into motion. Larger grains which are seldom moved have only superficial coatings or no coating at all (Bahamian ooids, Harris 1979). Mineralogy and microfabric (Fig. 4.20): The cortex of modern ooids consists of aragonite, Mg-calcite or calcite. Aragonite and Mg-calcite can co-occur in alternating laminae within one and the same ooid (Land
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Ooids: Structural Types
et al. 1979; Major et al. 1988). These mineralogies formed ooids in the Precambrian (Tucker 1984) and during much of the Phanerozoic. Cortical microfabrics: Most modern aragonitic ooids formed in high-energy marine environments exhibit a concentric or tangential microfabric, which is rarely well preserved in ancient concentric tangential ooids. Marine low-energy ooids are characterized by a radial-fibrous fabric which is also known from hypersaline environments and Low-Mg calcite ooids formed in freshwater environments. The original mineralogy of ancient ooids is inferred from preserved fabrics, owing to the fact that diagenetic calcification of aragonite typically results in obliteration of textural details, whereas ooids exhibiting radial arrangement of acicular crystals point to original calcite mineralogies (Kahle 1974). Radial ooids (Pl. 13/2, 4) are characterized by a cortical fabric consisting of radial-fibrous crystals. In most cortices this microfabric also exhibits concentric banding (radial-concentric, banded radial), but banding may be absent in small ooids with an average diameter of
Concentric (tangential) ooids
about 0.3 mm. The radial-fibrous structure of ancient calcitic ooids is considered to be primary because of the similarities to the microfabric of modern High-Mg calcite radial ooids from Baffin Bay, Texas (Land et al. 1979), the Amazon Shelf (Milliman and Barretto 1975), and from deep waters of the Great Barrier Reef (Marshall and Davies 1975). The significantly higher Mg content of radial calcite ooids as compared with tangential ooids may be suggestive of a High-Mg calcite precursor and a transformation of High-Mg calcite to Low-Mg calcite without destruction of microstructures. This primary radial microfabric must be distinguished from a secondary radial structure resulting from a diagenetic overprint on primary radial and concentric ooids (Marshall and Davies 1975). This structure transects the concentric laminae. The radially oriented large crystals may extend to the outer surface of the ooid or they may be limited to parts of the cortex. Traces of concentric layering can be preserved within the radially oriented crystals. Micrite ooids often occur together with other ooid types. The micritic fabric may be restricted to certain
Microfabric of the cortex
Mineralogy, modern examples
Environment
Concentric laminae consisting of tangentially arranged crystals whose long axes are aligned to the surface of the laminae. High microporosity
Aragonite: Bahamas, Yucatan, Abu Dhabi, Persian Gulf
Very shallow, warm lowlatitudinal seas; common in high-energy settings
(Great Salt Lake/Utah) Lacustrine-hypersaline Low-Mg calcite: Caliche ooids*
Radial (radial-fibrous) ooids
Laminae consisting of radially arranged crystals; long crystal axes perpendicular to the laminae surface
Terrestrial
Aragonite: Persian Gulf, Great Shallow marine, common in Barrier Reef, (Yucatan, low-energy settings Shark Bay, Mediterranean) Gulf of Aqaba
Sea-marginal hypersaline pool
Great Salt Lake/Utah
Lacustrine-hypersaline
Mg-calcite: (Baffin Bay/Texas)
Marine-hypersaline
Calcite and Low-Mg calcite: e.g. Cave pearls*
Non-marine
Micritic (random) Laminae composed of randomly Aragonite: Bahamas ooids arranged microcrystalline crystals or Laminae obliterated or absent, due to a pervasive micritization of the cortex
Shallow-marine
Fig. 4.20. Major structural ooid types. Concentric or tangential ooids originate commonly in high-energy settings; radial or radial-fibrous ooids are formed in moderate to low-energy environments. Well-preserved radial-fibrous fabrics of ancient ooids are generally regarded as primary features that have not been structurally changed over the course of transformation of High-Mg calcite to Low-Mg calcite. Micritic ooids can be generated by the random growth of crystals, or the micritic appearance of the ooids is caused by obliteration of original tangential or radial microfabrics by micritization and recrystallization of the cortex. Locality names put in brackets refer to places where ooids are rare; asterisks refer to terrestrial grains, which some authors call ooids, but which are treated as a separate grain category (pisoids, see Sect. 4.2.6) in this book.
Ooids: Structural Types
A
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B
Fig. 4.21. The drawings, taken from the fundamental paper by Sorby (1879), demonstrate two basic microfabrics found in ancient ooids and interpreted as being formed under different hydrodynamic conditions. A – Ferruginous tangential concentric ooids, which were repeatedly broken during formation. The nuclei are bioclasts (network structure: bryozoans) and fragments of ooids. The microfabric of the ooids is characterized by tangentially oriented crystals forming distinctly concentric laminae. The ooids occur together with small bioclasts and bivalve shells (top right) and angular terrigenous quartz grains. B – Single and compound radial-fibrous ooids formed around bioclasts, rounded clasts and peloids (black) as well as small oncoids (top left). The microfabric of the light-colored laminae is characterized by radially oriented fibrous crystals. Skeletal grains are echinoderms (top right) and mollusk shells. Note grain-support fabric and high open interparticle porosity. Middle Jurassic (A – Inferior Oolite, B – Great Oolite, Bathonian): Southern England. Primary sedimentary structures observed in the oolitic limestones of the Great Oolite Series suggest deposition in a tidal flat channel system (Klein 1965). The Great Oolite forms a hydrocarbon reservoir. The type of ooid fabric controls the distribution and amount of intragranular porosity (Sellwood and Beckett 1991).
laminae or it occurs everywhere in the cortex of originally calcitic or originally aragonitic ooids. Aragonite ooids can be transformed to micritic and microsparitic ooids by in-situ calcitization (Popp and Wilkinson 1983, Friedel 1995). In thin sections micrite ooids appear dark, featureless and microcrystalline. The micritic fabric results from the activity of microborers (Pl. 13/6) or from recrystallization, leading to a microgranular mosaic consisting of small calcite crystals (2–10 µm), and displaying neither concentric nor radial microfabrics. Fossils in ooids and oolitic limestones Oolitic limestones are either poor in fossils or they contain specific fossil assemblages. The composition of autochthonous fossil assemblages in grainstones is constrained primarily by sediment stability. A decrease in faunal diversity commonly correlates with increased sediment mobility (Feldman et al. 1993). The common absence or rarity of sessile or even mobile benthos in oolitic limestones suggests that periods of stabilization were not long enough for shelled organisms. Highly unstable and constantly shifting substrates are inhabited by only a few shelly benthos. Macrofossil assem-
blages can provide a powerful guide for estimating sediment stability when used in conjunction with the analysis of microfacies and sedimentary structures. Encrusting microfossils occurring upon ooids and enclosed within ooids indicate short-term stable conditions during the growth of the ooids (Fig. 4.22). Most common are encrustations by foraminifera (Pl. 13/3). Calcareous micro-algae in ooids are valuable environmental indicators (Dragastan and Richter 1999). Specific ooid types Fig. 4.23 summarizes a few ooid types characterized by specific morphology. Syndepositionally formed grains are cerebroid, eccentric, broken and regenerated as well as part of the distorted ooids. Some distorted ooids as well as half-moon and spiny ooids originate from diagenetic processes, the latter two by early diagenesis. Significantly deformed ooids are of tectonic origin. Modern ooids Modern ooids originate in marine and non-marine settings. Marine ooids occur in intertidal and shallow
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Modern Ooids
Fig. 4.22. Microfossils within and on ooids. A - Ooids offering a hard substrate for encrusting organisms. In areas dominated by mobile calcareous sands, ooids and bioclasts often are the only substrates available for encrusting organisms, particularly foraminifera. Note the multi-stage structure of the ooid. Ooid growth started on a nucleus formed by an ooid. The development of radial-fibrous laminae was disturbed by the encrustation of foraminifera, followed by a new set of laminae with radial-fibrous microfabric. The outer part of the ooid exhibits a tangential-concentric microfabric reflecting the change to more agitated water energy. The strong overgrowth by nubeculariid foraminifera indicates the absence of grain transport. Late Triassic (Carnian): Central Carinthia, Austria. B - Fossils as ooid nuclei. Due to the high mobility of ooid grains, oolitic limestones are commonly poor in autochthonous fossils. Exceptions are microfossils occurring as ooid nuclei, e.g. benthic foraminifera which provide clues to the age of oolites and their depositional environment. The nucleus of the radial ooid is an endothyranid foraminifera. Early Carboniferous (Kirchbach limestone, Viséan): Carnic Alps, Austria. Scale is 1 mm.
subtidal marine environments. Non-marine ooids are known from lacustrine and terrestrial sediments. Some of these ‘ooids’ are pisoids (Sect. 4.2.6). Ooids form in wave-agitated regions of the seas (high-energy ooids), and in marginal-marine or non-marine low-energy settings (‘quiet-water ooids’). Excellent examples of modern marine ooids (Box 4.14) include those from the Bahamas, the Persian Gulf, the Red Sea and the Great Barrier Reef. Very important studies on ooids formed in a hypersaline setting have been carried out in the Great Salt Lake in Utah, in coastal pools of the Gulf of Suez, and the southern Red Sea. The Quaternary ooid sand shoals of the Bahama Bank originate and accumulate in different subenvironments, including mobile fringes, sand flats, tidal channels, platform shelfs, open margins, and platform interiors. The shoals consist of long and narrow or spillover bars and tidal channels that trend nearly perpendicular to the length of the shoal. Tidal flow and wavegenerated current cause the ooid sands to be in motion over the entire width of the shoal. On average ooids range in diameter from about 0.25 to 1.00 mm and very rarely exceed 2 mm. The variations in the depositional settings are produced by a decrease in grain agitation from the ocean-facing margin of the interior of the ooid shoals. Local variations relate to topography and the stabilization of the sea bottom by vegetation. The main environments differ in the composition, abundance and
association of carbonate grains, the distribution patterns of organisms, as well as the type and frequency of physical and biogenic sedimentary structures. Ooids occur together with peloids, aggregate grains and skeletal grains. The distribution of ooid sand shoals is restricted to high-energy zones along open bank margins. Highenergy deposits are tidal-current or wind-wave-generated (Wanless and Tedesco 1993). Tidal-current generated sands are characterized by ooid growth in water depths ranging from less than 1 m to as much as 10 m. Well-developed tangential ooids originate in marine sand belts, tidal bar belts, floodtidal belts and in areas stabilized by algae and seagrass. Tide-dominated ooid sand bodies have been studied in the Bahamas and in the Persian Gulf (summarized in Halley et al. 1983). Examples of ancient oolites formed in tidally influenced zones are known from the Mississippian (Smosna and Koehler 1993) and Jurassic oolites of southern England (Klein 1965, Wilson 1968). Wind-wave-generated oolitic sands (southeastern Bahamas, Caicos platform) contrast strongly with tidalgenerated oolitic sands, restricted to shoals near the platform margins (northern Bahamas). Wind-wave generated sands are characterized by thick and widespread sheet-like ooid grainstones. Regular spherical, tangential ooids form stratified deposits on wind-wave agitated shallow subtidal shoreface settings on the platform interior, and on ocean-wave-agitated shorefaces facing the platform margin. Irregularly shaped ooids
Ooid Formation
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Box 4.14. Selected case studies of modern marine and non-marine ooids. Note that Mediterranean references also include studies of Early Holocene and reworked Pleistocene relict ooids. Marine environments: Bahamas – Ball 1967; Bathurst 1967, 1968; Carney and Boardman 1993; Cros 1979; Fabricius 1977; Gektidis 1997; Harris 1979, 1983; Illing 1954; Loreau 1982; Newell et al. 1960; Purdy 1963; Reitner et al. 1997; Strasser and Davaud 1986. Caribbean – Daetwyler and Kidwell 1959; Lloyd et al. 1987; Milliman 1969; Ward and Brady 1973; Zans 1958. Persian Gulf – Friedman 1995; Loreau 1982; Loreau and Purser 1973; Nazhat 1998. Gulf of Suez – Anwar et al. 1981, 1984; Sass et al. 1972; Sneh and Friedman 1984. Mediterranean Sea – Fabricius 1970; Richter 1976. Great Barrier Reef – Flood 1983; Marshall and Davies 1975. Shark Bay, Australia – Davies 1970.
which tend to deviate from their spherical shapes and are affected by biological degradation, cementation and encrustation, form banks and widespread subtidal sheets in areas of intermittent wind-wave-agitation across the platform interior. The ultrastructure of Bahamian ooids displays tangentially arranged, subhedral needle-like crystals or anhedral rod-like crystals. In contrast, ooids of the Great Salt Lake are characterized by crystals whose long axes are arranged perpendicularly to the ooid surface (Pl. 13/2). These radial-fibrous ultrastructures are common in marine and hypersaline ooids. Ancient records of radial-fibrous structures were initially regarded as diagenetic products of Bahamian-type aragonitic ooids (e.g. Shearman et al. 1970), but are now recognized as primary features. Modern depositional environments of marine aragonite ooids like the Bahamian ooids are said to be chemically not representative for ancient shallow-marine environments where calcite ooids have been formed (Sandberg 1975). Large and asymmetrical ooids, which are greatly different from Bahamian-type ooids, were described from the Laguna Madre in Mexico (Rusnak 1960). Laminae with radially and tangentially arranged crystals alternate in these ooids. The formation of these low-energy or quiet-water ooids (Freeman 1962) appears to be related to the low water turbulence and to periodic variations in salinity. Low-energy ooids formed in lagoons, coastal lakes and protected shallow environments with different salinity are common in the geological record. Low-energy and high-energy controls are not necessarily separated at a clear-cut level, as demonstrated by ooid sands occurring in lagoons along the Sinai beaches of the Gulf of Suez. The ooids occur in carbonate mud; they represent wind-blown material and
Hypersaline and freshwater environments: Saline lakes – Great Salt Lake, Utah: Eardley 1938; Halley 1977; Kahle 1974; Sandberg 1975; Reitner et al. 1997. Hypersaline lakes – Baffin Bay, Texas (marine-hypersaline): Land 1970. Laguna Madre, Texas: Freeman 1962; Rusnak 1960. Pyramid Lake, Nevada: Popp and Wilkinson 1983. Sea-marginal pools, Red Sea: Friedman 1978; Friedman et al. 1973; Friedman and Krumbein 1985. Dead Sea: Garber and Friedman 1983; Garber et al. 1981. Freshwater lakes – Davaud and Girardclos 2001; Wilkinson et al. 1980, 1984.
ooids formed under low and high energy conditions, differentiated by grain size and microfabric (Sass et al. 1972). Deep-water ooids: Modern and ancient ooids also are found in deep-marine settings. The bulk of these ooids represent allochthonous shallow-water material transported into slope and basinal settings as debris flows or turbidites. ‘Pelagic ooids’ (see Sect. 15.8) and ‘seamount ooids’ are found on the top of drowned platforms and seamounts (e.g. Hesse 1973). Controls on ooid formation The exact nature of the formation of individual ooids is still uncertain. It is generally accepted that most ooid deposits form in shallow waters which are regularly agitated over a long period of time by waves and/ or currents. Other environmental prerequisites include minimal siliciclastic input and generally warm temperatures, common in low-latitude settings. Carbonate ooids form in settings where at least five requirements must be met: • presence of nuclei, • bottom agitation to move grains, • a source of supersaturated water, • a process of water renewal, and • minimal amount of grain degradational processes. Major controls on ultra- and microstructures of ooid laminae are water chemistry, salinity and turbulence (Lippmann 1973; Loreau 1973; Fabricius 1977). The growth of radial-fibrous ooids in low-energy environments is apparently controlled by high or fluctuating salinity. This assumption is supported by experiments (Davies et al. 1978). A major and perhaps the essential
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Specific Ooid Types Cerebroid ooids (Graf and Lamar 1960; Carozzi 1962) Criteria: Indented periphery. Mottled appearance of the cortex. Sectors made of tangentially arranged laminae and radial micritic sectors which often start at former depressions of the nucleus surface. Occurrence: Marine and non-marine. Often associated with stromatolites. Interpretation: The indented periphery of the ooids has been explained as a result of a peripheral replacement. The formation of the micritic sectors is ascribed to bacterial dissolution and precipitation (Kahlke 1974). Asymmetrical and eccentric ooids (Gasiewicz 1984) Criteria: Small spherical and ovoid composite grains consisting of superficial ooids acting as eccentric nuclei for successively formed ooids whose growth takes place preferrentially upward. Alternation of very thin oolitic laminae (white) and micrite envelopes (black). Ooid size commonly < 0.30 mm. Rare in wackestones and grainstones, common in packstones, associated with abundant peloids. Eccentric ooids should not be confused with oomolds with eccentric nuclei. Occurrence: Low-energy shallow-marine lagoonal and lacustrine environments, periodically without agitation. Interpretation: The ooid laminae are explained as resulting from short periods of water agitation, the micritic envelopes correspond to longer non-turbulent periods. Eccentric growth is controlled by changes from suspension to saltation and traction processes at the sea floor. Top: A superficial ooid formed in suspension and deposited on the sea floor is the base of a micrite envelope and forms the eccentric nucleus of a subsequent new superficial ooid. Bottom: Ongoing ooid formation subsequent to rolling and turning. Broken and regenerated ooids (Carozzi 1961a) Criteria: Fragments of radial or tangential ooids that act as nuclei for new ooids are surrounded by ooid laminae. The ooids are characterized by alternating concentric ooid laminae and micritic laminae (white). Occurrence: Common in high- and low-energy settings, occurring together with partly cracked ooids and regular ooids. Interpretation: These ooids indicate multiple synsedimentary reworking and breaks in the formation of ooids, commonly associated with redeposition. Also known from ferruginous ooids (‘hiatus ooids‘, Berg 1944). Distorted ooids (Carozzi 1961b; Conley 1977) Criteria: Characterized by notched and stretched, sometimes flattened ooid grains, or grains connected by narrow apophyses, or series of grains linked in zigzag chaines parallel to bedding. The distortion has commonly preceded cementation and strong compaction. These ooids co-occur with other distorted grains (e.g. micritized intraclasts) in irregular pockets or in zones parallel to the bedding. Occurrence: Restricted to specific limestone horizons, affected by strong waves or currents. Common also in non-carbonate oolitic deposits. Interpretation: The shapes of distorted ooids illustrate a complete gradation from initial rupture and plastic deformation of soft grains to the rupture of rigid bodies. Explanations include fracture and plastic alteration of soft ooids in a turbulent environment, sediment sliding, or compaction. Deformation occurs when uncemented or poorly cemented ooids are buried (Pl. 13/8, Pl. 36/5). The term also is used for pitted ooids and cracked ooids which result from mechanical distortion and pressure solution. These ooids exhibit displaced and sometimes broken laminae (Pl.150/2). Half moon and shrunken ooids (Wherry 1916; Mazzullo 1977) Criteria: Ooids in which the interior cores have dropped to the bottom of the concentric outer layers, forming a geopetal fabric. Cross sections commonly exhibit a ’half-moon’ aspect characterized by an internal dividing line convex upward which separates an upper light part (frequently calcite-filled) and a lower dark part. Occurrence: Rare in carbonate-evaporite series, but also known from meteorically influenced carbonates. Interpretation: These ooids may be products of evaporite-carbonate or aragonite solution diagenetic processes (Carozzi 1963) and may indicate vanished evaporites in associated rocks (Folk and Pittman 1971; Folk and Siedlecka 1974). Geopetal ooids, however, can also result from the selective aggrading recrystallization of the ooid nuclei during a period of meteoric diagenesis (Mazzullo 1977). Spiny ooids (Davaud et al. 1990) Criteria: Characterized by external cortices exhibiting spines. The external cortices are detached from the underlying cortices near the points of contact between ooid grains, suggesting a postdepositional origin for the spines. The external cortices (white) are not stretched or broken and fit perfectly into the underlying cortices (densely arranged laminae) in undeformed areas. Occurrence: Lagoonal beach sands cemented in a vadose environment. Interpretation: The ooids are explained by tangential compressive deformation due to crystal growth in the outer cortices during early subaerial deformation. Spiny ooids differ from distorted ooids formed during burial compaction; these ooids display convex forms pointing into the pore space. Deformed ooids (Cloos 1947; Nissen 1964; Badoux 1970) Criteria: Originally spheroidal ooids are distinctly elongated, flattened and stretched. Microfabric structures recognizeable in traces only or completely destroyed. Occurrence: Described from folded carbonate series. Indicate tectonic style and timing of plastic deformations. Interpretation: Intensive tectonic stress results in elongations parallel to the schistosity planes.
Fig. 4.23. Common specific ooid types.
Description of Ooids
control on ooid growth is the distribution and composition of organic matter and the activity of microbes (Mitterer 1972; Suess and Futterer 1972; Reitner et al. 1997).
How to describe ooids? Box 4.15 summarizes the essential criteria for microfacies analyses of oolites. Diagenesis and porosity of ooids and oolites Early diagenetic changes of ooids include alterations by microborings and abrasion and processes related to the decay of organic matter. Dissolution processes occur during fabric-selective diagenesis which is controlled by the primary mineralogy. Aragonite within ooids deposited in shallow-marine shoal and shoreface
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environments is replaced by calcite. Dissolution starts with the dissolution within ooid laminae and continues to the formation of solution pores following the course of the laminae (Budd 1988). The loss of internal structures in many ancient ooids is attributed to a rapid solution of aragonite due to subaerial exposure, and facilitated by the decay of organic matter. Ooid cortices consisting of coarse blocky spar or small equant crystals are interpreted as a result of calcitization including the destruction of an originally porous aragonitic fabric and the formation of a coarse-grained calcite fabric (Rich 1982; Tucker 1984, 1985; Friedel 1995). Originally Mg-calcite radial-fibrous ooids transform to Low-Mg calcite via intergranular solution/precipitation films (Richter 1983; Heydari et al. 1993).
Box 4.15. How to describe ooids? Check your sample! Field criteria Depositional regime: Predominantly carbonate, mixed siliciclastic-carbonate or siliciclastic? Geometry of the oolite body: Mound-shaped or sheet-like? Restricted to well defined horizons or comprising a series of limestone beds? Limestone texture: Grainstone, packstone, wackestone? Or even bindstone? Sedimentary structures (lamination, bedding, cross-bedding, –> hydrodynamic conditions)? Association with lithoclasts and fossils? Laboratory criteria Abundance (high, medium, low, –> high abundance may indicate autochthonous ooid deposits) and distribution of ooids in the samples (concentrated and densely packed or scattered)? Single or polyphase ooids (–> redeposition)? Major structural ooid types (tangential, radial, micritic ooids; micritic ooids with microborings or signs of recrystallization? Radial ooids with fibrous crystals cross-cutting the laminae –> diagenetic alteration)? Specific ooid types (cerebroid; asymmetrical and eccentric –> changes in growth conditions; broken and regenerated –> synsedimentary reworking and temporal breaks in ooid formation, potential hiati in sedimentation; distorted –> disruption of normal ooid formation and/or diagenetic alterations; half-moon ooids –> clues to vanished evaporites; deformed ooids –> indication and measure of tectonic stress)? Dominating shape (spheroidal, ellipsoidal, other)? Size and sorting (ranges, average; uni- or bimodal; median, skewness –> differentiation of autochthonous and allochthonous ooids)? Nuclei (skeletal grain, lithoclast, peloids, mineral grain; single nucleus or several nuclei; preservation and rounding of nuclei; comparison of nuclei material with the non-ooid grains in the sample)? Nucleus/cortex ratio (–> clue to hydrodynamic conditions)? Cortex (homogeneous or inhomogeneous; thickness –> superficial or normal ooid)? Ooid laminae (thickness; alternation of light and dark laminae; signs of mechanical abrasion –> pointing to high-energy conditions; detached laminae –> compaction, burial diagenesis)? Inferred primary mineralogy based on the preservation (aragonite, calcite, bimineralic)? Fossils (acting as nuclei of the ooids, or encrusted on ooid laminae and surfaces; skeletal grains associated with the ooids)? Preservation of the ooids (partly or completely recrystallized, calcitized, dolomitized, silicified ooids; different preservation of laminae –> originally bimineralic composition; conspicuously black colored laminae or completely black ooids –> may indicate incorporation of organic matter under reduced conditions)? Cracks or ruptures within the ooids (–> shrinkage structures or burial features)? Oomolds (partly or completely leached ooids: abundance, distribution, internal sediment and cements; connecting pores).
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Ooids
Plate 13 Ooids: Key Criteria in Microfacies Studies Ooids are spherical particles with concentric laminae coating a nucleus. Modern marine ooids consist of aragonite or High-Mg calcite, or they are bi- or polymineralic. Calcareous ooids are classified according to (a) the primary microfabric (tangential, radial, micritic), (b) number of laminae (normal ooids with many laminae, –> 3; superficial ooids with few laminae, –> 7), and (c) shapes (–> 5, 8; compound ooids). Primary microfabrics are environmentally controlled: Tangential ooids originate predominantly in high-energy environments, radial ooids preferentially as ‘quiet-water ooids’ in low-energetic marine and non-marine environments (e.g. salt lakes, –> 2). Micrite ooids are caused by different processes (e.g. total micritization, –> 6; precipitation within biofilms, –> 5). Transport processes and diagenesis can change shape and composition of ooids (–> 1, 4 , 7, 8). Ooids are common constituents of shallow-marine carbonates, but also originate in non-marine settings, e.g. in hypersaline lakes and lagoons, and freshwater lakes. Wind-transported ooids form conspicuous eolianites. Genetic hypotheses focus on combined microbial-mechanical processes, chemical precipitation and chemical-physical processes. Current research favors strong controls by organic compounds (Morse and Mackenzie 1990). Some authors believe that microbes actively participate in the construction of ooid fabrics. Others suggest that microbes have a mainly destructive role or that they are only passively involved in building ooid layers by trapping particles or by providing a favorable substrate for nucleation of crystals. Ooids are of key importance in microfacies analysis: (a) They record paleoenvironmental conditions (water energy, transport, salinity, depositional sites), if details of microfabrics, shape and sorting are considered (Fig. 4.26; Pl. 120/1-2); (b) Ooids indicate changes in carbonate mineralogy and oceanographic conditions through the Phanerozoic; and (c) Oolitic limestones are important reservoir rocks for hydrocarbons, owing to high fabric-selective inter- and intragranular porosities, and contribute to the formation of ore deposits. 1 Tangentially structured, primarily Mg-calcitic ooids. The picture shows spherical ooids and ‘cerebroid’ ooids (arrow) with cones. The origin of cerebroid Mg-calcite ooids has been related to dissolution and precipitation by bacteria (Richter 1983). Cerebroid ooids are common in evaporitic settings (e.g. saline lakes), but also occur in marine environments. The photograph shows an original thin-section of Kalkowsky (1908) from the Rogenstein. Early Triassic (Lower Bunter): Bad Harzburg, Germany. The environment was a playa lake with variable salinity. Ooid-bearing beds are intercalated within mainly siliciclastic successions and occur in close association with laminated stromatolites. This association and the almost complete absence of fossils is characteristic for restricted brackish or saline environments. 2 Modern aragonitic radial and radial-concentric ooids. The small and partly distorted radial ooids have nuclei made of micritic clots. The cortex consists of radially arranged crystals. The larger radial-concentric ooids (top left) are subspherical in shape. Radial-fibrous crystals build the inner part of the ooid and occur within the concentrically arranged outer laminae. Nuclei are lithoclasts and fecal pellets of the brine shrimp. The onset of radial-concentric growth is sizecontrolled and indicates changes of low and high energy conditions. Great Salt Lake, Utah. This lake is regarded as a modern scenario for ancient salt lake microbialites and a place where bimodally sorted aragonitic radial ooids are being formed under hypersaline low-energy conditions. Ooid formation takes place close to the oxic/anoxic interface within the sediment, and appears strongly controlled by the composition of organic matter of biofilms attached to the ooids. 3 Radial-fibrous ooids periodically encrusted by sessile foraminifera. The biogenic encrustations took place during phases without movement of the ooids. Note the multi-stage development of the large ooid. Ongoing encrustation would result in the formation of an oncoid. Late Triassic (Carnian): Carinthia, Austria. 4 Radial-concentric quiet-water ooids and purely radial ooids (PRO). Note the radial-fibrous structure of the large ooids characteristic for formation in low-turbulence environments and the bimodal composition. Ooids exhibit moderate compaction (C) indicated by some interpenetrating grains. Breaks in ooid growth, regenerated broken ooids (left) and healed cracks (desiccation?), differences in shapes, and meniscus cements (arrows) indicate reworking in tidal zones (Simone 1974). SMF 15-R. Early Cretaceous (Neocomian): Monte Camposauro, southern Apennines, Italy. 5 Microbial ooids, formed in situ within a microbialite (see Fig. 4.24). Thin wavy peloid laminae alternate with ooids (arrows) and peloids. The small peloid nucleus is followed by a microspar rim and an outer micrite rim. The ill-defined laminated structure and shape of these ooids in combination with the occurrence of variously sized grains is a common feature of ooids formed in situ. Similar criteria are known from ooids of Late Jurassic sponge ‘reefs’. Cretaceous (Albian): Cantabria, Spain. 6 Micritized tangential ooids. The tangential fabric is only recognizable in peripheral parts of a few ooids. Note the very weak compaction. The depositional style reflects the situation 1 of the Carozzi model (Fig. 4.25), characterizing ooid formation and deposition in high energy zones. SMF 15-C. Oolite shoal, inner ramp environment. Early Carboniferous: Tournai, Belgium. 7 Weakly compacted grainstone with partly dissolved superficial ooids (‘oomolds’). Note the high intragranular porosity caused by dissolution of the originally aragonitic cortices and/or nuclei. Intergranular pores are filled with granular and drusy meteoric cements (see Pl. 29/3, 4). Pennsylvanian (Missourian): Utah, U.S.A. 8 Selectively dolomitized ooids. Note the distinct compaction of the grains indicated by the deformation of the ooids. The arrow points to distorted ooids (see Fig. 4.23). Cambrian: Yoho Valley, Rocky Mountains, British Columbia, Canada. –> 5: Neuweiler 1993
Plate 13: Ooids: Key Criteria
151
152
A
Compaction of Ooids
B
Fig. 4.24. Ooids originate in low- and high-energy environments (Fig. 4/19), but also grow in situ within microbial mats (benthic ooids), e.g. in hypersaline ponds (e.g. Friedman et al. 1985, Gerdes et al. 1994). The growth of the ooids is a result of cummulative processes by biofilm attachment and subsequent aragonite encrustation. A – Microbial ooid surrounded by sheats of filamentous mats. Dark laminae represent meshworks of degraded filament bundles of the cyanobacteria Microcoleus. Light laminae are extracellular polymeric substances produced by coccoid cyanobacteria and diatoms. B – Benthic ooid as a part of an aggregate grain. The ooid consists of microcrystalline aragonitic laminae with radially arranged crystals (light) and organically enriched laminae produced by biofilms (dark). Dark laminae are usually thinner than light laminae. In contrast to marine Bahamian ooids that are controlled by hydrodynamic conditions, the nuclei of benthic ooids are rarely allochthonous grains; they normally originate from the adjacent microbial community (e.g. cell colonies) or as metabolic products, including bubbles and liquid globules. Subrecent: Salina Janubio, Lanzarote, Canary Island, Spain. Scale is 100 µm. Courtesy of G. Gerdes, Wilhelmshaven.
Porosity: Aragonite ooids are susceptible to dissolution and formation of oomoldic porosity during freshwater diagenesis; moldic porosity is generally accompanied by low permeability. Primary calcite ooids are less susceptible to dissolution in meteoric waters, hence, preserved porosity is interparticle porosity, and permeability is generally higher. The understanding of the complex time and space relationships of different porosity types of oolitic limestones is critical for reconstructing the depositional and diagenetic history and for evaluating their importance as potential hydrocarbon reservoirs (Cussey and Friedman 1977; Carozzi 1981; Handford 1988; Sellwood and Beckett 1991). Primary porosity types of oolites include intra- and interparticle porosity. Secondary porosity is produced by early diagenetic intraparticle dissolution (oomoldic porosity) and late diagenetic fracturing. Oomolds (Friedman 1964): Dissolved aragonitic ooids are common in subaerially exposed Pleistocene limestones and are known from rocks as old as the Cambrian. Many oomolds exhibit an eccentric displacement of the non-aragonitic nuclei, often perpendicularly to the bottom of the mold, subsequent to the dissolution of the aragonitic laminae (Pl. 13/7, Pl. 29/3, 4, Pl. 150/ 1). The majority of observations indicates that oomoldic porosity suggests the presence of former aragonite. High-Mg calcite ooids transform to Low-Mg calcite
ooids with retention of fine fabric details and no indication of wholesale dissolution and creation of distinct moldic porosity. Compaction: Most oolitic limestones are somewhat compacted. The measurement of packing provides a tool for quantitatively estimating the degree of grain compaction before final cementation and lithification. Compaction measures derived from thin-section data of ooid grainstones include the packing density and the packing index (Coogan 1970). Both measurements consider the total number of grains and the number of grainto-grain contacts. Strong compaction of ooid carbonates is indicated by the frequency of interpenetrating grains, ooids with peripheral spalled-off laminae, flattened ooids (Pl. 36/5) and abundant, very closely packed ‘fitted’ ooids. Experimental compaction of ooids under deep-burial temperatures and pressure and varying salinity conditions shows appreciable reduction of bulk volume and porosity and the development of pressuresolution contacts. Plastically deformed ooids exhibit concav-convex and longitudinal contacts (Bhattacharyya and Friedman 1984). Dolomitization: Many limestones exhibit selective dolomitization of ooids starting with an initial precipitation along the periphery of the grains, followed by
Ancient Ooids
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Fig. 4.25. Model of the possible relationships between water energy, maximum and average grain size and the thickness of ooid cortices (superficial versus normal ooids). The model was developed for Great Salt Lake ooids (Carozzi 1957) and improved during the study of Mississippian shallow-marine ooid shoals in the Illinois Basin (Lacey and Carozzi 1967, Rao and Carozzi 1971). The maximum size of ooids and non-coated micritic grains is taken as a measure of the water energy level necessary to transport grains by traction or suspension processes. The numbers 1 to 6 describe different situations with regard to the hydrodynamic intensity which could favor the formation of ooids. Situation 1 and 2: The local agitation is greater than the maximum current intensity which brought the grains. Ooid formation affects all the grains available until the size of the largest ooid and that of the largest movable grain are (nearly) equal. The corresponding deposit is an oolite with a constant maximum size of normal and superficial ooids independant from the size of the available nuclei. Situation 3: The local water energy equals the maximum size of the non-coated grains. All smaller grains will be coated until their diameter reaches that of the largest non-coated grains. The sediment consists of normal and superficial ooids associated with equal-sized, rounded non-coated grains. Situation 4: The local water energy is lower than that of the currents which carry the largest grains into the environment. Grains which can not be lifted by the local agitation settle as non-coated grains. Other grains are coated, resulting in a sediment consisting of variously sized non-coated grains and small, predominantly normal ooids. Situation 5: The local water energy can just move the smallest grain resulting in very small superficial ooids, associated with non-coated rounded particles of different size. Situation 6: The local agitation equals or is less than the water energy needed to put the smallest grain in motion. The resulting deposit consists of rounded non-coated grains without ooids. Black circles: Non-coated carbonate grains (transported litho- and bioclasts, commonly rounded and micritic, brought into the area of ooid formation by currents); open circles: superficial ooids; circles with cross-hatched signature: normal ooids. After Carozzi (1960).
dolomite precipitation along porous cortical laminae (Zempolich and Bakker 1993; Pl. 40/8). The coalescence of dolomite crystals results in the formation of mimetic dolomite laminae and a concentric dolomite fabric (Pl. 40/6). Dolomite ooids without traces of concentric laminae may indicate dolomitization subsequent to a calcitization stage. Ancient ooids Non-carbonate ooids comprise iron ooids, phosphatic ooids (Swett and Growder 1982; Pl. 110/2, 4 , 6), siliceous ooids, and less frequent ooids consisting of silicate minerals (Meiburg and Kaever 1974; Ehrenberg 1993; Bloch et al. 2002), sulfates (Babel and Kaprzyk 1990) or sulfides (Schieber 1999). Ancient limestones with concentrically laminated iron ooids are of great economic importance. The iron minerals may be limonite, hematite, magnetite, siderite or chamosite. Iron oxides and chamosite are com-
mon in oolitic ironstones. Iron ooids in limestones are generally explained by the replacement of carbonate (aragonite?) ooids by iron and silica compounds, probably during very early diagenetic stages, perhaps preceding the inversion of aragonite to calcite. The onshore or offshore source of the iron (terrestrial weathering products, volcanic material, biochemical enrichment?) and the formation of oolitic ironstones is the subject of considerable debate (Gygi 1981, Kimberley 1983, 1994, Sturesson et al. 1999). Major problems are the process and environment of ooid formation (accretionary or concretionary; shallow-marine or nonmarine), the nature of iron minerals (diagenetic or primary), and the mechanisms of iron enrichment. Examples of modern iron ooids are known from lacustrine and marine environments (e.g. Lake Chad: Lemoalle and Dupont 1973; Mahengetang Island, Indonesia: Heikopp et al. 1996). Oolitic ironstones are known from the Precambrian and the Phanerozoic; they are particularly
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Ancient Ooids
Box 4.16. Selected case studies of ancient marine and non-marine (NM) ooids and ooid limestones. Pleistocene: Caputo 1993; Evans and Ginsburg 1987; Fabricius and Klingele 1970; Geno and Chafetz 1982 (NM); Hesse 1973; Jones and Goodbody 1984; Major et al. 1988; Richter and Neuser 1998; Wiedicke and Kudrass 1999 Tertiary: Roda 1965; Schoff 1937 (NM); Swirydczuk et al. 1979 (NM, low-salinity lake) Cretaceous: Gely and Loreau 1986; Nazhat 1998; Simone 1974, 1981; Strasser 1986; Tisljar 1983. Jurassic: Bausch 1961; Bersier 1937; Bosellini et al. 1981; Castellarin and Sartori 1973; Cussey and Friedman 1977; d’Argenio et al. 1975; Dickinson 1969; Dragastan and Richter 1999; Fabricius 1967; Heydari and Moore 1994; Koch et al. 1994; Kolckmann 1992; Loreau 1969, 1979; Martinis and Fontana 1968; Massari and Dieni 1983; Schaller and Koch 1996; Schloz 1972; Scudeler Bacelle 1976, 1980, 1983; Sellwood et al. 1989; Strohmenger et al. 1987; Swirydcuk 1988; Wilson 1968; Zempolich and Erba 1999. Triassic: Aigner 1984; Carozzi 1964 (NM); Fabricius 1967; Frantzen 1887; Friedel 1995; Gidon 1952; Haage 1970; Hornung 1998; Kalkowsky 1908; Labecki and Radwanski 1967; Milroy and Wright 2000 (nonmarine, freshwater lake); Paul 1982; 1985; Paul and Peryt 2000; Usdowski 1962, 1963; Young and Edmundson 1954; Zeng Yun-Fu 1983. Permian: Bausch and Wiontzek 1961; Clark 1980; Füchtbauer 1968; Herrmann 1956; Kerkmann 1966; Piatkowski 1977. Carboniferous: Burchette et al. 1990; Carozzi 1961; Choquette and Steinen 1985; Conley 1977; Dingle et al. 1993; Feldman et al. 1993; Henbest 1945; Hird and Tucker 1988; Hunter 1993; Keith and Zuppann 1993; Kettenbrinck and Manger 1971; Knewtson and Hubert 1967; Lacey and Carozzi 1967; Payton 1966; Wilkinson et al. 1984. Devonian: Poncet 1983; Reijers and Ten Haave 1983. Silurian: Richter 1983. Ordovician: Cantrell and Walker 1985; Carozzi and Textoris 1965; Davis 1966; Hyeong and Lee 1992; Schramm 1963; Stureson et al. 1999; Weaver 1992 Cambrian: Brown 1959; Carozzi 1963; Choquette 1955; Chow and James 1987; James and Klappa 1983; Mazzullo 1977; Poncet 1984; Sha and Jiang 1998; Swett 1965; Wherry 1916. Precambrian: Beukes 1983; Labecki and Radwanski 1967; Radwanski and Birkenmayer 1977; Simonson and Jarvis 1993; Singh 1987; Sumner and Grotzinger 1993; Swett and Knoll 1989; Tucker 1984, 1985; Zempolich et al. 1988.
abundant in two periods: from the Ordovician to the Devonian, and the Jurassic to the Early Tertiary (Young 1992). Iron oolites were formed in more than one mode. Microfacies studies of Devonian shallow-marine carbonates with iron ooids suggest that iron ooids originated in lagoonal or back-bar environments with temporarily phreatic conditions and fluctuating Eh-pH conditions. The iron was leached from pelitic sediments of the Old Red continent and was precipitated as concretionary iron hydroxide on bioclasts (Utescher 1992). The Middle and earliest Late Jurassic oolitic ironstones in central Europe originated in open, shallow to deeper marine environments (Gygi 1981). Iron was derived from the land by lateritic weathering. Iron ooids formed at the distal fringe of an argillaceous, terrigenous sediment below the normal wave base but within the reach of storm-generated waves. Ancient calcareous ooids: Ooids are known as far back as the early Precambrian (Beukes 1983; Simonson and Jarvis 1993: 2.5 billion years; ooids with radialconcentric structures). Ooids from the Early and the Late Precambrian (Radwanski and Birkenmajer 1977; Singh 1987; Swett and Knoll 1989; Tucker 1984, 1985;
Zempolich et al. 1988) exhibit distinct differences in size, shape and internal structures. This has been explained by fluctuation in the concentration of calcium carbonate in seawater during the Proterozoic (Grotzinger 1990). Estimate making according to the number of references (Box 4.16), oolitic carbonates were abundant in the Early Carboniferous and in the Middle and Late Jurassic, but seem to have been rare in Silurian and Tertiary deposits. Secular variations There is much evidence for temporal fluctuations in the mineralogy of marine carbonate precipitates during the Phanerozoic (Sect. 16.7.3). Sandberg (1983) postulated a secular variation in the mineralogical composition of ooids with aragonite (and High-Mg calcite) occurring in the Late Precambrian-Cambrian, Mid-Carboniferous through Triassic, and Tertiary to Recent, and Low-Mg calcite ooids in the Mid-Paleozoic and Jurassic-Cretaceous. The temporal variation was explained as a consequence of changes in the composition of ocean water and the major controls on the mineralogy (PCO and Mg/Ca ratio). Several authors prefer this model by Bates and Brand (1990). Many exceptions to the general pattern exist. 2
Ooids: Depositonal Sites
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Fig. 4.26. Depositional sites and environmental controls of ooids exemplified by a case study of shallow-water, near-coastal Purbeck facies (lowermost Cretaceous) carbonates of the Swiss and French Jura mountains (after Strasser 1986). Purbeck sediments comprise shallow subtidal, inter- and supratidal, normal marine, brackish, freshwater and hypersaline facies. Top: Inferred primary mineralogy, water energy, salinity, and depositional environment of ooid types characterized by different morphological features and cortical structures. All types show transitions. Type 1 and Type 4 may develop into superficial ooids, Type 2 to oncoids, and Type 3 to large asymmetric coated grains. 1 - spherical micritic ooids with thinly laminated tangential cortices forming well-sorted grainstones with up to 90 % ooids and only a few fossils (intertidal), 2 - irregular micritic ooids with thinly laminated cortices forming packstones (shallow marine to lagoonal), 3 - ooids with thinly laminated fine-radial cortices contributing to poorly sorted skeletal grainstones (deposition in quiet-water environments), 4 - ooids with a few fine-radial laminae occurring in well-sorted packstones and wackestones (brackish lagoonal environment), 5 - ooids with coarse-radial cortices contributing to well-sorted grain- and packstones (lagoons with varying water energy). Bottom: Environments where ooids may have been formed. Type 1 ooids formed in normal-marine shallow waters under high-energy conditions in small sand bars under the influence of tidal currents. Type 2 ooids originate in a quiet-water lagoonal environment populated by algae and cyanobacteria. The radial ooids of Type 3 indicate marine conditions differing in water chemistry from the environment of type 1 ooids, and frequent stirring-up of the sediment. Type 4 and type 5 ooids represent low-energy lagoons of varying salinity with charophycean algae and ostracods. Ooids were accumulated in sand bars by storms.
Wilkinson et al. (1985) found peak abundances of ooid formation during the Late Cambrian, Late Mississippian, Late Jurassic, and the Holocene. These periods coincide with periods of global transgressions and regressions. The reduction in ooid abundance in Middle Ordovician to Early Mississippian and Late Cretaceous to midTertiary might be explained by the absence of wave-
agitated environments during times of strong continental submergence or by a threshold formed by too high PCO levels. 2
Significance of ooids Two examples of the potential of ooid analysis in microfacies studies are shown in Fig. 4.25 and Fig. 4.26, and summarized in Box 4.17.
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Significance of Ooids
Box 4.17. Significance of ooids and oolites. Paleoenvironmental proxies Microfabrics, size distributions of ooids and sedimentary structures (e.g. cross-bedding, laminar bedding, fine laminae) of oolitic limestones reflect environmental conditions dominating water energy, salinity and water depths. Water energy levels and transport processes: Differences in high- and low-energy settings control growth and prevailing microfabrics of ooids, and determine the number of autochthonous ooids (deposited more or less at the place of origin) and allochthonous ooids (deposited after greater transport). As many ooids form in current-swept areas, they can accumulate in conditions different from those in which they were generated. Ooids of intertidal and shallow subtidal areas are frequently transported to adjacent subenvironments, but also to areas at a greater distance (e.g. slope, basins). Autochthonous ooids are characterized by rare or missing abrasion, the association of single and broken ooids, and the occurrence of ooids with same maximum size. These ooids exhibit a relatively simple transport history. Changes from radial to tangential concentric microfabrics combined with the crossing of a critical threshold of grain size allow the paleocurrents to be estimated (Heller et al. 1980). A diameter of 0.6 mm appears to mark the threshold for particles transported in suspension or as bed load (Miller et al. 1977). Allochthonous ooids are characterized by mixing of various-sized ooids (Pl. 46/6) and abrasion of laminae. These ooids are distinguished from ooids deposited in situ by microfabric, size, thickness of the cortices, and associated grains (Chow and James 1987). Evidence of redeposition is (1) the size, sorting, and skewness of ooids and associated grains, (2) the ratio of ooid content/mean diameter, and (3) the cortex/nucleus ratio (Carozzi 1983; Fig. 4.25). Low-energy conditions are generally inferred from the dominance of radial-fibrous microfabric, irregular and often asymmetrical shapes, and size parameters. Changes from low- to high-energy and from bed-load to suspension transport are reflected by microfabrics of ooids (Poncet 1984). Salinity: Freshwater ooids often differ from marine ooids in having exceedingly irregular surfaces and shapes, and commonly tangential or microsparitic laminae. Ooids formed in hypersaline settings appear to be characterized by the predominance of radial ooids as compared with other ooid types.Water depths: Some workers use the amount of ooids as measure of water depths (Fabricius 1967). Bahamian ooids are concentrated in very shallow waters. Their abundance in mobile ooid fringes and on sand flats (depth <1 m) ranges between 75 and >95% (Harris 1979). Sea-level fluctuations Sea-level fluctuations influence circulation patterns and submarine morphology, which in turn control the local conditions for ooid growth. This may be indicated by distinct compositional changes from skeletal to ooid grainstones or vice versa. Drops in sea level can lead to exposure and rapid cementation, causing the death of the ooid factory (Bosellini et al. 1981). Regional correlations Oolitic beds which can be laterally followed for long distances are valuable stratigraphic markers. An example are stratigraphically distinct Devonian ironstone horizons consisting of ferruginous ooids and bioclasts that were transported by storms several tens of kilometers over the open shelf (Dreesen 1989). Depositional settings Marine ooids are deposited in shoreface settings of ramps, on inner platforms, near outer margins of platforms, and as allochthonous sediments on the slope and in proximal or distal parts of basins. Specific settings can be recognized from the abundance, texture, ooid type, and the number of autochthonous and allochthonous ooids. Resedimented ooids, deposited as calciturbidites on the slope and in coalescent deep-sea fans, are often bimodal, poorly sorted and occur with lithoclasts and shallow-marine, platform-derived skeletal grains (Bosellini et al. 1981; Kolckmann 1992; Zempolich and Erba 1999). Paleoclimate and paleoceanography Ooids are often used as proxies for warm-water environments or tropical settings (e.g. Opdyke and Wilkinson 1990, Kiessling et al. 2002). The main assumption is that carbonate ooids favor warm waters with high salinities (Lees 1975). Oolitic deposits appear to be good paleoclimatic indicators. During the Phanerozoic they were usually concentrated in the tropics and lower subtropics, but not at or very near to the equator. Differences in the diagenetic style of ancient oolites are attributed to the prevailing climate. The paucity of early meteoric cements is explained as a result of an arid climate, whereas abundant meteoric cements should reflect a more humid climate (Hird and Tucker 1988). Integrated paleoceanographic and paleoclimatic models have been developed based on the mineralogical and geochemical criteria of ooids (Heydari and Moore 1993). Economic importance of ooid carbonates Hydrocarbon reservoir rocks: Because aragonite and calcite ooids behave differently with regard to dissolution and cementation, the original mineralogy as well as the spatial distribution of ooid sand bodies (sheets or shoals), and the geometry of ooid-bearing sequences have a significant influence on reservoir qualities of carbonate rocks (Swirydczuk 1988; Burchette et al. 1990; Hovorka et al. 1993). Oolites consisting of autochthonous and allochthonous ooids may represent different reservoir types due to significant differences of primary fabric-selective intra- and interparticle and secondary interparticle and fracture porosity (Carozzi 1983; Handford 1988). For important field case studies and a discussion of oolitic reservoirs see Keith and Zuppann (1993). Ore deposits: Carbonates with ferruginous ooids form important iron ore deposits in Europe and North America.
Pisoids and Vadoids
4.2.6 Pisoids and Vadoids – Simply ‘Larger Ooids’ or Carbonate Grains on their Own? Terminology Grains which are structurally similar to ooids, but have a diameter larger than 2 mm, are often called pisoids (from pisum, Latin for pea). Different names have been used for these grains, e.g. ‘diagenetic ooids’, caliche ooids (Fig. 4.27), cave pearls, vadose ooids etc. The size limit of 2 mm is highly artificial, but also considers, that most marine ooids do not grow to large sizes, because large grains can not be held in suspension long enough to be coated. Many pisoids are several centimeters in size, but others are smaller than the critical boundary allows. The term pisoid, has different meanings, depending on the genetic concepts inferred. Some authors use the name in order to describe rounded carbonate grains built by algae; these ‘algal pisolites’ however should be addressed as oncoids or rhodoids (Sect. 4.2.4). Other authors favor the term pisoid, when they want to express a non-marine origin for a coated carbonate or non-carbonate grain. It may be characterized by concentric layers made up of calcite or aragonite around a nucleus, commonly a lithoclast or pisoid fragments. The term vadoid introduced by Peryt (1983) underlines the substantiated or inferred formation of many (but not all) of these grains within a freshwater vadose or marine vadose environment, as indicated by specific cement types. Vadoid is a useful term if the vadose origin is sufficiently substantiated. A specific variety of pisoids, first recognized in back-reef strata of the Permian Reef Complex in New Mexico and West Texas, has been called walnutoid (Esteban and Pray 1983; Pl. 34/6, Pl. 126/1). The name is derived from the striking resemblance in size and shape to nuts of the Texas walnut tree. Comparison with other grains Pisoids are generally larger than most marine ooids, and characterized by dense or irregular laminations around a non-skeletal nucleus. Oncoids differ in the common presence of encrusting biota. Criteria Descriptive criteria of pisoids are the shape, nucleus, cortex, and structure of the cortex laminae as well as the size, sorting and texture of the grains. Shape: The shape of pisoids is environmentally controlled: Pisoids formed in continuously agitated water exhibit increasing roundness in the peripheral layers
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(e.g. cave pearls in high-energy splash pools, Pl. 14/8), or pisoids originating in geyser waters). Low-energy ‘quiet-water’ pisoids show irregular layers at the periphery and strong deviations from the spherical shape (e.g. pisoids formed in low-energy rimstone pools of caves). The shape may be asymmetrical, and exhibit downward or upward thickening (Pl. 14/6, Pl. 126/2) as well as lateral elongations whose laminae vanish within the surrounding sediment or connect single pisoids, indicating in-situ growth. Broken and recoated pisoids are common, indicating alternating periods of fracturing, reworking and pisoid growth. A common feature of vadoids is polygonal fitting resulting from vadose compaction. Nucleus and cortex: The nuclei are formed by fragments of pisoids and cement crusts, lithoclasts, peloids,
Fig. 4.27. Caliche pisoids. A - Holocene caliche with pisoids (center) and micritic crusts. Note the indistinct lamination around organically coated nuclei. The matrix contains abundant terrigenous quartz grains. Fuerteventura, Canary Islands. B - Ancient caliche with finely laminated layers which have been transformed into pisoids with indistinct boundaries. In contrast to stromatolites formed by biogenic activity, the laminar structures fill in the depressions of the microrelief. Such fabrics, together with in situ brecciation, desiccation cracks and the sparseness or absence of fossils, are good indicators of subaerial exposure of marine sequences and periods of nondeposition during which paleosols were formed in a semiarid climate. Middle Triassic (Wetterstein Formation, Ladinian): Reisskofel, Carinthia, Austria. Scale is 5 mm for A and B.
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Pisoids and Vadoids
Plate 14 Pisoids and Vadoids Pisoids are carbonate or non-carbonate grains structurally resembling ooids but differing in origin, environment, internal structure and often a larger size (several millimeters to centimeters). Discrimination between pisoids, oncoids and marine ooids is essential for recognizing non-marine or marginal-marine intervals, breaks in marine sedimentation (see Pl. 126) and the explanation of cyclic sequences (–> 1). Vadoids originate in a vadose freshwater (ranging in salinity from freshwater streams to hypersaline sabkhas) and in vadose-marine environments (e.g. supratidal areas, –> 4) . The grains are formed in various climate zones, in pedogene settings (caliche pisoids, –> 6; Pl. 128), karst regions and caves (spelean pisoids, ‘cave pearls’, –> 8), and comprise fluvial pisoids as well as marine-vadose pisoids (–> 1; see Pl. 126). Many carbonate vadoid grains are formed by chemical (e.g. cave pearls in speleothems), but also by microbially induced precipitation (e.g. caliche pisoids in humid subtropical and tropical areas). Phosphatic, ferrugineous, or siliceous pisoids are the result of the replacement of primary carbonate pisoids or concretional in-situ formations (–> 7). 1 Vadose-marine pisoids (‘walnutoids’, Esteban and Pray 1983), formed by accretionary growth on a shelf crest within shoaling upwards sequences in a protected hypersaline environment. The grains are linked by outer laminae and meniscus cement. Successive laminae differ in thinner darker micritic and clearer, microspar lamellae. Nuclei are peloids or fragments of preexisting pisoids. White areas represent neomorphic carbonate. Compare Pl. 126/1. Note that the existence of pisoids, inverse graded bedding, polygonal fitting as well as radial-fibrous fans and mammillary crusts need not necessarily indicate subaerial exposure. These pisolitic deposits are part of the shelf interior facies separating evaporitic dolomites on the landward side (lagoonal setting) and fossiliferous marine grain-, pack-, and wackestones on the seaward side (open shelf). Interpretations of the depositional environment ranges from caliche formation in continental or coastal-spray-zone settings (Dunham 1969; Scholle and Kinsman 1974), products of marine inorganic precipitation in inter- and subtidal settings (Esteban and Pray 1983), and grains formed in marine seepage or groundwater springs (Handford et al. 1984). The seep-spring hypothesis is currently favored. SMF 26. Late Permian (Carlsbad Group, Guadalupian): Mouth of Walnut Canyon, back-reef facies of the Permian Reef Complex, New Mexico, U.S.A. 2 ‘Coniatoids‘. The term has been suggested for carbonate crusts made up of uniformly thick coatings of aragonitic lamellae occurring in the supratidal zone of the Persian Gulf (Purser and Loreau 1973). Criteria: Very densely-spaced finely crystalline and fibrous laminae, sometimes forming discordant crusts on marine substrates. Note the irregular, non-spheroidal shape of the grains and the marine nuclei (aggregate grains). Supratidal environment. Early Carboniferous (Viséan): Czatkovice, Cracow area, southern Poland. Transgressive grainstone filling the relief of an underlying caliche horizon. 3 Broken and reworked pisoids indicating erosion and redeposition of vadoids in high energy environments. SMF 26. Early Cretaceous: Apennines, Italy, 4 Vadose pisoids. Note the meniscus cement (arrow) between the pisoids. Late Permian (Zechstein): Subsurface, northern Germany. 5 Coniatoids. As compared with marine ooids, these grains differ in their large non-skeletal nuclei (–> 2). Early Carboniferous (Viséan): Czatkovice, Cracow area, Poland. 6 Caliche pisoids formed on top of lacustrine stromatolites. Note the significant upward growth. These pisoids were formed within soils where vadose and shallow phreatic groundwaters became saturated with respect to calcium carbonate. Early Permian (Rotliegendes): Oberweiler Tiefenbach, Rheinland-Pfalz, Germany. 7 Siliceous pisoids. Bauxitic paleokarst surface. Note the shrinkage pores around the pisoids (arrows) and the morphological similarities with other vadoids (e.g. –> 1). Cretaceous: Subsurface Ras al Khaimah, United Arabian Emirates. 8 ‘Cave pearl’ (cave pisolite, spaeloid) of a modern speleothem. These non-attached calcite cave deposits originate by successive calcite precipitation from vadose-meteoric waters, forming strictly concentric lamellae around a nucleus; these vadoids are characterized by radial crystal structures. Note the slight deviations from spheroidal shape due to space competition with other grains. These grains grow in cupshaped depressions of caves. The grains are rotated by falling water drops (Homann 1969). Micritic pisoids with an irregular internal lamination have been interpreted as ‘subterranean oncoids’ because of calcite precipitation related to the metabolism of bacteria living within biofilms (Gradzinski 2001). Subrecent: Styria, Austria.
Plate 14: Pisoids and Vadoids
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or mineral grains. Nuclei formed by skeletal grains are rare, but may occur in peritidal settings. Pisoids may have more than one nucleus. Cortex: Many pisoids exhibit a distinct concentric structure consisting of very densely spaced dark and light laminae (Pl. 14/7, Pl. 126/1). The light laminae consist of microspar. Dark laminae are interpreted as microbially induced (Esteban and Pray 1983, Jones and MacDonald 1989). The number of laminae is variable and ranges from distinctly more than in ooids to only few. Caliche pisoids often have few laminae, freshwater and vadose pisoids many laminae (Esteban 1976), but many exceptions exist. Texture: Inverse grading is a common feature of autochthonous pisoids formed in situ. Common associations of pisoids are fenestral carbonates, tepees, diagenetic grainstones and vadose mammillary crusts, characterized by bent laminae at the bottom surface of pisoids. Diagenesis: Modern pisoids formed in hypersaline settings (e.g. Persian Gulf) consist of aragonite and/or High-Mg calcite. A brick-like microfabric of microsparitic laminae seen in many ancient pisoids is caused by meteoric-phreatic inversion of originally aragonitic laminae into granular calcitic laminae without a distinct dissolution phase (Assereto and Folk 1976). Experiments and studies of ancient peritidal carbonates suggest that intense fenestral alteration of original micritic sediments in the peritidal environment can lead to in situ formation of pisoids and diagenetic grainstones (Mazzullo and Birdwell 1989). Modern pisoids: Environments and controls Recent pisoids form under a wide range of environments including non-marine, transitional marine and shallow-marine settings (Box 4.18). Pisoids originate from Chemical precipitation from carbonate-saturated agitated water (e.g. cave pearls of speleothems; or thermal pisoids, e.g. Mammoth Hot Springs, Yellowstone). The concentric laminae originating from chemical precipitation are abraded by contacts with other grains. 2 Chemical and biochemical precipitation in low-energy environments (e.g. vadose pisoids, fluvial pisoids, pisoids formed in periodically desiccated and sometimes hypersaline tidal zones, and 3 concretionary growth in arid and semiarid pedogenic layers (caliche pisoids). 1
Environments of Pisoids
Box 4.18. Environments of pisoid formation. Non-marine • Pedogenic horizons (Burgess 1983; Chafetz and Butler 1980; Calvet and Julia 1983; caliche and soil pisoids see Sect. 2.4.1.1) • Speleothems (Kirchmayer 1964; Donahue 1969; Jones and MacDonald 1989; Tietz 1988; cave pearls see Sect. 2.4.1.3) • Cold- and hot-water springs and pools (travertine pisoids formed by inorganic precipitation and bacterially triggered carbonate deposition (Chavetz and Meredith 1983; Folk and Chavetz 1983) • Spring-fed pools (Schreiber 1981) • Freshwater and salt lakes (Jones and Wilkinson 1978; Handford et al. 1984; lacustrine pisoids see Sect. 2.4.1.7) • Fluvial environments (Braithwaite 1979; Jones and Wilkinson 1978; see 2.4.1.8) Transitional marine and shallow-marine • Beach zones • Coastal sabkhas and playa lakes (Jones and Renault 1994) • Supra- and intertidal zones (Purser and Loreau 1973, see 2.4.2.2) • Subtidal shelf crest
Several settings of pisoid formation are common in the geological record: Cave pisoids or ‘cave pearls’ are grains formed by chemical precipitation on a preexisting nucleus in a pool of water within a cave. They originate in high-energy splash pools and low-energy rimstone pools. The pisoids from splash pools are commonly spherical to subspherical (Pl. 14/8), exhibit a polished surface and typically have an exotic nucleus surrounded by distinct laminae exhibiting a radial microfabric, and yielding an extinction cross under crossed nicols. Pisoids from rimstone pools have an irregular shape with a rough surface, rarely have a nucleus, and commonly exhibit indistinct and irregular porous laminae. The crystal fabric of pisoids formed in speleothems is, at least partly, controlled by organic compounds and bacterial contribution (Jones and Renault 1994). Vadose pisoids may originate in hypersaline transitional environments or in meteoric environments. Characteristics of these grains are in situ growth, e.g. individual grains that coalesce and build composite grains (Pl. 14/1), and asymmetrical growth extending upwards (Pl. 14/6). Repeated desiccation of vadose pisoids can lead to fracture and grain breakage; reworked pisoid fragments are abundant in supra- and intertidal depos-
Ancient Pisoids
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its (Pl. 14/4). Features characteristic of vadose and phreatic pisoids are a radial fibrous fabric, perched inclusions, grain-to-grain contacts, polygonal fit of grains (Pl. 126/2), pisoids linked by cement crusts (Pl. 126/ 1), inverse grading (Pl. 126/2), tepee structures, and irregular non-spherical pisoid shapes.
cur? Pisoids have been interpreted variously as marine, schizohaline, lacustrine, fluvial, spring-surface fed surface pool, speleothemic, phreatic and vadose. Phreatic and vadose as well as schizohaline and marine records of pisoids are common from the Permian, Triassic and Jurassic (Box 4.19).
Vadose pisoids have many characteristics in common with caliche pisoids. Joint features are, for example, the downward elongation of laminae, and the development of fibrous or blocky calcite cements between the grains. Both hypersaline and caliche pisoids are associated with fenestral fabrics (commonly grainsupported for vadose pisoids, mud-supported for caliche pisoids). Features which can be used to differentiate vadose pisoids are (1) the regional geological setting of pisoid accumulation; (2) presence of fossils; (3) spar crystals with a triangular cross section, indicating precipitation from fresh water; and (4) fabrics which point to the original mineralogy (e.g. brick-like textures or elongated rays with square termination interpreted as aragonite relicts formed in hypersaline settings (Assereto and Folk 1976; Esteban 1976; Esteban and Pray 1983; Chafetz and Butler 1980; Hay and Wiggins 1980).
Primary marine pisoids are known from Paleozoic and a few Triassic and Jurassic occurrences. In contrast to the rather scarce record from the Phanerozoic, marine pisoids appear widespread in Proterozoic carbonates, commonly associated with ooids, stromatolites and intraformational conglomerates (Swett and Knoll 1989). This disparity is related to changes in the degree of calcium carbonate supersaturation caused by the Late Precambrian and Early Cambrian radiation of calcified algae and animals.
Caliche pisoids originate in situ at the surface or within surface layers by evaporation. The sizes ranges from about 2 mm to several centimeters. Many caliche pisoids are characterized by weak or obscure lamination, accretion and centrifugal growth, laminae of variable thickness, alternation of light and dark laminae, composite nuclei (pisoid fragments or lithoclasts), or absence of an observable nucleus, circumcracking structures, association with micritic crusts, a clotted peloidal texture and a mud-supported fabric (Fig. 4.27; Pl. 128/6). Fluvial pisoids and lacustrine pisoids typically show variable irregular shapes and sizes, alternation of finegrained detritic laminae and laminae consisting of coarse cement; association with laminated crusts, organic detritus and crystal sand. Unusual settings in which pisoids are formed are oilfield waters (Kemp 1959), landfills (Maliva et al. 2000), power stations (Knatz 1966), and old mines (Hahne et al. 1968). Ancient pisoids The interpretation of ancient pisoid limestones hinges on two questions: What processes caused pisoid formation and in what environment did formation oc-
The Permian Reef Complex case study: An instructive example of the varying interpretation of pisoid limestones in the ancient record are the pisoids occurring in bedded carbonates of the famous Permian Reef Complex (Pl. 14/1, Pl. 34/6, Pl. 126/1). Early workers considered the pisoids to have been formed by algae; they compared the grains with modern oncoids and assumed that the pisoids were formed and deposited in lagoons as free subaquatic sedimentary particles. Dunham (1969) and Thomas (1965) were the first to sugBox 4.19. Selected case studies of ancient pisoids. A comprehensive list of studies of modern and ancient pisoids and pisolites can be found in Peryt (1983). Pleistocene: Braithwaite 1983; Burgess 1983; James 1972. Tertiary: Gradzinski and Radomsi 1976; Misik 1980; Surdam and Stanley 1979; Swineford et al. 1958. Cretaceous: Enos 1974; Tisljar 1983. Jurassic: Melas and Friedman 1987; Goldberg and Friedman 1974. Triassic: Assereto and Folk 1976,1980; Bernoulli and Wagner 1971; Ciarapicca and Passeri 1983; Dozet 1991; Henrich 1984; Wagner 1987 Permian: Pisoids in the Permian Reef Complex, Texas – Dunham 1969; Esteban 1976; Esteban and Pray 1977, 1983; Loucks and Folk 1976; Kendall 1969; Mazzullo and Birdwell 1989 (discussion: Wright 1990); Pray and Esteban 1977; Peryt 1981, 1983, 1984; Clark 1980; Folk and Siedlecka 1974. Carboniferous: Gerhard 1985; Misik 1980; Wright 1981. Devonian: Wardlaw and Reinson 1971. Silurian: Kahle 1974; Frank et al. 1993. Precambrian: Cao Ruiji and Xu Yaosong 1983; Radwanski and Birkenmayer 1977; Simonson and Jarvis 1993; Swett and Knoll 1989.
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Significance of Pisoids
Box 4.20. Significance of pisoids and vadoids. Paleoenvironmental proxies Water energy levels: Differences in shape and lamination of vadoids (e.g. cave pearls) allow the evaluation of continuous or discontinuous and high- and low-energy environments. Transport processes: The analysis of the association of vadose pisoids with other coated grains allows transport processes within schizohaline peritidal zones to be recognized (Peryt and Piatkowski 1977). Salinity: Many pisoids found in peritidal carbonates were formed in waters of increased salinity. Other pisoids are indicative of non-marine salinity, e.g. freshwater or brackish waters. Sea-level fluctuations Pisolitic limestone beds intercalated within marine sequences may indicate breaks in sedimentation caused by drops in sea level, and associated with karstification phases. Pisoids are common constituents of the supratidal and intertidal members of shallowing-upward cycles. Depositional settings Caliche pisoids, cave pisoids, lacustrine pisoids and meteoric-vadose and marine-vadose pisoids can be used to recognize different subenvironments of terrestrial and marginal marine settings (Esteban and Pray 1983). Standard microfacies 25 underlines the importance of pisolitic carbonate rocks. Paleoclimate Micro- and ultrastructure of cave pisoids as well as caliche pisoids reflect seasonal changes and longer-term climatic changes. Economic importance Hydrocarbon reservoirs: Pisoid limestones are important reservoir rocks, e.g. in the Permian of West Texas and in the Mississippian of North Dakota (Gerhard 1985). The productive facies include pisolitic wackestones and grain- to packstones. Porosity is enhanced by subaerial and vadose dissolution related to pisoid formation.
gest a completely different interpretation of the pisoids. Pisoids were regarded as vadose concretions produced by subaerial vadose exposure and diagenesis in a caliche soil. The caliche hypothesis gained wide acceptance, but started being discussed again. Texas pisoids were compared with modern pisoids found in sabkha environments of the Persian Gulf (Scholle and Kinsman 1974). Assereto and Folk (1976) and Loucks and Folk (1976) concluded that the pisoids were produced by marginal marine supratidal environments. Intensive research by Lloyd Pray and his group led to a complete rejection of the caliche and the vadose origin hypotheses (Esteban and Pray 1976, 1983). The authors interpreted the pisoid depositional setting as a shelf crest of a carbonate marginal mound separating an evaporitic lagoon from a more open-marine area. Pisoids and pisolites formed from the intensive precipitation from very shallow hypersaline waters in a peritidal environment. Thick pisolite development commonly occurred as parts of shoaling upward sequences of inter-tepee depressions. As shown by the pisoids of the Permian Reef Complex, differences in shape, nuclei, lamination and texture of pisoids reflect differences in the depositional environment (Noe 1996). Mobile transported pisoids, formed on shallow subtidal environments, are characterized by spherical and ellipsoidal shapes, symmetrical cortices, nuclei of bioclasts and broken pisoids, and a normally graded texture. Pisoids formed in situ within
peritidal settings exhibit a polygonal fitted fabric, asymmetrical cortices, downward bent laminae, sediment inclusions in the peripheral laminae, and inverse grading. Similar criteria occur with pisoids formed in situ in the vadose zone (e.g. caliche pisoids), but these grains are distinguished by their association with meniscus and dripstone calcite cements and mammillary cement crusts. Non-carbonate pisoids Most modern pisoids consist of calcite or aragonite, but ferruginous, phosphatic and siliceous pisoids (Pl. 126/3) are also known (Carozzi 1960; Jenkyns 1970; Bhattacharyya and Kakimoto 1982; Burgess 1983; Soudry and Southgate 1989; Abreu 1990). Non-carbonate pisoids were formed in more than one mode (Sect. 4.2.5). The source of the iron of ferruginous pisoids has been related to volcanism, as well as to solution and weathering processes. Pisoids occurring within laterite and bauxite profiles are explained by dissolution and reprecipitation controlled by climatic variation, chemistry of the groundwater in terms of its pH and Eh, and differences in weathering of the parent rocks and in transport processes (Pl. 126/4). Significance of pisoids Box 4.20 summarizes the significance of pisoids in microfacies analysis.
Aggregate Grains
4.2.7 Aggregate Grains: Grapestones, Lumps and Other Composite Grains Originally separated ooids, bioclasts and other grains can be bound together by organic films, encrusting organisms and aragonite or Mg-calcite carbonate cements, forming composite grains. The best known modern examples are the sand-sized ‘grapestones’ of the Bahama Banks which constitute up to eighty volume percent of the total number of particles in a very large transitional area from agitated ooid shoals to the protected mud and peloid mud environments. Strong micritization of the component particles often makes it difficult to distinguish between aggregate grains and resediments, which are reworked chunks of partly consolidated sediment. Most aggregate grains form in very shallow areas where currents and waves are sufficient to remove fine-grained sediment, but not sand. Recent Bahamian-type aggregate grains occur in subtidal and intertidal shallow-marine environments with restricted or changing circulation. Grapestones, consisting of peloids, ooids and oncoids, and cemented by aragonite or cryptocrystalline Mg-calcite are also known from sea-marginal hypersaline pools (e.g. Red Sea: Friedman 1978). In hypersaline coastal ponds of the Black Sea, cementation within grain aggregates occurs by means of bacterial degradation of calcium sulphate (Baltres 1975). Grapestones and other varieties of aggregate grains are characteristic constituents of many Precambrian and Phanerozoic platform carbonates and valuable paleoenvironmental proxies. Terminology Aggregate grains were first observed in the Bahama Banks (Illing 1954; Purdy 1963), then in the Persian
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Gulf (Taylor and Illing 1969, Kendall and Skipwith 1969) and also in the Gulf of Batabano south of Cuba (Bathurst 1975). The grains are common in platform environments, but can also be associated with reefs (Milliman 1967). Different names characterize different shapes and the variable contribution of organisms and carbonate cements to the formation of these composite grains: Grapestones: Aggregates of spherical grains (often micritized ooids) whose external shape resembles microscopic clusters of grapes. Typical grapestones exhibit very little microcrystalline carbonate, but often interparticle carbonate cements (Pl. 15/2). Micritization of these cements may transform a grapestone grain into a lump grain (Pl. 62/1). Lumps: Aggregates with a smoother outline than grapestones; they commonly have hollow interiors and are often strongly micritized (Pl. 15/3). Botryoidal or oolitic lumps: Grapestone or lumps with a thin oolitic coating. Microbial/algal aggregate grains: Sedimentary particles are agglutinated by biofilms, cyanobacteria and algae (Pl. 15/5). The grains are common in lagoonal environments and often associated with microbial mats. Encrusted aggregate grains: These grains are characterized by strong biogenic encrustations with sessile foraminifera, calcareous algae, serpulids and other organisms (Pl. 15/4). The grains are common in protected lagoonal environments. A typical feature of aggregate grains is the irregular shape and the typically lobate outline (Pl. 15/1, 8). The size ranges between 0.5 and several millimeters, but lies commonly below 1 mm. Although many transitions exist between these varieties (Halley et al. 1983), the more common types are
Fig. 4.28. Multi-stage formation of aggregate grains (grapestones and lumps). Stage 1: Ooids and skeletal grains are bound together by microbial filaments, organic biofilms and encrusting organisms, e.g. foraminifera. Stage 2: Increasing calcification of microbiota contributes to the formation of a porous, cemented aggregate (grapestone) with a lobate outline. The grains are affected by microborers leading to micritization. Stage 3: Increased cementation (black) creates a smoother relief and reduces the intergranular porosity. Stage 4: Continuous micritization and cementation form a matrix-rich aggregate grain whose outline becomes successively more rounded. After Tucker (1990).
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Aggregate Grains
meaningfully differentiated because different types correspond to specific stages in the formation of aggregate grains as well as different energy levels.
grains. Coatings of aggregate grains by microbes and encrusters may result in the formation of cortoids and even oncoids.
Comparisons Grapestone type aggregate grains are commonly easily recognizable by shape and composition. If the grains have been transported and eroded along their surfaces, differentiating between aggregate grains and rounded intraclasts becomes more difficult. Rounded intraclasts occurring together with angular intraclasts are often larger than typical aggregate grains and may exhibit inclusions of microfossils. Extraclasts consisting of pieces of grainstone or wackestone differ from aggregate grains by truncation of the composite particles at the extraclast boundary and often a strikingly different composition as compared with associated
Controls on the formation of modern aggregate grains The Bahamian grapestone facies forms within relatively low-energy and low-nutrient environments with low sedimentation rates. Extensive grapestone deposits made of skeletal and ooid grains occur in very shallow water of the Great Bahama Bank margin at the bankward edge of flood tidal deltas and along the beaches of the leeward side of islands and keys beyond the surf zone. Aggregate grains tend to form in areas with some water circulation, winnowing of fines and mobilization of sediment, followed by periods of stabilization and cementation (Wanless 1981). They are absent in areas of dense vegetation.
Plate 15 Aggregate Grains (Grapestones, Lumps): Common on Shallow-Marine Carbonate Platforms Aggregate grains, first studied in the Bahama Banks, are composite grains consisting of ooids, skeletal material and peloids, bound together by organic films and aragonite or Mg-calcite carbonate cements. In thin sections many aggregate grains exhibit a typically lobate outline (–> 1, 4, 8) and are encrusted by foraminifera, calcareous algae or serpulids. Varieties of aggregate grains include ‘grapestones’ (–> 2), ‘lumps’, –> 3), and microbial/ algal aggregate grains, common in lagoonal environments (–> 4, 8). Common prerequisites for the formation of aggregate grains are stabilization of the individual particles during reduced phases of deposition in environments with restricted circulation, cementation on the sea bottom and periodical reworking. Biofilms, microbially induced early cementation between the particles, and the absence of mud are important controls. –> Note that limestone classifications do include aggregate grains in the intraclast or lithoclast categories! 1 Modern aggregate grains (‘grapestones’) consisting of densely packed, micritized ooids, peloids and foraminifera (arrow). The grains are bound together by organic films and thin cement rims. White areas within the large lobate aggregate grain are open pores. Berry Islands, Great Bahama Bank. The grapestone belt on the Great Bahama Bank covers a very large shallow area, deeper than 3 m, with a maximum depth of 10–15 m. 2 Pleistocene aggregate grains forming weakly lithified calcareous sands. Note thin cements on the surface of some grains (arrows). Most intergranular pores of the sediment are still open. E: Echinoid fragment. San Salvador Island, Bahamas. 3 Pleistocene aggregate grains (‘lumps’ because of the large amount of micrite in the grains). The grains are somewhat rounded, but the characteristic lobate outline is still visible. Aggregates consist of ooids and skeletal grains (e.g. foraminifera,). The calcitic bivalve shell (P) exhibits a layered prismatic microstructure. Late Pleistocene: Grotto Beach, San Salvador Island, Bahamas. 4 Encrusted aggregate grains typically showing very irregular outlines, distinctive invaginations and protuberances, and encrustations of miliolinid foraminifera (spar-filled chambers appearing as white bent elements, arrows). Intergranular pores are filled with submarine fibrous cement and vadose crystal silt. D: Dasyclad green alga. Carbonate Platform. SMF 17. Middle Permian: Velebit, Croatia. 5 Microbial lumps and small micritized dasyclad green algae (D). The net-like structures within the aggregate grains are cyanobacteria. Inner carbonate platform. SMF 17. Middle Triassic (Ladinian): Marmolada, Dolomite Mountains, Italy. 6 Aggregate grains occurring together with loose peloids. The lumps consist of peloids and a few skeletal grains. Outer carbonate platform. Early Jurassic: Djebel Assiz southwest of Tunis, northern Tunisia. 7 Aggregate-grain limestone. Carbonate platform. The sample was taken near a reef as shown by the abundance of reef biota within the aggregate grains. Note rounding and different size of many grains. Late Triassic (Bedded Dachstein Limestone, Norian): Gosaukamm, Austria. 8 Aggregate-grain grainstone. The grapestones consist of ooids and fossils. Some aggregate grains exhibit thin oolitic coatings (arrows,‘botryoidal lumps’). SMF 17. Middle Triassic (Anisian): Lagoonal limestone. Olang Dolomites, Italy.
Plate 15: Aggregate Grains
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Bahamian aggregate grains (Pl. 15/1, 3) demonstrate (a) a strong control of grapestone formation by microorganisms living in the interstices between the composite grains and in microborings (Fabricius 1977), and (b) continuous successions of increasing binding, agglutination and induration. Cyanobacteria, algae and fungi are effective in connecting and binding the constituent grains, trapping loose silt material (Cros 1979), destructively alterating the composite grains, and triggering cementation within intergranular voids. The formation of grapestones requires a sequence of depositional conditions, starting with the supply of the constituent grains (Fig. 4.28). After an environmental change, a subsequent set of depositional conditions must allow binding and cementation of grains. These changes correspond to periods of bottom mobility alternating with periods of bottom stability. Grapestone sites have been regarded as areas of nondeposition or at least reduced deposition. Marine cementation is strong because the process has millennia to operate on a limited amount of sedimentary material (Winland and Matthews 1974). Not all ancient limestones with abundant aggregate grains were formed by the processes outlined above: Constituents of the well-sorted grapestone sands of the Bahamas, may also correspond to sand and granulesized lithoclast fragments derived by erosion of firmground crusts (Demicco and Hardie 1994). Firmgrounds are sediments in the initial stages of hardground formation that can be easily broken up and transported as cohesive intraclasts (Jindrich 1969; Hardie and Garrett 1977; Dravis 1979). Carbonate aggregates, similar in morphology to ‘grapestones’ are common in lower slope fore-reef sediments, off Barbados. They were formed in situ by submarine lithification processes (Stentoft 1994). Ancient aggregate grains Carbonate grains which are similar to modern shallow-marine aggregate grains with respect to shape, size, composition and internal cements are already known from Precambrian limestones. Phanerozoic records of limestones with abundant and dominating aggregate grains are common in platform carbonates, but appear to be concentrated in specific time intervals including the Devonian, the Late Permian (Pl. 15/4, Pl. 34/7), the Middle and Late Triassic (Pl. 15/5, 7, 8), Jurassic (Pl. 15/6, Pl. 62/1) and the Cretaceous (Pl. 29/9, Pl. 43/1, Pl. 127/4). Limestones with abundant aggregate grains, peloids and only a few special skeletal grains are common members of platform carbonates formed in openmarine settings. Two common paleogeographic settings of larger grapestone depositional areas were the areas
Intra-, Extra-, and Lithoclasts
behind and adjacent to shelf-margin reefs (e.g. Pl. 121/ 4), and areas which are transitional on the seaward side to ooid shoals and green algal-foraminiferal sands on the landward side (Pl. 61/1). Significance of aggregate grains Aggregate grains, which can be interpreted as analogs of Bahamian grapestones, indicate low to moderate and changing water energy levels, tropical and subtropical warm-water conditions, low-nutrient environments and low sedimentation rates. The depositional setting of bedded limestones composed predominantly of aggregate grains usually corresponds to attached or isolated platforms. Grainstones predominantly consisting of aggregate grains are characteristic parts of platforms formed during sea level highstand phases. Economic potential: Carbonates with aggregate grains occurring in association with oolitic limestones are reservoir rocks (e.g. in the Cretaceous of the Near East). Because most grapestones originate in environments with no or very low terrigenous input, the noncarbonate residues of many limestones with abundant aggregate grains are conspicuously low. These limestones are important resources of chemically high-quality carbonate rocks (Flügel and Haditsch 1977).
4.2.8 Resediments: Intra-, Extra- and Lithoclasts – Insiders and Foreigners The redeposition of material derived from preexisting carbonate sediments and rocks is a common process in marine and non-marine environments. Reworking and redeposition can take place early and contemporaneous with sedimentation, or late and post-depositional. Syn- and postsedimentary reworking and erosion produce carbonate clasts. These clasts reflect environmental conditions within the depositional basins as well as the input of material brought into the basins. Constituent clast analysis of extrabasinal material has become a major tool of microfacies studies providing an excellent basis for reconstructing ‘lost carbonate platforms’ which are recorded only by their relics preserved in small clasts in limestones or abundant clasts of breccias (Sect. 16.3.2.1). Terminology The two major types of millimeter- to centimetersized redeposited grains are intraclasts and extraclasts (Folk 1959). An intraclast is a carbonate fragment of lithified or partly lithified sediment, derived from the erosion of nearby penecontemporaneous sediment from
Intraclasts
within the basin and redeposited within the same area (Pl. 16/7, 8). An extraclast is a fragment of carbonate rock derived from the erosion of an exposed ancient limestone on land outside the depositional basin in which it is found (Pl. 16/2, 4). These definitions look succinct, but are in practice somewhat tricky because of the difficulties involved in the precise recognition of the two categories in thin sections. Many authors, therefore, use the term lithoclast (Folk 1962) to characterize both, erosional detritus brought into the depositional area from an outside source (= extraclasts) as well as pieces of semilithified sediment which have been reworked syndepositionally (= intraclasts). The term limeclast (Wolf and Conolly 1964) which designates ‘limestone rock fragments and grains’ has the same comprehensive meaning. The term was introduced to replace the genetic term ‘intraclast’. Genetically, limeclasts embrace however, ‘intraclasts’ as well ‘extraclasts’. It should be noted, that the term lithoclast referred in the original definition only to angular fragments of reworked non-contemporaneous particles of extrabasinal lithified carbonate sediment. Other terms that describe genetic subcategories of intraclasts are plasticlasts, algal intraclasts, pseudo-intraclasts, and protointraclasts. The last term as well as caliche intraclasts refer to diagenetic products. Plasticlasts (Folk 1959) result from the reworking of weakly consolidated lime mud at the sea or lake bottom by waves or currents; the grains correspond to small mud flakes and mud pebbles (Pl. 16/7). These micritic particles are usually small (< 0.5 mm); they have the same composition as that of the micritic matrix. Algal intraclasts resulting from reworking of lime mud covered by microbial and algal mats are common constituents of inner ramp and platform carbonates. Pseudo-intraclasts (Wobber 1965) are produced by bioturbation. The breakup of the sediment by burrowing organisms causes variously sized, often angular micritic clasts, which are more or less still in place. Protointraclasts (Bosellini 1964) or autoclasts (Sander 1936) can be formed by early diagenetic movement of sediment related to submarine or subaerial dehydration of the sediment, but also to seismic shocks. The common characteristic is that the boundaries of the separated clasts still fit like pieces of a jigsaw puzzle. Diagenetically formed intraclasts which are still in place also occur in calcareous paleosols. These caliche intraclasts (Siesser 1973) are subaerially broken pieces of caliche crusts (Esteban and Klappa 1983; Francis 1986; Pl. 128/6). Calclithite is a terrigenous or carbonate-cemented rock consisting of more than 50% carbonate extraclasts,
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which are derived from the erosion of older lithified limestones (Folk 1959). The grain size of these rocks ranges from silt to gravel, but is commonly formed in the sand or conglomerate sizes. Many calclithites are typically stream or alluvial fan deposits. Calclithites record rapid erosion and signify a high-relief source area, as shown by the high content of carbonate extraclasts in the lithic sandstones of the Alpine Molasse basins.
4.2.8.1 Intraclasts: Origin and FaciesDiagnostic Types Intraclasts are commonly found in shallow-marine environments, but also are transported to deepwater. Shallow-marine environments in which intraclasts are formed are characterized by wave-dominated regimes and tides that continuously rework carbonate. ‘Rip-up clasts’ are common in inter- and supratidal settings but also occur in subtidal environments where scouring undermines lithified mud beds (Exuma Sound, Bahamas). Intraclast grainstones are often interpreted as deposits formed by storm wave erosion and reworking of various sediment types occurring in shallow-marine environments. The occurrence in cross-bedded channel fills supports this interpretation. Intraclasts can also originate by desiccation of carbonate muds in supratidal environments (Ainardi and Champetier 1976). Intraclasts formed by seismic events are discussed in Sect. 12.4. Box 4.21. Selected records of ancient and modern ‘black pebbles’. ‘Black pebbles’ are known from ancient paleosols, karstic sinkholes and reworked speleothems, lacustrine ponds, beachrocks as well as from tidal and subtidal carbonates. A model for blackening and blackpebble sedimentation was described by Leinfelder (1987): See Fig. 4.29. Holocene: Barthel 1974; Houbolt 1957; Maiklem 1967; Shinn 1973; Shinn et al. 1969; Strasser 1984; Strasser and Davaud 1986; Sugden 1966; Van Straaten 1954; Wagner and van der Togt 1973; Ward et al. 1970. Pleistocene: Beach and Ginsburg 1980; Montenat 1981. Tertiary: Freytet 1982; Häfeli 1966; Platt 1992; Shinn and Lidz 1988. Cretaceous: Barthel 1974; Cotillon 1960; Francis 1986; Lang and Tucci 1997; Rey 1972; Strasser and Davaud 1983. Jurassic: Enay 1980; Seyfried 1980; Leinfelder 1987. Triassic: Bechstädt 1975, 1979; Bechstädt and DöblerHirner 1983; Eppensteiner 1965; Henrich 1984; Piller 1976; Wignall and Twitchett 1999. Carboniferous: Wilson 1967. Ordovician: Sardeson 1914.
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Microlithoclasts Limestones predominantly consisting of small angular to subangular micritic grains (sometimes described as ‘peloids’, Sect. 4.2.2) appoximately equal to or smaller than 0.1 mm are common in deeper-marine shelf, slope and basin settings (Pl. 114/2, Pl. 137/1). These ‘small intraclasts’ (Wilson 1967) differ from peloids in their low sphericity and greater variation in size and shape. Reef clasts Destruction of lithified reef carbonate leads to the production of reef-derived detritus ranging from millimeter-sized debris to huge blocks up to several tens of meters in size. The smaller fraction consists of mostly angular and subangular lithoclasts, which are commonly deposited on fore-reef slopes or transported to lower slope and basinal settings (e.g. calciturbidites). Because the distribution of many reef organisms (framebuilders, foraminifera, algae) is limited to specific parts of the reefs, the biotic composition of the lithoclasts can be used to reconstruct the source area (Pl. 16/6, Pl. 27/ 2, Pl. 114/3, Pl. 115/2).
Black Pebbles
Black pebbles The term refers to special types of carbonate clasts and grains predominantly characterized by the distinctive gray or black color of the particles. Although originally summarized as ‘black pebbles’ (cailloux noirs) two types should now be differentiated: (a) black grains in the sand-size range and (b) dark gray to black angular limestone pebbles (black lithoclasts) millimeters or centimeters in size. Both types occur together with nonblackened counterparts and may be found within the same limestone bed. Black pebbles and black grains are common in lithoclastic packstones (limestone breccias and conglomerates) and oncolitic limestones (forming nuclei of oncoids) but also occur in mud- and wackestones. Black grains: Modern and ancient black grains have been recorded from intertidal and subtidal environments. The grains are usually rounded (whole or broken fossils, peloids, ooids). Together with nonblackened grains they exhibit an often graded salt-and-pepper texture and are common above unconformities and diastems. Intertidal and subtidal blackening of carbonate
Fig. 4.29. ‘Black pebbles’: Black-colored limestone clasts are a valuable tool for identifying partial or complete subaerial exposure of limestones (Strasser and Davaud 1983). The figure displays models for blackening (top) and black-pebble deposition (bottom), based on the study of the Upper Jurassic Ota limestone of the Lusitanian Basin, Portugal (Leinfelder 1987). Blackening is mostly due to plant material. Organic material infiltrates soft or slightly lithified sediments as well as solution cavity fillings within cemented limestones. Black pebbles are in part transported and found in alluvial fans, beach conglomerates and lagoonal settings represented by oncolitic limestones and lithoclastic packstones. Other black pebbles are parautochthonous relicts of brecciated black crusts that were reworked during a transgression.
Significance of Intraclasts
grains was reported from many Holocene carbonates of tropical and nontropical marginal marine environments (Florida, Bahamas, Persian Gulf, Great Barrier Reef/Australia, Tunisia; North Sea) as well as from hypersaline lakes (Ward et al. 1970). Both organic matter and/or iron or manganese sulfide as well as clay minerals have been quoted as the blackening agents (Lang and Tucci 1997; Maiklem 1967). Pilkey et al. (1969) proved the necessity of reducing conditions for blackening in an experiment. The salt-and-pepper texture may be caused by laterally migrating tidal channels that erode blackened grains and redeposit them together with nonblackened grains, or it may be the result of landward migrating mudbanks. Erosion and redeposition transport blackened grains and lithoclasts into various sedimentary environments and, therefore, they are found in near coastal areas and supratidal and inland ponds (Fig. 4.29, Box 4.21 and Box 4.22). Black lithoclasts: Black lithoclasts originate in subaerial environments. Most black lithoclasts are angular
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and show evidence of having been fractured. The clast occur at the top of pedogenic sequences and are not graded. The organic matter absorbed on and infiltrating into the grains is derived from rotted plants (high terrestrial plants, algae) and affects mainly unconsolidated carbonate sediments deposited in terrestrial, supratidal, intertidal and subtidal settings (Platt 1992). Black pebbles are associated with their original environment, such as soils (Bechstädt and Döhler-Hirner 1983; Francis 1986). They are also known from tempestites (Aigner 1982) and beach rocks (Pl. 33/3). Blackening of lithoclasts derived from soilstone crusts in the Florida Keys has been explained as being caused by ‘instantaneous’ forest fire heating (Shinn and Lidz 1988). Heating experiments show that limestone pebbles can be blackened at temperatures between 400 °C and 500 °C. Cementation of subtidal sediments within the vadose zone, followed by brecciation and resedimentation can result in the formation of ‘black pebbles’ (Bechstädt 1975). Lithoclasts deposited in depressions
Box 4.22. Significance of intraclasts in microfacies analysis. Paleoenvironmental proxies Paleocurrent analysis: Intraclast orientation patterns are useful in evaluating wave action and tidal currents in shallow subtidal and peritidal environments (Pl. 18/5). It seems likely that the long axes of intraclasts are aligned transversely to the flow (Lindholm 1980). Imbrication of flat pebbles indicates upcurrent orientation. Orientation patterns of mud clasts in turbidites allow bottom currents to be recognized. Seismic shocks: Intercalations of intraformational flat-pebble conglomerates and laterally continuous intraclastic carbonates in subtidal shelf carbonates were interpreted as indicators of possible tsunami effects related to seismic shocks (Kazmierczak and Goldring 1978; Rüffer 1996). Sedimentation patterns Discontinuous sedimentation: Repeatedly bored and encrusted lime clasts in micritic shelf carbonates can be used to reconstruct sea-bottom conditions (e.g. soft-, firm- or hardground) and lithification stages (Chudziekiewicz and Wieszorek 1985). Hardground intraclasts indicate repeated breaks in sedimentation and may mark discontinuity surfaces. Significance of ‘black pebbles’: The occurrence of ‘black grains’ indicates submarine diastems and unconformities. ‘Black lithoclasts’ in multicolored breccias point to subaerial exposure and the existence of a vegetation cover. Black pebbles occurring within breccias associated with marls and marly dolomites of lagoonal sequences point to cyclic emersion and karstification producing residual sediments. In general, evidence of blackened carbonate material in limestones should stimulate one to think of breaks in marine sedimentation, partial to complete subaerial exposure, reworking of carbonate strata, changes in sea-level fluctuations, existence of terrestrial plant growth and possibly even forest paleo-fires, as well as the existence of nearby terrestrial environments. Reconstruction of lost platforms and sea-level fluctuations Clast analysis (see Sect. 16.3) is of particular value in evaluating platform- and reef-derived lithoclastic carbonates and breccias deposited on the slope or at the base-of-slope (Pl. 17/1). The analyses require field studies (abundance and compositional variations of clasts), microfacies differentiation, and biostratigraphical dating of clasts and enclosing matrix (Pohler and James 1989; Stock 1994; Belka et al. 1996). Clues to diagenetic processes Intraclasts are common in Proterozoic shelf and ramps. The influence of storm wave erosion and subsequent erosion is more probable in the formation of these intraclasts than fair-weather processes, because of the extensive sea-floor cementation caused by microbialites carbonates (Sami and James 1993). Reworked dolomite crusts in tidal carbonates may be indicative of early supratidal dolomitization (Germann 1965).
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Resediments: Intra-, Extra-, Lithoclasts
can be dissolved by organic acids and blackening is induced by bitumen or pyrite. Hardground intraclasts A specific type of lithoclasts are hardground intraclasts (Pl. 21/4, Pl. 139/10) found in deep shelf or basinal settings, sometimes called ‘subsolution clasts’. These grains are mm- to cm-sized irregularly-shaped, predominantly micritic clasts. The surface is impregnated with Fe-Mn oxides (Pl. 23/8) and differs therefore in color from the matrix. Microborings as well as biogenic encrustations (foraminifera, serpulids) are common. The clasts reflect repeated phases of erosion, encrustation/impregnation and accretion. The clasts were originally interpreted as a result of submarine dissolution at the sea bottom during inter-
Plate 16
vals of nondeposition (subsolution clasts: Hollmann 1964), but a multistage origin appears to be more reliable, including early lithification of some of the carbonate mud, followed by exhumation of the lithified parts by weak currents, erosion and encrustation at the surface, redeposition within the mud and additional continued accretionary growth in the sediment. In accordance with this interpretation, the hardground lithoclasts are therefore intraclasts. Partial lithification of modern deep marine carbonates muds and subsequent formation of lithoclasts are known from the Mediterranean (Müller and Fabricius 1974). Another interpretation is reworking of lithified carbonate mud sediments by bioerosion or extreme water energy (Fürsich 1979). This explanation is reliable if associated hardgrounds exhibit burrows.
Resediments: Intra-, Extra- and Lithoclasts
A clast is an eroded and redeposited particle. These processes can take place within the sedimentary basin producing ‘intraclasts’ (–> 7, 8). Erosion of an older lithified sediment from outside the basin in which this particle now becomes deposited produces an ‘extraclast’. The term ‘lithoclast’ was originally coined for erosional detritus brought from an outside source into the basin (= extraclast), but is now frequently used for both intra- and extraclasts, when the genetic origin is unknown. Processes producing intraclasts are currents (e.g. storms); desiccation of muds; local sliding; burrowing and grazing activities of organisms; redeposition of carbonates formed in association with coastal vegetation (‘black pebbles’ –>1, 3) and diagenetic changes in volume during dehydration, compaction, or leaching. ‘Clast analysis’ (the study of microfacies, fossils and geochemical signatures of the clasts) has become an essential approach in basin analysis. It provides data on the long-term history of sedimentation and basin development. 1 Lithoclast limestone formed by storms (proximal tempestite). Clasts are intraclasts (black pebbles) reworked by storms and deposited in floatstone and rudstone layers, alternating with micrite layers. Note the distinct black color of the pebbles. Early Tertiary (Paleocene): Lower Austria. 2 Well-rounded carbonate extraclast distinctly differing in microfacies and diagenetic features from the host sediment. Skeletal grains within the clast are algal spores and shells. Fossils in the host rock are fusulinid foraminifera (F). Note the truncation of the gastropod shell at the boundary of the extraclast. Early Permian: Carnic Alps, Austria. 3 Poorly sorted intraclast wackestone. Note the differences in size, shape and orientation of the clasts, indicating strong reworking of semi-lithified sediment. The dark clasts are ‘black pebbles’ originating from the erosion of pedogenic coastal substrates overgrown by plants (e.g. seagrass or mangroves) and rich in organic matter. The arrow points to spar-filled casts of rootlets. These intraclasts are indicative of the existence of subaerial environments and near-coast depositional sites. Former seagrass zones can be recognized, former mangrove zones are more difficult to decipher (Plaziat 1995). Late Jurassic: Switzerland. 4 Large storm-derived non-carbonate extraclasts (siltstone) together with carbonate peloids. Early Carboniferous (Late Tournaisian): Czatkovice quarry, Cracow Upland, southern Poland. 5 Accumulation of planktonic foraminifera, occurring as clasts in pelagic wackestone containing identical foraminifera. The clast may be produced as coprolite. Late Cretaceous: Northern Tunisia. 6 Fine-grained peloidal grainstone clast indicating submarine erosion of lithified carbonates adjacent to a reef zone recorded by debris with diagnostic reef organisms (Tubiphytes, arrows). Late Permian: Sosio, western Sicily, Italy. 7 Incipient tidal intraclasts. This coarse lithoclastic floatstone, characterized by unfossiliferous micrite clasts, formed as lag deposit in tidal channels. Late Jurassic: subsurface, Kinsau, southern Germany. 8 Precambrian intraclasts, together with ooids (arrow). Altyn Formation: Logan Pass, Glacier Park, Montana, U.S.A.
Plate 16: Resediments: Intra-, Extra-, Lithoclasts
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4.2.8.2 Extraclasts: Strange Foreigners The presence of extrabasinal carbonate and non-carbonate clasts in limestones is often regarded as an unlucky mishap despite the high potential of extraclasts in reconstructing paleobasins and paleosource areas (Zuffa 1980). Common features of extraclasts seen in thin sections are: • extraclasts are often either conspicuously round or angular; • fossils and other grains within the extraclast are truncated at the extraclast boundary; • extraclasts should contain stratigraphically ‘older’ fossils than the matrix; • many extraclasts exhibit criteria pointing to recrystallization, dolomitization or tectonic deformation (e.g. calcite-filled veins) prior to their deposition; • carbonate extraclasts often occur together with noncarbonate clasts such as sandstone or chert fragments (Pl. 114/1) or together with volcanic material (Pl. 26/2). Extraclasts occurring in limestones originate from the drainage of river systems in near-coastal regions, rock-falls at coasts, deposition of erosional material by debris and turbidity flows but also by eolian transport. Less common are transport as stomach stones from reptiles, or ice-transport (dropstones; e.g. Hladil 1991). More common are pebbles transported by seaweeds or drift wood and incorporated within a root framework. Examples were described from Mesozoic shelf and pelagic settings (Mayr 1953; Birkenmaier et al. 1960; Massari and Savazzi 1981; Leinfelder 1994). A summary was given by Etzold and Maus (1990). Of high interest for the reconstruction of paleogeographic and paleoceanographic patterns are exotic clasts (e.g. crystalline rock fragments) found in limestones (Jansa and Carozzi 1970; Lantschner et al. 1996; Rajchel and Myszkowska 1998).The accumulation of extraclasts may result in the formation of breccias (see Sect. 5.3.3). Diagenetic clasts Pressure solution can result in the formation of diagenetic lithoclasts (Huber 1987). An example is shown in Pl. 47/6: The continued solution of the micritic matrix of a former lime mudstone has led to the formation of a marly matrix yielding solution-resistant particles which correspond in shape, size variation and sorting to lithoclasts. The diagenetic origin, however, is shown by stylolites and argillaceous solution residues.
Extraclasts
Basics: Carbonate grains Alexandersson, E.T. (1972): Micritization of carbonate particles: processes of precipitation and dissolution in modern shallow-marine environments. – Universitet Uppsala Geologiska Institut Bulletin, 7, 201-236 Bathurst, R.G.C. (1966): Boring algae, micrite envelopes and lithification of molluscan biosparites. – Geological Journal, 5, 15-32 Beales, F.M. (1958): Ancient sediments of Bahamian type. – American Association of Petroleum Geologists, Bulletin, 42, 1845-1880 Bosellini, A., Ginsburg, R.N. (1971): Form and internal structure of recent algal nodules (Rhodolites) from Bermuda. – Journal of Geology, 79, 669-682 Bosence, D.W.J. (1983a): Description and classification of rhodoliths (rhodoids, rhodolites). – In: Peryt, T.M. (ed.): Coated grains. – 218-224, Berlin (Springer) Bosence, D.W.J. (1983b): The occurrence and ecology of recent rhodoliths - a review. – In: Peryt, T.M. (ed.): Coated grains. – 225-242, Berlin (Springer) Chafetz, H.S. (1986): Marine peloids: a product of bacterially induced precipitation. – Journal of Sedimentary Petrology, 56, 812-817 Chafetz, H.S., Butler, J.C. (1980): Petrology of recent caliche pisolites, spherulites and spelothem deposits from central Texas. – Sedimentology, 27, 497-518 Chudzikiewicz, L., Wieczorek, J. (1985): Bored and encrusted clasts in the Lower Kimmeridgian carbonates at Sobków (SW margin of the Holy Cross Mts., Poland). – Annales Societatis Geologicae Poloniae, 55, 195-206 Dahanayake, K. (1977): Classification of oncoids from Jurassic carbonates of the French Jura. – Sedimentary Geology, 18, 337-353 Dahanayake, K. (1978): Sequential position and environmental significance of differential types of oncoids. – Sedimentary Geology, 20, 301-316 Demicco, R.V., Hardie, L.A. (1994): Sedimentary structures and early diagenetic features of shallow marine carbonate deposits. – SEPM Atlas Series, 1, 265 pp. Dunham, R.J. (1969): Vadose pisolite in the Capitan reef (Permian), New Mexico and Texas. – Society of Economic Paleontologists and Mineralogists, Special Publication, 14, 182-190 Esteban, M.C., Pray, L.C. (1983): Pisoids and pisolite facies (Permian), Guadalupe Mountains, New Mexico and West texas. – In: Peryt, T.M. (ed.): Coated grains. – 503-537, Berlin (Springer) Fabricius, F.H. (1967): Origin of marine ooids and grapestones. – Contributions to Sedimentology, 7, 1-113 Folk, R.L., Chafetz, H.S. (1983): Pisoliths (pisoids) in Quaternary travertines of Tivoli, Italy. – In: Peryt, T.M. (ed.): Coated grains. – 474-487, Berlin (Springer) Gerdes, G., Dunajtschik-Piewak, K., Riege, H., Taher, A.G., Krumbein, W.E., Reineck, H.-E. (1994): Structural diversity of biogenic carbonate particles in microbial mats. – Sedimentology, 41, 1273-1294 Gischler, E., Pisera, A. (1999): Shallow-water rhodoliths from Belize reefs. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 214, 71-93 Gygi, R.A. (1981): Oolitic iron formations: marine or not marine? – Eclogae geologicae Helvetiae, 74, 233-254 Gygi, R.A. (1992): Structure, pattern of distribution and paleobathymetry of Late Jurassic microbialites (stroma-
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Basic Reference: Carbonate Grains
tolites and oncolites) in northern Switzerland. – Eclogae geologicae Helvetiae, 85, 799-824 Hottinger, L. (1983): Neritic macroid genesis, an ecological approach. – In: Peryt, T.M. (ed.): Coated grains. – 38-55, Berlin (Springer) Jones, B., Goodbody, Q.W. (1984): Biological factors in the formation of quiet-water ooids. – Bulletin of the Canadian Petroleum Geology, 32, 190-200 Kalkowsky, E. (1908): Oolith und Stromatolith im norddeutschen Buntsandstein. – Zeitschrift der deutschen geologischen Gesellschaft, 60, 68-125 Kobluk, D.R., Risk, M.J. (1977): Calcification of exposed filaments of endolithic algae, micrite envelope formation and sediment production. – Journal of Sedimentary Petrology, 47, 517-528 Leinfelder, R. (1987): Formation and significance of black pebbles from the Ota limestone (Upper Jurassic, Portugal). – Facies, 17, 159-170 Logan, B., Rezak, R., Ginsburg, R.N. (1964): Classification and environmental significance of algal stromatolites. – Journal of Geology, 72, 68-83 Loreau, J.P., Purser, B.H. (1973): Distribution and ultrastructure of Holocene ooids in the Persian Gulf. – In: Peryt, T. (ed.): Coated grains. – 279-328, Berlin (Springer) Macintyre, I.G. (1985): Submarine cements - the peloidal question. – Society of Economic Paleontologists and Mineralogists, Special Publication, 36, 109-116 Monty, Cl. (1981): Spongiostromate vs. porostromate stromatolites and oncolites. – In: Monty, Cl. (ed.): Phanerozoic stromatolites. Case histories. – 1-4, Berlin (Springer) Opdyke, B.N., Wilkinson, B.H. (1990): Paleolatitude distribution of Phanerozoic marine ooids and cements. – Palaeogeography, Palaeoclimatology, Palaeoecology, 47, 135-148 Perrin, C. (1992): Signification écologique des foraminifères acervulinidés et leur role dans la formation de facies recifaux et organogènes depuis le Paléocene. – Geobios, 25, 725751 Peryt, T.M. (1981): Phanerozoic oncoids: an overview. – Facies, 4, 197-214 Peryt, T.M. (1983): Vadoids. – In: Peryt, T.M. (ed.): Coated grains. – 437-449, Berlin (Springer) Prager, E.J., Ginsburg, R.N. (1989): Carbonate nodule growth on Florida’s outer shelf and its implications to fossil interpretations. – Palaios, 4, 310-317 Reid, R.P. (1987): Nonskeletal peloidal precipitates in Upper Triassic reefs, Yukon Territory (Canada). – Journal of Sedimentary Petrology, 57, 893-900 Reid, R.P., Macintyre, I.G. (1988): Foraminiferal-algal nodules from the eastern Caribbean: growth history and implications on the value of nodules as paleoenvironmental indicators. – Palaios, 3, 424-435 Reid, R.P., Macintyre, I.G. (2000): Microboring versus recrystallization: further insight into the micritization process. – Journal of Sedimentary Research, 70, 24-28 Richter, D.K. (1983): Calcareous ooids: a synopsis. – In: Peryt, T.M. (ed.): Coated grains. – 71-99, Berlin (Springer) Riding, R. (1983): Cyanoliths (Cyanoids): Oncoid formed by calcified cyanophytes. – In: Peryt, T.M. (ed.): Coated grains. – 276-283, Berlin (Springer) Sandberg, P.A. (1975): New interpretations of Great Salt Lake ooids and ancient non-skeletal carbonate mineralogy. – Sedimentology, 22, 497-538 Scholle, P., Ulmer-Scholle, D. (2004): A color guide to the petrography of carbonate rocks, grains, textures, porosity,
diagenesis.– American Association of Petroleum Geologists, Mem., 77, 460 pp. Simone, L. (1981): Ooids: a review. – Earth Science Review, 16, 319-355 Strasser, A. (1984): Black-pebble occurrence and genesis in Holocene carbonate sediments (Florida Keys, Bahamas, and Tunisia). – Journal of Sedimentary Petrology, 54, 1097-1109 Strasser, A., Davaud, E. (1983): Black pebbles of the Purbeckian (Swiss and French Jura): lithology, geochemistry and origin. – Eclogae geologicae Helvetiae, 76, 551-580 Wehrmann, A., Freiwald, A., Zankl, H. (1995): Formation of cold-temperate water multispecies rhodoliths in intertidal gravel pools from Northern Brittany, France. – Senckenbergiana maritima, 26, 51-71 Wignall, P.B., Twichett, R.J. (1999): Unusual intraclastic limestones in Lower Triassic carbonates and their bearing on the aftermaths of the end-Permian mass extinction. – Sedimentology, 46, 303-316 Further reading: K081 (peloids), K084 (cortoids), K085 (oncoids and rhodoids), K082 (ooids), K088 (pisoids and vadoids), K083 (aggregate grains), K086 and K087 (intraclasts and extraclasts).
4.3 Morphometry of Carbonate Grains Morphometrical criteria of sedimentary grains are helpful in evaluating aquatic and eolian transport processes. Transported carbonate grains are characterized by morphological shape, grain size and preservation criteria. Grain shape refers to all aspects of the external morphology, including the gross form (equidimensionality or sphericity), roundness and surface structures (smallscale roughness or smoothness). Form reflects variations in the proportions of grains. Roundness is a measure of the sharpness of grain corners, that is, the smoothness. Surface texture refers to small-scale microrelief markings, such as pits or ridges. Form and roundness tend to be positively correlated; particles that are highly spherical in shape also tend to be well rounded. Roundness can also be correlated with grain size (see Sect. 6.1). Most parameters for estimating form are based on the relative lengths of the three main axes of grains. The most widely used measure for expressing form is sphericity (Wadell 1932; Krumbein 1941; Sneed and Folk 1958). Grain shape is described qualitatively in terms of resemblance to geometric shapes and quantitatively with the help of numerical shape parameters (Syvitski 1991; Hilbrecht 1995). Automated grain shape analysis is performed on two-dimensionally projected silhouette images of sand or coarse silt grains (Pye 1994).
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4.3.1 Intentions and Methods Rounding of carbonate grains is common in grainstones as well as in fine to coarse breccias and conglomerates (Sect. 5.3.3). In grainstones, only fossils and extraclasts are good abrasion indicators. Ooids and peloids are round to begin with, and intraclasts are rounded almost immediately. Rounding is the result of many complex factors. The roundness of fossils and bioclasts depends on the inherent roundness, the roundability of the skeletal grains (related to the microstructure and architecture of shells), grain sizes (coarser particles round faster than finer ones), the vigor of surf action, and the length of time an area is exposed to surf action. Bioclasts that can be used in morphometric studies of carbonate grains include foraminifera, calcareous algae, mollusk shells and echinoderms. The roundness is a function of grain composition, size, type of transport process, and transport distance. Hard, resistant grains, such as quartz, are rounded less readily during transport than less durable grains such as carbonate grains and limestone pebbles, which become well rounded by fluvial transport over distances of just a few kilometers in contrast to quartz pebbles. Downstream roundness increases more rapidly for limestones than for other lithologies (Mills 1979). Experimental abrasion of limestone pebbles indicate no significant difference with regard to the effect of the abrasion of fine- or coarse-grained limestones during calculated transport over distances between 6 and 142 km (Kuenen 1956). Less durable micritic limestones tend to reach their ultimate ellipsoidal shapes in the final stage of transport. Methods: Many parameters have been proposed to describe the shape and roundness of rock particles (Reviews: Barrett 1980; Diepenbroek et al. 1992; Hofmann 1994), but many of them are based on three-dimensional measurements and can not be applied directly to two-dimensional thin-section data. Most shape studies of modern sediments concern quartz grains or sedimentary pebbles. Because large grains and pebbles are commonly more easily rounded than small grains, many models developed for pebble- to cobble-sized grains must be considered with caution in the context of microfacies analyses of carbonate grains. However, a standardized characterization of the morphometric data of carbonate grains is recommended. Grain roundness in thin sections is commonly estimated by reference to a two-dimensional visual comparison scales or charts. These consist of sets of grain images of known roundness and are recommended for rapid estimates of rounding and sphericity. The charts of Krumbein (1941) and Krumbein and Sloss (1955;
Morphometry
Fig. 4.30), Powers (1953), and Folk (1955) are the most widely used comparison scales. Descriptive terms for roundness classes used by most authors and by the Shell Standard Legend (1995) are very angular, angular, subangular, subrounded, rounded and well rounded. The sphericity of grains is described as very elongated, elongated, slightly elongated, slightly spherical, spherical, very spherical (Shell Standard Legend 1995). The chart enhances rapid measurement and closely spaced discrimination of grain characteristics (examples in Pl. 26/ 3, Pl. 27/2, Pl. 33/1, 2, Pl. 114/1, 3, Pl. 115/2). Another useful roundness scale, particularly for calcareous shell material, was provided by Pilkey (1967; Fig. 4.31), who studied sands from the southeastern U.S. Atlantic coast. The author concluded that the roundness of shell material larger than 2 mm obviously reflects wave energy. High-energy beaches produce strikingly well rounded skeletal grains, whereas much of the coarse shell material in low-energy coastal settings
Fig. 4.30. Comparison chart for estimating roundness and sphericity of sand-sized grains (adapted from Krumbein and Sloss 1963). Roundness is roughly shown as an increase (from left to right) in the shape from angular to round (angular, subangular, subrounded, rounded, well rounded). Sphericity is the degree to which the shape of a sedimentary particle approaches that of a sphere. Sphericity values correspond to the ratio of the nominal diameter of the grain to the dianmeter of a circumscribing sphere (generally the longest diameter of the grain). A perfect sphere has a sphericity of 1.0. Roundness values are derived from the Wadell coefficient, computed as the ratio of the average radius of the curvature of the particle image to the radius of the maximum inscribed circle. A perfectly rounded particle (an ooid) has a roundness value of 1.0. Less rounded grains have values less than 1.0. The chart is used in the description of transported bioclasts and clasts in limestones. When a match or near match is found, both the roundness and the sphericity of the particle have been determined. Measurements are made directly from thin sections or from photographs. It is helpful to examine at least 100 carbonate grains per sample and to take a critical look at the type of grains. The letters accompanying the images facilitate the rapid assignment to a specific roundness/sphericity class. See Pl. 114/3 and Pl. 115/2 for examples.
Morphometric Data
Fig. 4.31. Roundness scale for calcareous shell material (after Pilkey et al. 1967). Class 1 fragments exhibit no apparent rounding of edges or corners. Classes 2 and 3 are gradations between the fresh fragments and the highly rounded, essentially cornerless fragments of class 4. Note that the images shown in this chart not only correspond to transport processes but also reflect preservational bias and various taphonomic controls. Grains in class 3 can be the result of the activity of predators, such as crabs, or of the destruction of shells by storms. Class 4 may indicate multiple reworking or rounding in a high-energy environment.
remains unbroken. In contrast, a distinct energy-roundness relationship does not exist for sand-sized calcareous bioclasts. This can be explained by the effect of gradual wearing down rather than fragmentation of coarser material. An increase in rounding of skeletal grains is also known in traverses crossing lagoonal to high-energy environments of modern reef tracts (Folk and Robles 1964, Swinchatt 1965) and ancient reefs.
4.3.2 Significance of Morphometric Data for Carbonate Grains Morphometric data assist in evaluating paleoenvironmental conditions (e.g. in Fig. 4.32) and diagenetic processes, and can additionally be used in the descriptive discrimination of samples, of both geological and archaeological material (e.g. temper grains in ceramics, Sect. 19.5). Environmental indicators: Sand-sized carbonate
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particles can not be directly compared with quartz grains or pebbles with regard to morphometric characteristics, because the particles vary in their ability to resist mechanical destruction and chemical solution, and because carbonate grains and quartz grains differ in their initial shapes. However, changes in grain shapes in association with changes in grain sizes, sorting and modal composition allow one to identify different hydrodynamic levels (Plumley et al. 1962; Catalov 1972; Skupin 1973; see Sect. 12.1.1) and the maturity of calcareous sands (Folk 1962; see Fig. 8.6). A morphometric index alone generally does not allow environments to be clearly distinguished, but the roundness and sphericity values of bioclasts coupled with grain size and sorting offer good results in differentiating environments with different hydrodynamic controls, e.g. reef platforms. The presence of well-rounded extraclasts in limestones points to material subjected to fluvial transport or rounded in coastal environments prior to final deposition (Pl. 16/2). Diagenesis: Variously shaped skeletal grains react differently with regard to sedimentary packing. Knowledge of the inherent grain volume of carbonate sands composed of differently shaped grains allows the dynamic and inhibitory factors that control grain volumes in grainstones to be studied quantitatively. These data assist in understanding relations between burial diagenetic factors, compaction and porosity development. Basics: Morphometry of carbonate grains Barrett, P.J. (1980): The shape of rock particles. A critical review. – Sedimentology, 27, 291-303 Depenbroek, M., Bartolomä, A., Ibekken, H. (1992): How round is round? A new approach to the topic ‘roundness’ by Fourier grain shape analysis. – Sedimentology, 39, 411-422 Folk, R.H. (1955): Student operator error in determination of roundness, sphericity, and grain size. – Journal of Sedimentary Petrology, 25, 297-301 Folk; R.H. (1962): Spectral subdivision of limestones types. – In: Ham, W.E. (ed.): Classification of carbonate rocks. – American Association of Petroleum Geologists, Memoir, 1, 62-84 Folk, R.H., Robles, R. (1964): Carbonate sands of Isle Perez, Alacran Reef Complex, Yucatan. – Journal of Geology, 73, 255-292 Hilbrecht, H. (1995): Computergestützte Methoden in der Morphometrie. – Berliner Geowissenschaftliche Abhandlungen, E16, 765-779 Hofmann, H.J. (1994): Grain-shape indices and isometric graphs. – Journal of Sedimentary Petrology, A64, 916-920 Krumbein, W.C. (1941): Measurement and geological significance of shape and roundness of sedimentary particles. – Journal of Sedimentary Petrology, 22, 64-72 Kuenen, Ph. H. (1956): Experimental abrasion of pebbles. 2. Rolling by current. – Journal of Geology, 68, 427-449 Mills, H.M. (1979): Downstream rounding of pebbles, a
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Morphometry: Example
Fig. 4.32. Morphometry values of carbonate grains coupled with grain size and sorting offer the possibility in differentiating environmental conditions. The sample (Late Jurassic Sulzfluh limestone: Graubünden, Switzerland) exhibits aggregate grains (AG), intraclasts (IC), extraclasts (EC), small ooids (O) and skeletal grains including benthic foraminifera (F), gastropod shells (G) and dasyclad fragments (D) enclosed within cortoids (C). Aggregate grains, intraclasts and extraclasts differ in morphometric criteria: Roundness of aggregate grains is low, their sphericity is high (fields B and G of Fig. 4.30). Intraclasts are rounded and well-rounded and show different sphericity (fields E, O and S of Fig. 4.30). Extraclasts have low roundness and low sphericity values mostly corresponding to field L of the comparison chart. Similar differences exist for grain size: Aggregate grains range between 0.20 and 2 mm, most grains are smaller than 0.50 mm. Intraclasts have sizes between 0.50 and 1 mm, extraclasts between 0.50 and 5 mm. Sorting of the sediment, estimated using the comparison charts depicted in Fig. 6.2 and Fig. 6.3, is poor to very poor. Relative frequency of grains expressed by constituent ranking (see Sect. 6.2.1.4) places aggregate grains in the first position, followed by intraclasts, extraclasts and skeletal grains. The petrographic diversity (Sect. 6.2.1.4) of the skeletal grains is low. The sediment is characterized by a high compositional maturity (Sect. 6.2.1.4), most grains are reworked particles. Scale is 1 mm. Interpretation: The variations in roundness and sphericity of grains, the wide ranges in grain size as well as poor sorting indicate mixing and reworking of grain populations. Roundness and sphericity of aggregate grains correspond to their original shape and indicate only little transport. In contrast the morphometric data of intra- and extraclasts point to transport of particles within the depositional environment. The fenestral fabric of the large extraclast in the center suggests a source in a low-energy protected environment. The depositional fabric of the sample (lithoclastic grainstone), grain size ranges and sorting correspond to the Energy Index IV-3 (see Sect. 12.1.1.2) characterizing a sedimentation of reworked grains in a moderately agitated environment. Shallow water depth of grain source areas is indicated by the dasyclad fragments (Clypeina) and by the lituolinid foraminifera (see Pl. 62/5, 6 and Sect. 10.2.2.1). All microfacies criteria attribute the sample to a differentiated carbonate platform. Poor sorting, mixing of resediments and the deposition of eroded grains together with finer grains point to a deposition within tidal channels. quantitative review. – Journal of Sedimentary Petrology, 49, 295-302 Pettijohn, F.J., Potter, P.E., Siever, R. (1987): Sand and sandstone. Second edition. – 553 pp., Berlin (Springer) Pilkey, O.H., Morton, R.W., Luternauer, J. (1967): The carbonate fraction of beach and dune sands. – Sedimentology, 8, 311-327 Powers, M.C. (1953): A new roundness scale for sedimentary
particles. – Journal of Sedimentary Petrology, 23, 117-119 Sneed, E.D., Folk, R.L. (1958): Pebbles in the lower Colorado River, Texas, a study in particles morphogenesis. – Journal of Geology, 66, 114-150 Wadell, H. (1932): Volume, shape and roundness of rock particles. – Journal of Geology, 40, 443-451 Willetts, B.B., Rice, M.A. (1983): Practical representation of characteristic grain shapes of sands: a comparison of methods. – Sedimentology, 30, 557-565 Further reading: K107
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5 Microfacies Data: Fabrics
Having discussed matrix and grain types of carbonate rocks in previous chapters, this chapter focuses on the fabrics constructed by rock constituents. The term fabric includes textural as well as structural criteria. Most fabrics reflect depositional controls or early diagenetic processes. This chapter starts with the description of common depositional and diagenetic fabrics (Sect. 5.1), continues with criteria on the recognition and significance of discontinuity structures (Sect. 5.2) which are of major importance in sequence stratigraphy, and closes with a review of predominantly postdepositional features (Sect. 5.3).
5.1 Depositional and Diagenetic Fabrics The order of the discussion in this chapter can be used as guideline for studying fabrics in thin sections. A typical investigation should start with an evaluation of topand-bottom criteria, then grain orientation, packing patterns, and indications of lamination and bedding, and finally an assessment of burrowing textures, which in turn may contribute to the formation of nodular structures, along with various diagenetic processes. 5.1.1 Geopetal Fabrics Geopetal fabrics (Sander 1936) indicate the relationship between top and bottom at the time the rock was formed. Recognizing geopetal fabrics in limestones is crucial for understanding the depositional and postdepositional history of carbonate rocks, particularly changes in original sedimentary dips of beds by compaction and tectonic tilting of beds. The comprehensive manual for clarifying top-and-bottom criteria published by Shrock in 1948 is still a very valuable reference. Common geopetal features in carbonate rocks Geopetal structures occur on different scales ranging from meter-scaled crossbedding to millimeter-
scaled fabrics that are the topic of microfacies studies. The most abundant microscopic geopetal criteria are summarized in Fig. 5.1. These fabrics are produced by sedimentological, biological and diagenetic processes. Bedding plane surfaces: The top and bottom surfaces of limestone beds exhibit many geopetal sedimentary structures, e.g. ripple marks, mud cracks, and trace fossils (see Sect. 3.1.1.2). A simple but compelling example are fossils concentrated on the surface of bedding planes (Fig. 5.2). Sedimentary structures: Thin sections of finegrained limestones sometimes show fine-scaled crossbedding and cross-lamination which attenuate and flatten downdip and are truncated above (Harbaugh 1959). Unconformity planes may truncate underlying features. Mechanical deposition on free surfaces: Sand-sized grains can be trapped in bottom reliefs, saucer-shaped shells (Pl. 127/2), or internal cavities in the sediment (Sandfang: Sander 1936). Some examples are shown in Pl. 17/1, 2. Grains may also settle on flat surfaces of oncoids (Pl. 17/2), on flat surfaces of shells, or on lithified discontinuity surfaces (Sect. 5.2). Grain grading characterized by coarse grains, e.g. bioclasts and lithoclasts, at the base of the bed and fine grains at the top is an often used, but not always unequivocal geopetal criterion. Normally graded beds are most commonly found at the base of turbidites (Pl. 17/ 1), but are also produced by waning currents (e.g. ebb tides), burrowing organisms in intertidal and shallow subtidal environments (Rhoads and Stanley 1965), trapping of sediment by sea grass (Wanless 1981), and the effect of hurricanes. Reverse (inverse) grading with the finest material at the base and the coarsest at the top is not uncommon (Pl. 17/1) and is a typical feature of high-concentration gravity flows. It is common in grainand debris-flow deposits, but is also found in numerous kinds of deposits including beach, base-surge and overwash sediments. A frequently employed explanation for the origin of reverse grading is the Bernoulli principle, whereby dispersive pressures near the boundary of deposition tend to drive large particles upward (see Fisher and Mattinson 1968; Sallenger 1979; Naylor
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_5, © Springer-Verlag Berlin Heidelberg 2010
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1980 for discussion). Reverse grading in larger reef cavities may also result from seawater pumping through voids (Scoffin 1993). Internal geopetals: The most commonly used geopetals are fossils, half-moon ooids (Fig. 4.23), borings (Pl. 80/7) and cavities within the sediment, characterized by fine-grained internal sediment fillings at the base and sparry calcite at the top. These fabrics represent frozen spirit levels (Wasserwaagen) which record the water level in the cavity at the time of deposition of the internal sediment. Partial infilling occurs in many fossils, including among others calcareous algae (Pl. 122/2), brachiopods, bivalves, gastropods (Pl. 28/1, Pl. 89/7, 8), serpulids (Pl. 92/1) and cephalopods (Richter 1968). Particularly favorable conditions for the formation of intraskeletal geopetal fillings are low sedimentation
Internal Geopetals
rates (e.g. in stratigraphically condensed horizons) and rapid burial of fossils (e.g. by storm events). Important internal geopetals in micritic reefal limestones are related to cavities with a flat bottom and serrated roofs, known as ‘stromatactis’ (Pl. 17/3, Sect. 5.1.5.3). The fine-grained micritic internal sediment in geopetal cavities consists of micrite, micrite with peloids, or carbonate silt (Aissaoui and Purser 1983). Micrite and peloidal micrite may represent infilled sediment, but some of these internal ‘sediments’ are in fact fine-grained magnesian calcite cements, particularly in carbonate slope deposits (Wilber and Neumann 1993). Other micritic infillings represent microbial sediments. Gray silt-sized internal sediment may correspond to ‘vadose silt’ (Sect. 7.4.2.1). Internal fillings and finegrained matrix sediment often differ distinctly in texture and crystal size (Pl. 28/1).
Plate 17 Geopetal Fabrics of Limestones: Top-and-Bottom Criteria in Thin Sections Both sedimentary and diagenetic features are used to determine the top and bottom of carbonate rocks. Paleontological criteria are also important (in-situ growth structures, light-dependant algal crusts). Geopetal fabrics aid in determinating of the stratigraphic order of successions, reconstructing the dips of paleoslopes and reconstructing of complex reworking and depositional patterns (–> 6). Widely distributed geopetals are particles and mud trapped in bottom reliefs or in intra- or interskeletal cavities (–> 4, 5), grains deposited on flat surfaces (–> 1, 2) and normal grain grading with coarse particles at the bottom and fine grains at the top (–> 1). Truncated microcross-bedding and valves deposited with the curvature facing upwards (convex up) are further top-and-bottom criteria. Abundant geopetals are fossils filled with sediment at the bottom and calcite at the top (–> 3). 1 Grain grading. Calcareous turbidite. The figure shows the lower and the middle section of a turbidite bed characterized by significant changes in packing, sorting and grain size. Geopetal fabrics are represented by peloids deposited on the surfaces of the bivalve shells (lower part of the picture) and vertical grading of lithoclasts and peloids. Note ‘normal’ (lower part) and ‘reverse’ grading of peloids and lithoclasts (upper part), the latter caused by infilling of small grains in interstices of larger grains. The sample is from a section composed of thin-bedded micritic limestones and calciturbidites. The turbidite beds consist of peloids, lithoclasts and benthic biota derived from a nearby coral reef and deposited in a shallow basin. Late Jurassic (Tithonian): Kapfelberg near Kelheim, southern Germany. 2 Deposition of grains on free surfaces. Oncoid rudstone. Geopetals are represented by the uniform deposition of grains on flat surfaces of aligned oncoids (arrows). The nuclei of the oncoids are phylloid algae (Pl. 58). Early Permian: Carnic Alps, Austria. 3 Internal geopetals. Fenestral mudstone. Geopetals: Basal infilling of silt-sized sediment (arrows) in synsedimentary cavities (stromatactis type). Peritidal environment. Late Triassic (Hauptdolomit, Norian): Northern Calcareous Alps, Austria. 4 Geopetal infilling. Storm wave layer consisting of a coquina with brachiopod (B) and trilobite shells (T). Geopetals: Infilling of fine lime mud within and between the fossils after deposition of the shells. Note the divergent orientation of brachiopod valves indicating event deposition, and the multiple cementation of large shelter pores (SP). The sample is an example of a ‘storm wave shell concentration’ (Fürsich and Oschmann 1993), characterized by good preservation of skeletal elements, bimodal sorting (complete shells and shell debris), horizontal lateral orientation and absence of signs of abrasion, bioerosion or encrustation. The limestones originated in subtidal environments between reef mounds (see Pl. 143). Early Devonian (Emsian): Anti-Atlas, southern Morocco. 5 Geopetal deposition of peloids, fine quartz grains and microfossils in a gastropod shell. Note the compositional coincidence of infill and surrounding sediment. Early Permian: Carnic Alps, Austria. 6 Reoriented geopetals. Gravity-flow breccia. The geopetal fabric in intergranular voids (black arrows) is distinctly different from the original sedimentary geopetal fabric (white arrows). This indicates resedimentation of lithified sediment subsequent to submarine cementation as part of a megabreccia (Flügel et al. 1993). The pore-filling sequence consists of a white calcitic crystal seam followed by light gray internal sediments, overlain by micrite and (dark) dolomitized crystal silt. The remaining pore space was occluded with granular calcite (white). Middle Permian (Murgabian): Sosio, Sicily, Italy.
Plate 17: Geopetal Fabrics
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Reoriented Geopetals
Fig. 5.1. Common geopetal fabrics in thin sections of limestones.
Geopetal infilling of fine-grained sediment in laminated microbial and algal carbonates and bioclastic grainstones is a common feature, see Deelman and Decoo (1976) for an experimental analysis of floored interstices in sand-sized carbonate sediments. Intraand interskeletal deposition of fine-grained sediment may be caused by storms (Pl. 17/4). Growth patterns of sessile organisms: Sessile colonial organisms (e.g. calcareous algae, sponges, corals) tend to grow upward guided by light and the availability of nutrients. The growth patterns of these organisms and other fossils preserved in life position are excellent top-and-bottom criteria (Pl. 41/1). Biogenic encrustations on clasts or shell fragments can be used as geopetals if the fragments are not reworked. Abundant encrustations on lithified surfaces are valuable geopetals. Oriented shells: Convex-upward oriented valves are commonly considered as the common depositional pattern occurring in low-energy environments (Futterer 1978; Allen 1990). However, bottom currents and bioturbation as well as differences between large and small-sized valves can considerably change this pattern (Salazar-Jiménez and Frey 1982). Diagenetic criteria: Gravitative and pending cements (Pl. 34/6, Sect. 7.4.2.1) are important geopetal indicators in transitional shallow-marine, marginal-
lacustrine and pedogenic carbonates. Geopetal infillings of dolomite crystal molds with vadose silt during dedolomitization are further top-and-bottom criteria (Pl. 39/8). Reoriented geopetals Synsedimentary sliding and slumping of weakly lithified carbonates can lead to the reorientation of geopetals, particularly of buried shells filled with internal sediment. Burrowing may cause the redeposition of shells (e.g. bivalves, cephalopods, brachiopods) subsequent to the lithification of internal sediment. Internal sediment in intraskeletal voids may lithify more rapidly than the surrounding fine-grained sediment and quickly fill the remaining void space with cement, thus resulting in reworking and displacement of geopetally filled fossils (Wieczorek 1979). Storms may rework geopetally filled shells (Pl. 102/ 1). The solution of sulfates may also produce overturned geopetal fabrics (Folk et al. 1993) Significance of geopetal fabrics in limestones Effect of compaction on reef slopes: Measurement of geopetal fabrics in reef cavities allows the effect of postdepositional processes such as differential compaction and basinward tilting of reef and foreslope sediments to be determined (Saller 1996).
Biofabrics and Grain Orientation
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facies studies and has been successfully applied to foraminiferal and mollusk limestones. Biofabrics of foraminiferal limestones: Clues to potential reservoir qualities Larger foraminifera from different systematic groups contribute to the formation of bioclastic limestones, particularly in the Late Carboniferous and Permian, Late Cretaceous and Cenozoic. Excellent examples of foraminiferal biofabrics are Early Tertiary limestones (Aigner 1985, Racey 1994). The life cycle of Tertiary larger foraminifera includes smaller sexual and larger asexual tests. The ratio between these tests differs in autochthonous and allochthonous accumulations and can be used to recognize hydrodynamically different ramp and shelf settings (Fig. 5.3; Pl. 18/3). Accumulations of Cenozoic larger foraminifera are particularly common in shallow-ma-
Fig. 5.2. Many limestone beds exhibit concentrations of fossils on the top surface of the beds.The photograph shows brachiopods, tentaculitids and ostracods on the bedding surface of fine-grained shelf limestones. Late Silurian (Ludlow): Gotland Island, Sweden. Detail from Samtleben et al. 2000. Scale is 1 cm.
Sediment transport on sedimentary slopes can be deciphered from upside-down geopetals (Pl. 26/7) and strongly diverging geopetals (Pl. 17/6). Recognizing tectonically tilted or overturned beds: Cavity fillings and incomplete fillings in fossils are used to determine the initial dip of tilted and faulted beds (Shrock 1948).
5.1.2 Biofabrics and Grain Orientation Many bioclastic limestones exhibit characteristic biofabrics characterized by specific patterns defined by orientation, packing and sorting of skeletal grains. Biofabrics as well as the statistical orientation of bioclastic and intraclastic grains reflect the hydrodynamic and sedimentological history of skeletal carbonates. Biofabric analysis is a very promising tool for micro-
Fig. 5.3. Biofabric of a Tertiary foraminiferal limestone. The sediment is characterized by an accumulation of nummulitid larger foraminifera (predominantly Nummulites deserti and Nummulites fraasi) forming a bioclastic packstone. Because small asexually produced A-forms and large, sexually produced B-forms respond differently to transport processes, biofabrics can be used to reconstruct depositional conditions and settings, and to estimate primary porosity and permeability. Most foraminiferal tests of the sample represent smallsized A-forms of Nummulites and Assilina tests of equal size. The matrix is micrite. Note imbrication of grains (arrow). The good sorting of the bioclasts is due to transport and indicates redeposition. Eocene (Esna Formation, Ilerdian): Ferafra Oasis, Western Desert, Egypt. Scale is 5 mm.
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rine carbonate environments. Larger foraminifera are ‘biologically standardized’, e.g. the size and shape of their tests depends on biological factors (Seilacher 1973). Since larger foraminiferal tests have high initial porosity, low bulk density, and are easily transported, they can form ideal hydrocarbon reservoirs as seen in the Eocene nummulitid limestones in Tunisia, Libya and northern Oman where reservoir quality is primarily controlled by biofabric type and the degree of fracturing.
Orientation patterns of bioclastic and intraclastic carbonate grains Orientation patterns of grains seen in thin sections reveal depositional factors (e.g. hydrodynamic variables) as well as diagenetic changes during compaction (Bathurst 1987). Fossil orientation: Fossils generally exhibit good orientation patterns and may provide valuable paleo-
Orientation of Fossils
current information on local and regional scales, but only elongated fossils respond systematically to the aligning effects of currents, and are therefore usually used in basin analysis studies (Potter and Pettijohn 1977). For many fossils ambiguities exist as to whether they are oriented transversely or perpendicularly to current patterns (e.g. brachiopods, gastropods). The relationship between fossil orientation and current direction is mainly controlled by the shape and uniformity of particle density. Common fossils used in field-based orientation studies are bivalve and gastropod shells, orthocone cephalopods (Sundquist 1982; Wendt 1995), belemnites, and crinoid columnals (Schwarzacher 1963). Systematic mapping of fossil orientation is commonly based on bedding plane studies, but observations of microfossils in thin sections following the plane of stratification or the bedding-perpendicular plane are also useful in recognizing hydrodynamic conditions on the sea bottom (Stauffer 1962). An impressive example of the contribution of microfossils to basin analysis was
Plate 18 Biofabrics, Grain Orientation, Bedding and Lamination: Indications of Depositional and Climatic Controls Biofabric (–> 3) and grain orientation patterns (–> 1, 5, 7) as well as bedding and lamination (–> 2, 4, 6) reflect depositional and climatic conditions and indicate biological and diagenetic controls during sedimentation. Internal bedding of limestones is recorded by distinct compositional changes (–> 2) and current-aligned orientation patterns of elongate particles (e.g. imbrications, cross-structures, –> 1, 3). Biolaminated fabrics are characterized by thin laminae differing in grain size, texture, biota (–> 4, Pl. 50/2), mineral composition, clay content, and the amount of organic material. Ideally, laminae represent ‘sedimentation units’ formed in short-term intervals. The classification of bedding and lamination is based on structural and textural criteria (Demicco and Hardie 1994, see Sect. 3.1.1.2). 1 Grain orientation: Imbrication structure, indicating current energy. The current ‘direction’ is from right to left. Most grains are cortoids comprising coated shells (CS), fusulinids (F), and echinoderms (E). Open marine platform. Early Permian: Forni Avoltri, Carnia, Italy. 2 Bedding: Interlayering of fine wavy laminae consisting of sand-sized grains (peloid packstone, P) and carbonate mud (mudstone, M). This texture is called ‘heterolithic bedding’ and indicates a depositional environment where current flow varies considerably. Jurassic: Northern Alps. 3 Cross-bedding of foraminiferal shells (Nummulitidae: Assilina and Nummulites). Note the radial hyaline wall structures. The biofabric characterized by chaotically stacked tests, linear contact imbrication und cross-bedding indicate deposition by short-term high energy events. Early Tertiary (Eocene): Agathazell, southern Bavaria, Germany. 4 Biolamination: Alternation of fine wavy laminae corresponding to microbialites (MB), sponges (S), and sediment. Late Triassic (Dachstein limestone, Norian): Begunjscica Mountains, northern Slovenia. 5 Grain orientation: Parallel orientation of tabular limestone clasts caused by current transport. Smaller grains are micrite intraclasts and peloids. Precambrian (Altyn Formation): Logan Pass, Glacier Park, Montana, U.S.A. 6 Lamination consisting of microspar laminae and peloidal laminae. Equant coating of the small peloids by cement rims perhaps points to a bacterial origin of these grains occurring within a peloid/micrite crust (see Pl. 8/5). Reefal limestones. Middle Triassic (Ladinian): Seiser Alm, Dolomites, Italy. 7 Lamination due to parallel oriented densely-packed ‘filaments’ (arrows, see Pl. 113/2). The microfabric of undisturbed hemipelagic laminae is often characterized by a preferred horizontal orientation of elongated particles (Krinsley et al. 1998). Early Jurassic: Djebel Assiz, Northern Tunisia.
Plate 18: Biofabrics, Grain Orientation, Bedding and Lamination
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provided by King (1948) who used fusulinid orientations for interpretating the directions of ancient shorelines. Tentaculites and styliolinids are often currentaligned, but preferred orientations can vary considerably (Wendt and Belka 1991, Hladil et al. 1996). Other current-aligned microfossils are shown in Pl. 74/8 (foraminifera), Pl. 78/2 (sponge spicules), Pl. 94/7 (trilobites), Pl. 57/7 (gymnocodiacean algae), Pl. 112/1 (sponge spicules), Pl. 102/6 (shells), Pl. 105/2 (dasyclad algae), and bivalve shells. Orientation of intraclasts: Penecontemporaneous erosion and redeposition of sediment, indicated by intraclastic beds, is a common feature of peritidal and shallow subtidal carbonates. Measurements of the azimuths of the long axes of intraclasts on bedding planes and in oriented samples assist in reconstructing tidal currents and wave movement (Lindholm 1980). Imbrication of intraclasts or elongated skeletal grains characterized by the preferred dipping of disk-shaped or elongate fragments at an angle to the bedding can be used as indicators of current-flow directions (Pl. 18/1, 3). Flat pebble conglomerates containing imbricate intraclasts are typical in modern shallow-marine carbonates, and also known from ancient carbonate platforms (Whisonant 1987). Imbricate intraclasts document episodic events. Microfacies analysis reflects the source sediments of the intraclasts and may indicate whether the material was transported landward or seaward.
5.1.3 Bedding and Lamination Fabrics Bedding and lamination are caused by changes in depositional, biological and diagenetic controls. Depositional factors primarily include changes in sedimentation rates and in the composition of the sediment, and alternations of sedimentation and non-sedimentation phases. Biological controls are predominantly a result of the interaction of microbes and microbial mats with their physical and chemical environment and their influence on binding and trapping of the sediment. Diagenetic processes leading to bedding structures are dissolution and cementation within the sediment. Types, descriptive criteria and the significance of bedding and lamination structures are discussed in Sect. 3.1.1.2. Although bedding and lamination must be studied on an outcrop scale, a thorough investigation of oriented thin sections is important for recognizing hidden sedimentary structures (e.g. small-scale cross bedding) and cyclic patterns. This chapter describes some of the more common lamination and bedding types seen in
Bedding and Lamination Fabrics
thin sections that are important in microfacies analyses (see Fig. 3.4). Depositional controls on fine-scaled lamination and bedding are recorded by conspicuously planar laminae (Fig. 3.3), distinct changes in texture, composition and grain size (Pl. 18/2), parallel orientation of grains (Pl. 18/7), or abrupt changes in microfacies (Pl. 18/6). The latter can reflect short-term depositional events (e.g. storms, turbidites, mass flows) causing event stratification (Wheatcroft 1990; Seilacher 1991). Depositional processes leading to the formation of laminae are settling-out of grains from suspension, bottom flows and bedload deposition, and debris and grain flows. Common depositional bedding types are parallel bedding, cross-bedding (Pl. 18/3) and normal or reverse graded bedding (Pl. 17/1). For reverse grading see Sect. 5.1.1. Biological controls are indicated by wavy or crinkled laminae, irregular alternations of micritic and sparry layers (Fig. 3.3, Fig. 3.5, Pl. 18/4, Pl. 50/1), intercalations of sedimentary layers and peloidal fabrics (Pl. 18/ 6, Pl. 50/2), and relicts of cyanobacteria (Pl. 53) and other microbes. Microbial mats form on bedding surfaces. Most mats have a characteristic filamentous morphology, the filaments have rigid walls and polymeric sheets which act as sites of agglutination of sedimentary grains and external calcification. Of major importance for the creation of a laminated pattern (biolaminations, Gerdes and Krumbein 1987; Gerdes et al. 1991) is the capacity of the mat-constructing biota to migrate vertically to escape burial by sediments and to recolonize the newly deposited surface. In terms of ecology, biolaminites correspond to successions with formative and consuming stages. Formative stages include primary production via photosynthesis and chemosynthesis. Consuming stages are characterized by anaerobic microbial communities with heterotrophic and anoxygenic phototrophic bacteria. Heterotrophic bacteria trigger the formation of carbonates, silicates and sulfides, and the growth of cements which increase the preservation potential of biolamination structures. The thickness of biolaminated stacks ranges from millimeters to centimeters and even decimeters. Resulting biolaminated structures have been called ‘cryptalgal’ structures (Aitken 1967; Monty 1976). Bioturbation commonly alters or destroys bedding and lamination structures and contributes to a homogenization of the sediment (Reineck 1963, see Sect. 5.1.4). Exceptions exist however: A specific type of biologically controlled bedding is biogenic stratification
Burrowing and Bioturbation Fabrics
produced by infaunal deposit feeders (e.g. worms and crabs living on tidal flats) during feeding and burrow excavation leading to vertical size sorting of the sediment (Meldahl 1987). This bedding is distinct from biogenic graded bedding (Rhoads and Stanley 1965) commonly formed in intertidal sediments by the feeding activity of polychaete worms in that the transition from the lower coarse-grained interval is sharp and not graded. Diagenetic controls are revealed by parallel dissolution seams (Pl. 37/6), horse tail structures (Pl. 37/8), stylolite sets (Pl. 37/3) and significant differences in carbonate and clay contents. Stylobedding and stylolamination are common features of shallow and deep marine limestones (Sect. 7.5.2). Other criteria of diagenetically caused or at least enhanced bedding are cemented layers (Pl. 50/6), stratified cementation, hardgrounds (Sect. 5.2.4.1, Pl. 23) and selective compaction. Diagenetic bedding is a common feature in chalks (Ekdale and Bromley 1988) and in argillaceous limestone and marly sequences of open marine platforms (Bathurst 1987). Diagenetic controls are of major importance in models explaining the origin of marl-limestone alternations in terms of repeating zones of diagenetic carbonate dissolution and cementation (Ricken 1986). These processes considerably enhance primary carbonate variations and can thereby generate rhythmic bedding (Sect. 16.1.1.1). Significance of depositional and biogenic lamination Depositional laminae, produced by variations in sediment input and primary production, are characteristic of hemipelagic sediments and occur in shallow and deep-water settings. Sediment is supplied seasonally and/or episodically. The preservation and composition of laminations of fine-grained carbonate rocks are excellent proxies for paleoclimatic conditions, oxic or anoxic bottom and pore waters, and ecologic constraints of benthic organisms. Thin-section image analysis is a common method of identifying and quantifying sedimentary laminae in the context of paleoclimatic studies of both marine (particularly lagoonal carbonates, Park and Fürsich 2001) and lacustrine sediments (Francus 2001, Trichet et al. 2001).
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fabrics (Ekdale et al. 1984; Bromley 1996; Donovan 1994; Fig. 5.4). These organisms occur in niches at different levels above and within the substrate. Tiering of infauna is reflected by vertically changing ichnofossils. Ichnofabrics are recorded by both discrete trace fossils as well as complicated structures, e.g. bioturbation fabrics. Burrowing is a prominent process in modern subtidal environments, often causing a lack of internal layering. Burrows and bioturbation provide information on conditions of life in sediments (Curran 1991) and, thus, about the depositional environment. They are also valuable indications of synsedimentary and diagenetic deformations (Gaillard and Jautee 1987). Compaction can be directly measured from deformed, originally cylindrical bioturbation tubes using the compaction law developed by Ricken (1987), see Sect. 7.5.1 and Pl. 49/3. Burrows are formed within soft unconsolidated sediments (muds, sands, firmgrounds) by the activity of animals during feeding, resting or migration. The size of burrow structures ranges from millimeters to a few centimeters and up to decacentimeters. The distinction between burrows and borings (biogenic structures drilled into hard sediment or shells) is not always easy to discern (Pl. 23/1). Roots molds and casts (caliche, tidal flat, subtidal settings) can be mistaken for burrows (Pl. 20/6). Most distinctive is the downward branching pattern where there is a marked and systematic decrease in tube diameter at junctions. Roots may have carbonaceous linings and exhibit branching out into clusters of root hairs. Bioturbation, burrowing and bio-retexturing The term bioturbation was proposed for ‘all kinds of displacement within sediments and soils produced by the activity of organisms and plants’ (Richter 1936). The term designates the churning and stirring of sediments by organisms resulting in the destruction of sedimentary structures (e.g. bedding). Because bioturbation is often very loosely used, a new term has been
5.1.4 Burrowing and Bioturbation Fabrics Organisms that travel across, burrow into, rest upon or ingest sediment create trace fossils exhibiting characteristic shapes and patterns and well-defined ichno-
Fig. 5.4. Distinct burrow in a mudstone destroying the dark organic-rich layers. Siliciclastic components enriched within the burrow. Late Devonian: Menorca. Scale is 2 mm.
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proposed. Bio-retexturing is defined as the processes by which burrowing biota may modify or transform original textures by relocating significant volumes of unlithified sediment into or out of the sedimentary layers (Pedley 1992). Occurrence and origin of bioturbation Bioturbation occurs in shallow- and deep-marine environments as well as in lacustrine settings and even in eolian environments (e.g. carbonate dunes: Ahlbrandt et al. 1978). Burrows are the most common features observed in shelf limestones (e.g. Pl. 135/1). Burrowing occurs in carbonates both in intertidal and especially subtidal settings, and in deep-marine bathyal environments where a high density of endobionts in conjunction with low sedimentation rates may cause multiple episodes of burrowing and churning-up of the sediment. Bioturbation is ubiquitous in outer-ramp settings where crustacean-dominated bioturbation results in intensive bio-retexturing of the primary depositional fabrics and masks inorganic process-related structures. Infaunal burrowing organisms are particularly frequent in calcareous muds deposited in shallow-marine
Burrowing and Bioturbation
quiet-water sheltered embayments, e.g. lagoons, muddy tidal flats) and deeper water environments below fairweather and storm-wave base. Recent shallow-marine carbonate environments in Florida and the Bahamas are dominated by burrows of the shrimp Callianassa, which affect up to 75% of the sediment. Repetition of burrow excavation and storm infilling results in the formation of a storm-derived sediment textures dominated by mottled packstone with coarse skeletal patches (Shinn 1968; Tudhope and Scoffin 1984; Curran and White 1991; Wanless and Tedesco 1993). Burrowing is common in many slope sediments (Henderson et al. 1999; Pl. 137/1, 6). The occurrence of pelagic bathyal and abyssal deep-sea bioturbation (Pl. 19/ 2, 3; Pl. 139/1) in the upper few tens of centimeters of the sea bottom depends on the nature of the seafloor sediment, the oxygen content of the bottom and of the pore waters, and the nutrient resources (Ekdale et al. 1984; Wetzel 1984; Meadows and Meadows 1994). Because the rates of deep-sea bioturbation are high, in contrast to the low sedimentation rates, deep-sea bioturbation structures often represent multiple episodes of burrowing (Magwood and Ekdale 1994).
Plate 19 Burrowing and Bioturbation: Sea-Bottom Conditions and Post-Sedimentary Diagenetic Changes Burrows are formed within soft unconsolidated sediments (muds, sands, firmgrounds) by the activity of animals. The term ‘bioturbation’ designates the churning and stirring of sediments by organisms resulting in the destruction of sedimentary structures (e.g. bedding). Bioturbation fabrics are typified in terms of spatial arrangement and geometry of burrows, burrow density (using semi-quantitative ‘bioturbation indices’), burrow abundance, and texture and filling of burrows. These criteria can be studied in the field but can also be applied to large thinsections. Burrows and bioturbation are the most common features observed in shelf limestones and deep-marine carbonates. The common criteria of burrowed limestones is a mottled appearance differing in color and texture (–> 6), or by dolomitization from the host rock. 1 Indications of bioturbation are the circular swirls (arrows) of skeletal debris (trilobite and echinoderm fragments), the patchy distribution of these structures, and variations in packing densities of grains. Well-oxygenated subtidal mid-shelf environment. Red bedded limestones. SMF 9. Ordovician: Öland Island, Sweden. 2 Burrowed radiolarian wackestone. Note the difference in the packing of radiolarian tests in the infilling of the burrow and outside the burrow. Bathyal environment. Late Jurassic: Northern Alps. 3 Burrowed calcisphere packstone. Note different packing of the spheres (corresponding to calciodinoflagellate cysts). The calcisphere content exceeds 35%. The irregular texture is partly caused by synsedimentary sliding of the sediment (Voigt and Häntzschel 1964). Open-marine deep-shelf environment. Red bedded limestones alternating with flint layers. Late Cretaceous (Pläner marl, Late Turonian): Northern Germany. 4 Burrow filling consisting of benthic foraminifera and shell bioclasts. Note absence of compaction indicating firmground substrate consistency. Late Jurassic (Sequanian): Beze, France. 5 Burrow within microbial crusts, infilled with oncoids and coated grains. Upper slope environment. Late Triassic (Raibl Beds): Schlern, Dolomites, Italy. 6 Distinct mottling due to bioturbation in a deep-shelf lime mudstone with pelagic ostracods (see Pl. 93/3,4). Note differences in color shades and compaction of burrows indicating softground substrate consistency. Early Carboniferous (Late Viséan): Betic Cordillera, southern Spain. 7 Burrowing indicated by skeletal grains (predominantly trilobite debris) forming twisted structures. Bedded intermound mid-shelf limestones. Early Devonian (Emsian): Anti-Atlas, southern Morocco.
Plate 19: Burrowing and Bioturbation
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Animals involved in burrowing include a wealth of organisms living for a lifetime in one place or sporadically searching for nutrients and protection (e.g. worms, echinoids, bivalves, arthropods). Burial velocity appears to be a major control to organisms contributing to bioturbation fabrics (Boudreau 1994). Diagnostic criteria of bioturbation fabrics in limestones Small- to medium-scaled burrows diameter ranges from <1 mm (micro-burrows, Das and Rao 1992) to several centimeters. They are often backfilled with pellets (Pl. 117/2), are lined (Pl. 92/10, have internal sprites, are often branched or interconnected (Pl. 25/4), and may have distinct and consistent geometrical patterns and orientations. Bioturbation is indicated by mottled fabrics (see Pl. 19/6, Pl. 139/1), inhomogeneous micrite texture (Pl. 135/1), variously colored burrow structures (Pl. 137/6), whorled alignment of shell material (Pl. 19/1, Pl. 23/1) and differences in grain sizes of burrow infillings as compared with the non-burrowed parts of the limestone (Pl. 19/4). Small- and large-scale burrows may be infilled by differently textured sediment (e.g. wackestone textures within packstone, grainstone within mudstone). Bioturbated and non-bioturbated parts within the same limestone bed can exhibit contrasting textures which are caused by the homogenization and transformation of former grainstones by bioturbation into wackestones or packstones, and the preservation of the grainstone texture outside the burrowed areas. Small burrows in grainstones often exhibit organic and agglutinated linings. Linings are necessary in order to ensure efficient water and oxygen circulation for the organisms. Description and quantitative classification of bioturbation fabrics Important criteria for differentiating bioturbation types are the spatial arrangement and geometry of burrows, the estimated or measured degree of bioturbation (low, medium, high), texture of burrows (considering degree of reworking, breakage of bioclasts), filling (same or different material as compared with the matrix), and differences in diagenetic criteria between burrows and matrix (cement, dolomite). Bioturbation structures are typified in terms of spatial arrangement and geometry of burrows, burrow density (using a semi-quantitative ‘bioturbation index’: Taylor and Goldring 1993; Droser and Bottjer 1986), burrow abundance and texture and filling of burrows. These criteria should be studied in the field but also can be applied to split core samples and large thinsections. Most classification schemes for the amount
Burrowed Limestones
of bioturbation involve semiquantitative categories. These categories are divisions between two extremes: 0% bioturbation (no evidence of biogenic mixing of sediments) and 100% bioturbation (complete biogenic mixing of sediments). The categories between these end members in a continuous spectrum of bioturbation are defined by the ranges of biogenic disturbance, e.g. 1– 10% bioturbation, 10–40% bioturbation, and so on. Names have been given to such categories reflecting some ordinal scale (e.g. ichnofabric index 2, ichnofabric index 3) or general descriptions that express the relative progression of bioturbation (e.g. very slightly bioturbated, slightly bioturbated). Francus (2001) described a method of quantifying bioturbation based on a thin-section image-analysis technique. The method is consistent with semiquantitative methods, but provides higher solutions of biological processes. Secular variations of Phanerozoic bioturbation patterns Recent work demonstrates that late Neoproterozoic sea floors were characterized by well developed microbial mats and poorly developed vertically-oriented bioturbation, thus providing a fairly stable, relatively low water content surface and sharp water-sediment interfaces characterized by firmground substrates. During the Cambrian microbially bound sea floors became increasingly scarce in shallow marine settings. This is largely due the evolution of burrowing organisms along with an increasing vertically oriented and bedding disruptive component as a result of bioturbation (McIlroy and Logan 1999). The Phanerozoic-style sea floor is characterized by a blurry water-sediment interface, greater water content, and lack of well-developed microbial mats (Droser and Bottjer 1993; Bottjer et al. 2000). Large changes in the extent and depth of bioturbation took place in the Late Neoproterozoic and Early Paleozoic, reflecting the increasing influence of bioturbation upon soft sediment fabrics through the Phanerozoic (Bottjer et al. 2000), accompanied by morphological adaptions of organisms living on soft substrates (Thayer 1983). Deeper oxygenation of the sediment with increased mixing via infaunal activity would also have affected shallow burial diagenetic processes of carbonates. Significance of burrowed limestones Bioturbation in carbonate sediments influences the composition of the sediment, controls diagenetic processes and is of importance in evaluating the reservoir potential.
Bioturbation and Diagenesis
Destruction of original structures: Stratification is frequently devastated by deep burrowing (Enos and Perkins 1977; Bromley 1990). Vertical bioturbation of shelf sediments is largely responsible for the homogenization of the sediment and can cause considerable stratigraphical contamination of skeletal grains (Sarnthein 1972), which should be considered in using microfossils seen in burrowed limestones as biostratigraphic markers. Bioturbation may speed up or slow down the degradation of carbonate grains. A turnover of sediment may cause a decrease in grain size through breakage which in turn affects potential dissolution of grains. Burrows and sediment mixing can bring previously buried grains to the surface or can preserve grains in deep and sheltered burrows. Burrowing will strongly influence the taphofacies (Brashaw and Scoffin 2001). Substrate conditions: Burrows are important in recognizing marine substrate types (Sect.12.1.3). The consistency of bioturbated sediment may be found in five states (Goldring 1995): Soupground (incompetent, water saturated sediment), softground (muddy sediment with substantially compacted burrows, see Pl. 19/6), looseground (silt or sand grade sediment with grainor mucus-stabilized walls of burrows, Pl. 19/1, Pl. 92/ 10), firmground (stiff substrate with minor burrow compaction, Pl. 19/4, Pl. 147/1) as well as hardground, rockground and shellground (best recognized by encrusting and boring biota and mineralized crusts, borings cut evenly across grains and matrix alike). Oxygenation: Burrowing is used to evaluate the oxygen distribution and chemical conditions within the sediment and at the sediment-water interface (Sect. 12.1.5). The degree and intensity of bioturbation is an excellent indicator of level-bottom oxygenation and for the estimation of ancient oxygen-deficient environments. Bioturbation tends to increase the subduction of organic matter below the sediment-water interface. Infaunal benthic animals have critical requirements for oxygen in their respiratory processes. A progression of trace fossils with specific life habits exists ranging from reducing settings to well-oxygenated substrates. Particle flux: Bioturbation is an indication for the flux of organic matter and nutrients near the sea bottom, because the intensity of biological mixing in marine sediments is related to the flux of organic matter to the sea bottom. Depositional setting: Bioturbation and burrowing are common in inner and mid-shelf environments. Ver-
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tical burrows are more frequent in intertidal areas. Anastomosing burrows are more common in subtidal sediments. Deep-marine burrows differ from shallow-marine by ichnotaxa association and spatial distribution. Sedimentation rates: The type and abundance of bioturbation fabrics can be used as evidence of high or low sedimentation rates. The ecological stratification of trace fossils within the taphonomically active zones (infaunal tiering) aids in explaining discontinuity surfaces and lack of sedimentation (Monaco et al. 1996). Impact of bioturbation and burrowing on diagenesis: Burrowing and bioturbation create inhomogeneities in permeability and porosity of the sediment, control migration paths of fluids and gases (Libelo et al. 1994), and accelerate elemental cycles (e.g. of carbon, sulphur: Berner and Westrich 1985). Bioturbation has a strong impact on the nitrogen cycle in marine sediments and the distribution of organic matter contaminations. The physical network of burrows within the sediment and the increased roughness of the surface caused by burrow openings are likely to be responsible for increased transfer of dissolved and particulate material between the sediment and the water column (Meadows and Meadows 1994). Differences in permeability and porosity in turn influence the style and course of dolomitization. The finer crystal size of burrow fills as compared with micritic matrix may be one factor controlling the often observed selective dolomitization of burrow fills (Zenger 1992, Pl. 19/6). Bio-retexturing profoundly controls diagenetic processes because it can selectively increase individual bed permeability where coarse burrow backfilling is dominant. Conversely, locally it can inhibit early sea floor cementation by admixing micritic sediment into grain support fabrics. Feeding, burrowing and irrigation increase solute transport and solid phases reaction rates, thus leading to rapid carbonate dissolution (Green et al. 1992). This may explain fluctuations in the frequency of preserved micro- and macrofossils in bioturbated limestones as compared with non-bioturbated carbonates. Economic significance: Bioturbation influences the distribution of permeability and porosity in limestones (Dawson 1981). Burrowing acts on reservoir properties, e.g. as in Early Tertiary limestones offshore Central Tunisia (Hauptmann 2000). The tight network of branching burrows resulting from the activity of echinoids, crabs and bivalves is responsible for the complex fractal distribution patterns of dissolution versus
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Open Space Structures
cementation as well as for partial dolomitization and silicification.
see Pl. 20/2, Pl. 41/2). Two major types of open-space structures – Stromatactis (Dupont 1881) and birdseyes (Ham 1952) – can be distinguished using the following criteria:
5.1.5 Birdseyes, Fenestral Fabrics and Stromatactis Many limestones are characterized by open-space structures (Wolf 1965). The term has been suggested for variously sized cavities in carbonate rocks filled with calcite and/or internal sediment. The term is a non-genetic, designating specific sedimentary and diagenetic voids usually in fine-grained micritic limestones (predominantly mudstones, packstones and bindstones). The voids are conspicuous constituents of the limestones and form characteristic fabrics, some of which resemble blind windows within a dark wall and are, therefore, called fenestral fabrics (Tebbutt et al. 1965; Plate 20
Stromatactis Shape
Flat to undulose base, digitate and dentate roof Filling Mud at the bottom (often planar) and/or calcite spar at the top Orientation Subhorizontal or irregular Geometry Size
Birdseyes Circular, oval, irregular Calcite spar or evaporite mineral Irregularly or parallel to bedding planes Single individuals
Network consisting of interconnected cavities or single individuals Several mm to tens of cm, A few millimeters larger than birdseyes
Open-Space Structures: Birdseyes, Stromatactis, and Fenestral Fabrics
Open-space structure is a non-genetic term designating specific sedimentary and diagenetic voids in carbonate rocks (predominantly mudstones, packstones and bindstones). The mm- to cm-sized voids are conspicuous constituents of the carbonate rocks and form specific fabrics. Some of these fabrics resemble windows and are, therefore, called fenestral fabrics (Tebbutt et al. 1965). The voids are differentiated according to shape and size. Millimeter-sized, irregularly distributed spheroidal spar-filled voids are called birdseyes (–> 7, 8), often associated with microbial bindstones. Larger millimeter- to centimeter-sized voids with flat bases and digitate roofs are known as stromatactis. These voids are commonly filled with fine-grained sediment at the base and sparry cements at the top. Birdseye limestones are formed preferentially in supratidal and intertidal environments, stromatactis-type voids occur more commonly in subtidal environments. Both types are formed by various processes (Box 5.1 and Box 5.2). Significance: Open-space structures are valuable microfacies criteria for indicating ancient depositional settings and environmental conditions, and controlling the porosity and permeability of carbonate rocks. 1 Stromatactis structure in peloidal wackestone. The cm-sized structures occur within a 1.5 m thick horizon in the uppermost part of slope sediments close to a microbial-sponge reef. They are interpreted as the result of an internal erosion of a laminar organic framework (Matyszkiewicz 1993). The roofs of the cavities are lined with thin, marine phreatic isopachous cements, the center with blocky calcite crystals formed under burial or meteoric conditions. Note the basal internal sediment consisting of tiny peloids. Late Jurassic (Oxfordian): Mlynka quarry NW Crakow, southern Poland. 2 Laminoid fenestral fabric (LF-A fabric) Laterally elongated open spaces within pisoid layers. Supratidal environment. SMF 21. Late Triassic (Carnian): Nesza quarry, Budapest, Hungary. 3 LF-B II fabric in bindstones with reworked pisoids. The fabric is characterized by a large amount of detrital sediment. Supratidal/intertidal environment, inner carbonate platform. Early Jurassic (Pantokrator Limestone): Peloponnesus, Greece. 4 Fenestral fabrics within cyanobacterial mats. Note the occurrence of filaments (arrow). The association of cryptalgal structures and fenestral fabrics indicates a peritidal environment. Early Carboniferous (Late Tournaisian): Czatkovice quarry, Crakow Upland, southern Poland. 5 Fenestral fabric. Note pisoids, black pebbles (center), and internal micrite (gray). Supratidal environment. Early Jurassic (Pantokrator Limestone): Peloponnesus, Greece. 6 Not birdseyes! Open bifurcated voids (plant roots and rootlets) within recent caliche. Root molds can be mistaken for burrows but differ in downward branching and decrease in diameter. French Reef, Bahamas. 7 Birdseyes. Small calcite-filled voids in a laminated dolomite mudstone. Planar fine laminae composed of couplets of peloid-rich layers alternating with continuous dark mudstone layers. This is a common texture in peritidal carbonates. Lagoonal intertidal environment. Late Triassic (Hauptdolomit, Norian): Steinplatte, Tyrol, Austria. 8 Close-up of birdseyes. Note sediment infilling. Late Triassic (Hauptdolomit, Norian): Steinplatte, Austria. –> 1: Matyszkiewicz 1993
Plate 20: Birdseyes, Fenestral Fabrics, Stromatactis
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These cavities and fillings are believed to be predominantly of synsedimentary origin and should not be confused with spar-filled interparticle pores, growthcavities in reefs, solution voids and spar-filled fractures, or mineralized lichens and spar-filled molds of plant roots found in caliche crusts (Klappa 1978). Birdseyes occur predominantly in supratidal and intertidal carbonates, stromatactis in limestones formed in deeper and shallow subtidal environments. Other terms in use for limestones with spar-filled voids are dismicrite (disturbed micrite: Folk 1959), describing a micritic limestone with less than 1% grains and containing many birdseyes, and loferite (Fischer 1964), characterizing a limestone riddled by spar-filled shrinkage pores (Sect. 16.1.1.3). Both terms are partly synonymous with birdseyes limestone but the terms birdseyes and stromatactis should be reserved for describing the shape and geometry of the cavities, whereas loferite may be used for carbonate rocks abundant in various kinds of fenestral fabrics formed in a peritidal environment.
5.1.5.1 Birdseyes Birdseyes and birdseye limestones (Pl. 20/7, 8, Pl. 123/ 2) are old names used by American field geologists to describe calcite-filled voids that sparkle in the sunlight. Apparently, the shape and sparkle of these features reminded someone of birds’-eyes. The name was introduced for very small spar-filled voids in Ordovician micritic shelf limestones (Ham 1952). Birdseyes occur as isolated bubble-like vugs 1 to 3 mm in diameter, or as planar isolated vugs 1 to 3 mm high by several millimeter in width forming fenestral fabrics. The structures are often associated with microbial and algal mats. They are found in carbonate rocks from the Precambrian to the Recent and originate predominantly in supratidal, sometimes in upper intertidal settings. Origin of birdseyes Birdseyes have different origins (Box 5.1). Modes 1, 2, 5 and 6 produce birdseyes arranged more or less concordant to bedding planes. Irregularly distributed open-space structures may originate by modes 3, 4 and 7 to 9. Nearly all modes are known from modern shallow-marine environments, some have been verified by experiments (Shinn 1983). The term birdseyes often is rather loosely used, but should be restricted to millimeter-scale spar-filled vugs exhibiting rounded shapes and concentrated in peritidal environments (Shinn 1983). Small vertical to near-
Fenestral Fabrics
vertical tubular voids filled with calcite, in places associated with birdseyes, are most likely root tubes or burrows and should be named ‘pseudo-birdseyes’ (Shinn 1982). These tubular structures of constant size often occur in tidal and subtidal fine-grained limestones (Wu 1983). 5.1.5.2 Fenestral Fabrics Birdseyes (but also stromatactis-type structures) are important constituents of fenestral fabrics. Tebbutt et al. (1965) suggested the name fenestra or window to characterize ... ‘a primary or penecontemporaneous gap in rock framework, larger than grain-supported interstices’. Fenestrae may be open cavities, or may be completely or partially filled by surface-derived internal sediment, diagenetic internal sediment (e.g. crystal silt) or cements. Fenestrae have no apparent support in the framework of primary grains composing the sediment. In that respect, they differ from growth-framework pores or from solution voids. The term fenestra is now in use for any kind of a hole of any shape in a rock or unconsolidated sediment that is not of an intergranular nature (Demicco and Hardie 1994). It seems, however, better to restrict the term to fabric-selective textures, and to use the term void for non-fabric selective cavities. Terminology of fenestral fabrics Fenestrae are arranged concordant to stratification, cross-bedded, or irregularly distributed. Arrangement Box 5.1. Proposed origin of birdseyes structures. Direct or indirect organic interactions 1 Spar-filled interspaces in and between algal and cyanobacterial mats: Ham 1952; Philcox 1963; Deelman 1972 2 Cavities supported by algae and enlarged by receding tidal waters: Wolf 1965 3 Cavities formed by burrowing organisms: Fruth and Scherreiks 1975 4 Gas bubbles caused by decaying organisms: Korn 1932; Hodgson 1958; Cloud 1960; Shinn 1968 Inorganic 5 Desiccation, shrinkage pores in tidal zones: Fischer 1964 6 Air inclusions formed during subaerial exposure and subsequent tidal reinduration: Bolliger and Burri 1970; Deelman 1972 7 Different behavior of capillary and non-capillary sediment in comparison to water: Deelman 1972 8 Leaching of anhydrite 9 Late diagenetic selective recrystallization (‘pseudosparitization’): Wolf 1965; Zorn 1971
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Fenestral Fabrics
and shape is used in classifying laminoid-fenestral fabrics (LF fabrics) characterized by elongate horizontal fenestrae within fine-grained or grain-supported sediment (Fig. 5.5): (a) LF-A fabrics – with distinct laminoid fenestrae (Pl. 132/6, Pl. 142/13), and (b) LF-B fabrics – with irregular laminoid fabrics. The latter type is subdivided into low-detrital fabrics (LB-1) with voids embedded in micrite, and high-detrital fabrics with voids occurring in sediment rich in grains (LB-2, Pl. 8/6, Pl. 122/1). Origin and distribution of fenestral fabrics The shape and orientation of fenestrae reflect different origins: (1) Laminoid fenestrae (Pl. 41/2) are of polygenic origin and may be caused by (a) wetting and drying of carbonate mud in supratidal settings (Shinn 1968), (b) drying out of the surface of cyanobacterial mats resulting in wrinkling, lifting and separation from the adjacent sediment (Hardie and Ginsburg 1977; Davies 1970; Logan et al. 1974), and (c) degassing of decaying organic material connected with compaction of subspherical gas bubbles. Laminoid fenestrae are common in modern intertidal to supratidal settings. However, these structures
develop (as do cyanobacterial mats) in a variety of subaqueous marine and non-marine environments, too. (2) Tubular fenestrae originate in modern intertidal and shallow-subtidal environments as burrows or as root holes, but are also formed by upward escape of gas bubbles biogenically produced within the sediment. (3) Spherical to subspherical fenestrae are produced by air and gas bubbles trapped during the deposition of the host sediment or they are generated by post-depositional decay of organic matter. Air bubbles can be transported within the sediment by rising groundwater. A particular type of mm- to cm-sized spherical and subspherical open fenestrae are keystone vugs (Emery 1945), see Pl. 29/9. The vugs are larger than ordinary birdseyes and are produced by trapping of air-bubbles during storm deposition in the squash zone on beaches or in the sheetwash zone on tidal flats or playas (Shinn 1968; Dunham 1970; Deelmann 1972). (4) Irregularly formed and distributed fenestrae (Pl. 40/4) may form by (a) desiccation, (b) soft-sediment deformation of other originally more regular-shaped fenestrae, (c) gas bubbles, (d) evaporite molds, (e) burrows, (f) burial of pustular cyanobacterial mats, (g) dewatering of gel-like carbonate muds, and (h) replacement of grains. Some of these ‘irregular’ structures are not fenestrae but rather intergranular pores of grain- and packstones. Shinn (1983) proposed the term ‘pseudo-fenestrae’ for these structures. Note that shelter pores in subtidal lithoclast grainstones can strongly resemble vuggy laminoid fenestrae (Goldhammer et al. 1987, Hardie et al. 1986).
5.1.5.3 Stromatactis
Fig. 5.5. Subtypes of laminoid-fenestral fabrics. Type A and type B are distinguished by shape and distribution of the sparfilled fenestral voids (here white), type B-I and B-II by the relative amount of voids and sediment. Adapted from Tebbutt et al. (1965) and Müller-Jungbluth and Toschek (1969).
These structures, characterized by elongated cavities with irregular tops and flat bases, were first observed in Frasnian and Early Carboniferous carbonate mounds of Belgium (Dupont 1881, 1883) and were considered as recrystallized fossils. Most authors now agree, that the name stromatactis should refer to a spar-filled (and sometimes mud-filled) body commonly embedded in micritic limestones that originated from centripetal cementation of cavities by fibrous or radiaxial cements. Some authors restrict the term to a spar network formed by interconnected cavities and exhibiting a reticulate distribution (Bourque and Boulvain 1993), others also incorporate non-interconnected spar bodies occurring underneath platy skeletons of stromatoporoids or corals (Tsien 1983; Flajs and Hüssner 1993). These structures resulted from cementation of synsedimentary shelter cavities or cavi-
194
ties formed by the collapse of mud underneath a rigid object acting as an umbrella during water escape and mud compaction. Stromatactis is abundant in micritic limestones, but geometrically similar open-spacestructures also occur in cemented grainstone crusts (Macdonald et al. 1994). The internal sediment in the cavities of stromatactis commonly is fine-grained micrite, sometimes with microfossils, peloidal micrite, or laminated silt-sized sediment. Genetically it corresponds to the infill of finegrained sediment in a reticulated cavity system or to diagenetic products caused by recrystallization of mud and deposition of solution products (Aissaoui and Purser 1983; Flajs and Hüssner 1993; Vigener 1996; Aubrecht and Szulc 2000; see Sect. 7.4). A common feature of internal sediments is a vague and indistinct lower boundary.
Origin of Stromatactis Fabrics
Common cements occurring within stromatactis cavities are radiaxial, fibrous and granular calcite. Fibrous and radiaxial cements are considered as being of early synsedimentary origin, formed a few centimeters or decacentimeters beneath the depositional surface. Early diagenetic fibrous cements first originate in small voids enlarged by erosional and solution processes. Other cement types are formed during shallow and deeper burial phases. Early cementation contributes greatly to rapid lithification and stabilization of carbonates formed at the flanks of mounds. Stromatactis fabrics occur in patchy distribution or as layers which alternate with limestones consisting predominantly of micrite (Fig. 5.6). These alternations, observed in bedded Devonian stromatactis carbonates, have been interpreted as reflecting Milankovitch cycles
Box 5.2. Proposed origin of stromatactis fabrics. Organic 1 Recrystallized fossils: Dupont 1881; Scoffin 1987 2 Permineralized, unfossilizable organism (e.g. algae): Lecompte1937; Tsien 1983 3 Cavities formed after the decay of various soft-bodied organisms in the sediment: Bathurst 1959; Philcox 1963; Otte and Parks 1963; Lees 1964 ; Davies and Nassichuk 1990; Lees and Miller 1985; Warnke and Meischner 1995 4 Burrow network such as Callianassa: Shinn 1968 5 Action of various burrowing organisms: Shinn 1968 6 Non-skeletal algae: Coron and Textoris 1974 7 Thrombolites, microbial: Pratt 1982 8 Diagenetic CaCO3 replacement of large microbial accretions: Tsien 1985; Weller 1989 9 Winnowing of sediment between microbial framework mats on the sea floor: Pratt 1982, 1986 10 Cementation of cavities left after the decay of microbial mats: Flajs and Hüssner 1993; Kaufmann 1998; Vigener 1996 11 Collapse of material within organic tissue of probable microbial origin: Lees 1964, 1988; Lees and Miller 1985 12 Partial cementation of cavities created by collapse of unlithified sponge tissue: Bourque and Gignac 1983; 1986; Beauchamp 1989; Bourque and Boulvain 1993; Vigener 1996 13 Early marine cementation of cavity networks created by exacavation of uncemented material in partly indurated, decaying sponges and organic mats: Bourque and Boulvain 1993; Dicken 1996 14 Recrystallized bryozoans: Lowenstam 1950; Textoris and Carozzi 1964 15 Sheltering of sediment by bryozoan sheets: Schwarzacher 1961; Lees and Miller 1995 16 Action of plant roots: Bechstädt 1974 Inorganic 17 Primary cavities and/or dissolution cavities: Kukal 1971 18 Cavities produced by dewatering and compaction of a thixotropic mud: Heckel 1972; Macdonald et al. 1994; Bernet-Rollande et al. 1982 19 Excavations between layers of early-cemented crusts on the sea floor, differential cementation of host peloidal mudstone into successive, closely-spaced hardgrounds: Bathurst 1980, 1982; Wallace 1987 20 Downslope creeping and slumping of partially cemented mud on mound flanks: Schwarzacher 1961; Bathurst 1982; MacDonald et al. 1994 21 Internal erosion of cyanobacterial growth cavities on a partly lithified slope: Matyszkiewicz 1993 22) Enlargement of pre-existing cavities (e.g. vertical water and gas escape structures) in unconsolidated bioclastic lime mud by internal erosion and internal reworking: Wallace 1987; Shen and Zhang 1997; Vennin 1997 23 Forcing apart and fracturing of variably bound and lithified sediment by episodic earthquakes: Pratt 1995 24 Subaerial solution (e.g. karstification) resulting in the formation of voids subsequently filled with carbonate silt and cement: Semeniuk 1971 25 Recrystallized patches of the host lime mud: Black 1952; Orme and Brown 1963; Wolf 1965; Ross et al. 1975 26 Pressure solution connected with low-grade metamorphism: Logan and Semeniuk 1976 27 Creation of cavity systems by hydrothermal endo-upwelling: Belka 1998
Stromatactis
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Fig. 5.6. Stromatactis limestone. Slope of a mudmound. Note the difference in color and microfacies between the red lime mudstones with spar-filled stromatactis structures and the overlying red and pink echinoderm wackestones, interpreted as storm-transported sediment. The stromatactis cavities are filled with several generations of fibrous calcite cement. The matrix contains sponge spicula. Early Devonian (Emsian): Notre Dame de Cros quarry, Montage Noire, southern France. Scale is 5 cm. After Vigener (1996).
(Flajs et al. 1996). The spar network is often preferentially aligned along the bedding planes and can be continuous for tens of centimeters to a few meters.
Origin of stromatactis The origin of stromatactis is still controversial (see reviews by Bathurst 1982; Flajs and Hüssner 1993; Lees 1988; Boulvain 1990; Monty 1995). Central problems are the reasons for the shape, how the cavities were formed, and the filling mechanism of the cavities. Up to now more than twenty-five, partly related possibilities have been discussed for processes leading to the formation of stromatactis (Box 5.2). The fundamental question is to ascertain whether stromatactis reflects a taphonomic process (e.g. the decay of a former softbodied organisms) or is a product of physical and chemical processes (e.g. winnowing, internal erosion, dissolution). Two models, both assuming dominant organic control, are currently under discussion. A key question is the relative role of sponges versus microbes in creating cavity systems: (a) Sponge model (Bourque and Boulvain 1993): The model (Nos. 12 and 13 in Box 5.2) integrates the origin of stromatactis (decay of sponge soft tissue and
subsequent internal erosion) with that of the host limestone and also considers the origin of the red color of the limestone (a feature common in many Cambrian to Devonian stromatactis carbonates). The model postulates that the formation of the stromatactis mounds was controlled by sponges, whereas stromatactis and the common red color of many stromatactis mounds are early-diagenetic features formed a few meters below the substrate-water interface. The model was developed for Cambrian to Devonian red stromatactis mounds formed in deeper shelf settings. Criteria used as indications of former sponges are the occurrence of stromatactis together with peloidal relicts of sponges and the abundance of sponge spicula in limestones with stromatactis. (b) Microbial model (Pratt 1982; Flajs and Hüssner 1993; Kaufmann 1996): This model (Nos. 6-11 in Box 5.2) underlines the role of microbial mats in contrast to that of metazoans. Criteria taken as indications for microbial origin are the occurrence of stromatactis in extended layers attaining sizes of several hundred square meters, the joint occurrence of stromatactis and cryptalgal structures, the presence of a reticulated former cavity system, the digitation of the structures towards the top, the association with peloidal fabrics as well as scarce relics of cyanobacterial structures.
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Fig. 5.7. Stromatactis cavities on the flanks of a Waulsortian mudmound. Section normal to bedding plane. The elongated cavities are filled with gray internal sediment at the base and recrystallized radiaxial carbonate cements at the top. Early Carboniferous: Ireland. Scale is 5 mm.
A combination of the microbial with the sponge model seems reliable for the genetic interpretation of Early Devonian stromatactis mounds and shelf limestones (e.g. Dieken 1996; Vigener 1996). High-frequency sea-level fluctuations governing nutrient and sediment input are considered as major controls for the formation of these stromatactis limestones. Models assuming a predominantly inorganic control of the formation of stromatactis cavities (Nos. 1923 in Box 5.2) concentrate on the frequency of stromatactis limestones on the flanks of carbonate buildups, sometimes underlying mass-flow deposits (Steiger and Jansa 1984; Matyszkiewicz 1993) and thus indicating the possibility that slumping of partly lithified lime mud was of major importance (Bathurst 1982). Paleoenvironmental setting of stromatactis Devonian stromatactis structures are common in mudstones and wackestones of deep-shelf mud mounds, but also occur in shallow shelf settings (e.g. detrital crinoid limestones). The former type is commonly associated with sponge-microbial assemblages, the latter with trilobites and brachiopods (Dieken 1996). Flanks of Paleozoic and to a minor degree Mesozoic mounds and reefs as well as slopes of platform margins are common settings of subtidal stromatactis formation. Mounds with abundant stromatactis structures originate in different quiet-water environments: (a) below wave base (Bourque et al. 1986), (b) near the storm wave base (Kaufmann 1996), (c) between the fair-weather wave base and the storm-wave base (Brachert et al.
Zebra Limestones
1992; Flajs and Hüssner 1993), (d) below storm wave base (Bourque and Boulvain 1993; Kaufmann 1996). A common assumption is that stromatactis structures usually originate in mounds most likely formed below the photic zone, because of the abundant lack of calcareous algae. Stromatactis-bearing mounds are widely distributed in bathymetrically different ramp settings, particularly in the Devonian and Early Carboniferous of Europe, North America and North Africa (e.g. Carboniferous ‘Waulsortian’ facies, Fig. 5.7), but also originated at the top of submarine volcanic swells. Note that stromatactis is not only restricted to deeper subtidal settings, as shown by the infilling of shrinkage voids corresponding in shape and structure to stromatactis with ‘vadose silt’ (Pl. 17/3) and the association of birdseyes and stromatactis in loferite-type bedded limestones. Zebra rocks A conspicuous fabric found in peritidal carbonates of the Lofer cyclothems and somewhat reminescent of layered stromatactis structures has been called zebra fabric (Fischer 1964). These structures range in width Box 5.3. Selected case studies of stromatactis structures in carbonate mounds. Note the apparent absence of Triassic stromatactis mounds and the scarceness of Mesozoic records as compared with the Paleozoic (see Neuweiler et al. 2001). Cretaceous: Reitner 1986; Camoin and Maurin 1988; Neuweiler 1993 Jurassic: Matysziewicz 1993, 1997; Mathur 1975 Permian: Blendinger et al. 1997; Davies and Nassichuk 1990; Otte and Parks 1963 Late Carboniferous: Keim and Brandner 1999 Early Carboniferous: Bathurst 1959; Black 1952; Dupont 1883; Orme and Browne 1963; Hodgson 1968; Lees 1964; Lees and Miller 1985, 1995; Miller 1986; Schwarzacher 1961; Lees 1964; Warnke and Meischner 1995 Devonian: Aubert 1989; Bathurst 1980; Belliere 1953; Boulvain and Coen-Aubert 1992; Bourrouilh 1987; Brachert et al. 1992; Dieken 1996; Dupont 1881; Flajs and Hüssner 1993; Flajs et al. 1996; Gnoli et al. 1981; Heckel 1972; Kaufmann 1998; Kerans 1986; Kerans et al. 1986; Kukal 1971; Lecompte 1936, 1937; Macdonald et al. 1994; Monty et al. 1982; Philcox 1963; Playford 1984; Schwarz 1927; Tsien 1983, 1985; Vigener 1996; Wallace 1987; Weller 1991; Wendt et al. 1993 Silurian: Bourque and Gignac 1983; Textoris and Carozzi 1964; Lowenstam 1950 Ordovician: Bathurst 1982; Grover and Read 1978;;Middleton 1988; Ross et al. 1975; Stenzel and James 1995; Krause 2001; Woods 2001 Cambrian: James and Gravestock 1990
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Zebra Limestones
from microscopic to decimeter-sized, and may be meters long. They may have simple cement fills or complex banding of cement and internal sediment. Interbedded dark micritic limestones and cm-sized, often crinkled white calcite layers produce characteristic zebra-like stripes in the rock. This fabric also occurs in ancient beachrocks (Gray and Adams 1995) as well as in Paleozoic carbonate mud mounds (Bosence and Bridges 1995, Fig. 5.8). Although originally explained as parallel sets of sheet cracks, current explanations are just as vague as those for stromatactis, and include collapse of thixotropic carbonate gels (Heckel 1972), lithified crusts with winnowed-out intercrust layers, cement-filled sheet cavities (Bathurst 1980), superposed microbial mats (Pratt 1982), or, soft sediment dilation and deformation associated with slumping.
Krause (2001) put forward an interesting interpretation of zebra fabrics and stromatactis structures of a small Ordovician mudmound in Nevada. He compared zebra and stromatactis limestones with structures triggered by gas clathrate hydrates. In this scenario, the carbonate cement layers would have precipitated and stabilized the zebra and stromatactis fabric, both by consolidating and dissociating gas clathrate hydrates, and oxidizing and reduzing associated gases.
5.1.5.4 Significance of Birdseyes, Fenestral Fabrics and Stromatactis Open-space structures and fenestral fabrics should be evaluated in terms of paleoenvironment and relative water depth, diagenetic history and associated reservoir rock properties: Paleoenvironmental proxies: Birdseyes as well as laminoid-fenestral fabrics occur preferentially in shallow near-coast supratidal and upper intertidal environments. Birdseyes fabrics are not exclusively limited to marine environment. They are also known from lacustrine carbonates and even from eolianites, where rainstorm-induced fenestrae may occur high above sea level (Bain and Kindler 1994). Abundant stromatactis structures indicate deeper subtidal environments but stromatactis formed in intertidal settings must also be taken into account. Interand subtidal setting of stromatactis are distinguished by the composition of internal sediments, carbonate cement types and associated biota.
Fig. 5.8. Zebra limestone. The spectacular pattern of the rock is characterized by alternations of dark wackestones and white parallel layers consisting of several generations of calcite. Some authors consider this fabric as a variation of stromatactis fabrics, but zebra rocks should be discussed on their own despite coincidences in supposed genetic interpretations and the time range of occurrence (Cambrian to Jurassic). Polished section. Lower Late Devonian (Frasnian): Harz, Germany. After Weller (1991). Scale is 1 cm.
Diagenesis and reservoir potential: The formation and filling of open-space structures is critical in early and burial diagenesis and porosity history (Bourque and Raymond 1994; Grover and Read 1978; Mazzullo and Birdwell 1989). Fenestral porosity (see Sect. 7.3.1) is commonly associated with supratidal, mud-dominated bindstones and mudstones. Primary fenestral porosity may be high, up to 65% (Enos and Sawatsky 1981), and permeability in birdseyes limestones can be increased significantly by dolomitization opening communication between fenestral pores through intercrystalline porosity in the matrix dolomite. Industrial use of fenestral limestones: Devonian reefs in Belgium and France yield spectacular red-colored stromatactis limestones which are widely used as decorative stones and for works of art.
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5.1.6 Nodular Fabrics Bedded nodular limestones exhibit a conspicuous fabric characterized by centimeter- and decimeter-sized, often ‘rounded’ nodules floating within a usually micritic matrix. The overall appearance sometimes resembles a conglomerate. Descriptive criteria of nodular fabrics Criteria important for the discussion of the origin of nodular fabrics relate to nodules, matrix, and spatial arrangement of nodules. Carbonate/clay ratio, grain size distribution, and burrowing are additional criteria. Nodules are described with regard to their composition, shape, boundaries and size. Composition (microfacies, fossils within and at the periphery of the nodules) can be the same as or different from that of the surrounding matrix. Many nodules appear ‘rounded’ to ‘subrounded’. ‘Subangular’ and ‘angular’ nodules may indicate reworking and redeposition, these nodules may form conglomerates or even breccias. Apparent rounding, however, can also be caused by solution
Nodular Fabrics
processes (Pl. 21/3). Boundaries between nodules and matrix are inconspicuous or distinct. Distinct boundaries are often marked by a dark clay-rich seam (Pl. 21/2). Isolated nodules exhibiting black ferruginous and manganese coatings often represent ‘hardground clasts’. Nodule size (largest diameter) and orientation may be of interest in the discussion of a possible transport, e.g. along a slope relief. Nodular limestones are usually mud-supported. The matrix corresponds to a mudstone or a wackestone, often with small bioclasts. Variations in the color of matrix and nodules are due to differences in the amount of non-carbonates, Fe-hydroxides, sulfides, and organic carbon. Origin of nodular limestone fabrics The origin of the nodular fabric has been explained by diagenetic, sedimentary and tectonic processes. Diagenetic processes include solution processes as well as cementation and nodule growth within the sediment. Sedimentary models underline the role of transport and redeposition. Tectonic explanations favor the forma-
Plate 21 Nodular Fabrics Nodular fabrics of limestones exhibit scattered or densely packed nodules within a matrix of like or unlike character. Nodules are generally cm-sized and often have ‘rounded’ outlines. Contacts between nodules and matrix or between adjacent nodules are indistinct (indicating bioturbation and diagenetic overprint) or sharp (indicating mechanical transport and sorting of nodules, pebbles or lithoclasts) and sometimes accentuated by darker seams (caused by pressure solution). Many nodular limestones exhibit striking red colors which have led to the term ‘Rosso Ammonitico’ for pelagic and hemipelagic red and variously colored limestones in the AlpineMediterranean region (Farinacci and Elmi 1981). Red nodular limestones often are associated with condensed pelagic limestones and resedimented beds and occur on submarine rises (structural highs and volcanic ridges). Nodular fabrics are common in deeper-marine and slope environments (e.g. Devonian ‘griotte facies’ and Tournaisian ‘calcaire noduleux’ of France, Jurassic ‘Ammonitico rosso’ facies of the Alpine-Mediterranean region) but also can be formed in shallow-marine and non-marine settings. The Early Jurassic nodular Adnet limestone of Salzburg (Austria) shown on this plate is explained by early diagenetic partial cementation and later compactional pressure solution (Pl. 139; Böhm 1992). 1 Red nodular limestone. Initial stage of nodule formation, interpreted as local and partial cementation of the lime mud. Note the indistinct boundaries (arrows) between the nodules. Nodules correspond to bioclastic wackestones. Skeletal grains are echinoderms and brachiopod shells. Kiefer-Bruch, Adnet near Salzburg, Austria. 2 Nodular limestone. Boundaries are marked by dark clay seams (arrows) resulting from incipient pressure solution. The sediment is a fine-grained bioclastic wackestone. Larger fossil fragments are brachiopods (B) and bivalve shells. ScheckBruch, Adnet, Austria. 3 Nodular fabric becomes more pronounced by pressure solution. Stylolitization produces an iden-supported fabric (see Pl. 37/7). Adnet, Austria. 4 Red crinoid-shell wackestone. A nodular fabric is formed by redeposited hardground intraclasts (arrow, see Pl. 139/10). Kiefer-Bruch, Adnet, Austria. 5 Beginning nodule formation explained as partial cementation within isolated areas of the bioclastic lime mud. Early Jurassic: Wolfgrub-Bruch, Adnet, Austria. 6 Typical nodular limestone. Darker nodules are red-colored due to Fe-oxides and a higher organic content. Note the high degree of rounding. Adnet, Austria.
Plate 21: Nodular Fabrics
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200
tion of nodular fabrics by shear processes affecting limestone-marl/clay alternations. Reviews of genetic interpretations were given by Gründel and Rösler (1963) and Böhm (1992). Most models have been developed for Ordovician and Silurian (Möller and Kringan 1988), Devonian (Wendt 1982) and Jurassic (Elsmy 1981) nodular limestones. The models can be classified into three groups: 1 Nodule formation caused by solution processes and explained as (a) Relicts of an intensive submarine dissolution of carbonate at the sea bottom (subsolution: Hollmann 1962, 1964), or (b) Result of the different solubility of clay-rich and carbonate-rich sediments during late diagenetic processes connected with pressure solution (Bathurst 1987). 2 Nodule formation caused by localized cementation within the sediment: Explanations of cementation seeds and the sources of carbonate necessary for cementation include: (a) Bacterially-induced nodule growth (Jeans 1980) (b) Formation of synsedimentary concretions by concentration of carbonate caused by decaying remains of organisms (Kukal 1975) (c) The term pseudonodular limestone. This refers to Upper Jurassic sponge limestones composed of densely-packed, rounded-oval to subangular carbonate nodules up to 5 cm in size. The formation of pseudonodules is explained as selective lithification of carbonate crusts and associated siliceous sponges, followed by fragmentation of already lithified fragments caused by an increase in sedimentary load (Matyszkiewicz 1994, 1997). (d) Early diagenetic dissolution of carbonate within the reduction zone caused by H2S- and NH4-rich pore waters, followed by an ascendent flow of the pore water into the oxidation zone due to compaction, and leading to the precipitation of carbonate around crystallization centers (Gründel and Rösler 1963). Jenkyns (1974) modified this model by explaining the carbonate-rich solution as aragonite dissolution at the sedimentary surface, followed by a descendent flow of calcite-supersaturated pore water, causing calcite precipitation a few centimeters below the sediment surface. The model is supported by the observation of patchy cementation by High-Mg calcite in subrecent carbonate mud in the Ionian Sea and on Bahamian slopes, and is regarded as a possible precursor to nodular fabrics (Müller and Fabricius 1974; Mullins et al. 1980).
Nodular Limestone Fabrics
(e) Local cementation of burrows, partly induced by the organic matter within the burrows (Fürsich 1973; Abed and Schneider 1980). Bioturbation is regarded as a major process in forming hydroplastic pseudonodules (Elmi and Ameur 1984) and distributing nodules (Möller and Kvingan 1988). 3 Formation of nodules through mechanical processes including (a) Early diagenetic down-slope creeping of limestone/marl sequences and tearing up of partially lithified limestone beds (Seyfried 1978). (b) Tectonic processes such as sedimentary boudinage (Varol and Gokter 1994) or shearing of limestone/ clay sequences due to tectonic (Bourbon 1982; Burkhard 1990; Schweigl and Neubauer 1997). No single mechanism or sequence of mechanisms can explain how all nodular limestones are formed. Therefore, each case should be considered individually. Erosion and redeposition of hardened nodules formed only a few decimeters below the sediment-water interface seem to be a reliable interpretation for the genesis of many ancient nodular limestones. Many authors describe nodular fabrics in limestones as resulting from early diagenetic concretionary cementation. However, submarine dissolution on the seafloor, a combination of bioturbation, cementation and redeposition as well as pressure-solution should also be taken into account as possible factors. Significance of nodular limestone fabrics Evaluation of paleowater depth and paleotopography: Nodular limestones are bound to particular environments which developed in specific episodes of rifting of furrows and basins and in specific paleogeographic situations (Wald et al. 1983; Elmi and Ameur 1984). Settings include flat seamounts, steep slopes and flat bottoms of basins and narrow furrows. Möller and Kvingan (1988) drew attention to the relationships between paleotopography, paleodepth, depositional environment and the potential of calcareous sediments for forming nodular limestones during diagenesis. Many Paleozoic and Mesozoic nodular limestones were formed on subtidal deep swells. Spectacular examples are the nodular limestones of the ‘Ammonitico rosso’ facies in the Jurassic of the Alpine-Mediterranean region (see Sect. 15.8.2.2, Pl. 139). Most nodular limestones occur in deeper-marine depositional settings. Some nodular fabrics, produced in connection with burrowing, however, are formed in
Basics: Depositional and Diagenetic Fabrics
shallow subtidal settings (Abed and Schneider 1980, 1982) or on shallow pelagic ridges (Wendt et al. 1984). Sedimentation rates: Studies of nodular fabrics, bioturbation patterns and the non-carbonate content of the limestones assist in evaluating sedimentation rates, which should generally be low regardless of the genetic model favored. The time involved in the forming nodular limestone sections is long, as indicated by the common occurrence of condensed faunas representing several biozones (see Sect. 5.2). Basics: Depositional and diagenetic fabrics Geopetal structures Aissaoui, D.M., Purser, B.H. (1983): Nature and origins of internal sediment in Jurassic limestones of Burgundy (France) and Fnoud (Algeria). – Sedimentology, 30, 273289 Futterer, E. (1978): Studien über Einregelung, Anlagerung und Einbettung biogener Hartteile im Strömungskanal. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 156, 87-131 Rhoads, D.C., Stanley, D.J. (1965): Biogenic gradded bedding. – J. Sed. Petrol., 35, 956-963 Salazar-Jiménez, A., Frey, R.W. (1982): Concavity orientations of bivalve shells in estuarine and nearshore shelf sediments, Georgia. – J. Sed. Petrol., 52, 565-586 Sander, B. (1936): Beiträge zur Kenntnis der Anlagerungsgefüge (Rhythmische Kalke und Dolomite aus der Trias). – Tschermaks Mineralogisch-Petrographische Mitteilungen, 48, 28-139, 141-209 Sander, B. (1951): Contributions to the study of depositional fabric. Translated by R.B. Knopf. – 207 pp., Tulsa (American Association of Petroleum Geologists) Shrock, R.P. (1948): Sequence in layered rocks. A study of features and structures useful for determining top and bottom or oder of succession in bedded and tabular rock bodies. – 507 pp., New York (McGraw Hill) Wanless, H.R. (1981): Fining-upwards sedimentary sequences generated in seagrass beds. – J. Sed. Petrol., 51, 445-452 Wieczorek, J. (1979): Geopetal structures as indicators of top and bottom. – Annales de la Societé géologique de Pologne, 49, 215-221 Further reading: K105 Biofabrics and grain orientation Aigner, T. (1985): Biofabrics as dynamic indicators in nummulite accumulations. – J. Sed. Petrol., 55, 131-134 Hladil, J., Cejchan, P., Gabasova, A., Tideaborsky, Z., Hladikova, J. (1996): Sedimentology and orientation of tentaculite shells in turbidite mudstone to packstone: Lower Devonian, Barrandian, Bohemia. – J. Sed. Petrol., 66, 888-899 Kidwell, S.M., Holland, S. (1991): Field description of coarse bioclastic fabrics. – Palaios, 6, 426-434 Lindholm, R.C. (1980): Intraclast orientation in Cambro-Ordovician limestones in western maryland. – J. Sed. Petrol., 50, 1205-1212
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Potter, P.E., Pettijohn, F.J. (1977): Paleocurrents and basin analysis. Second edition – 425 pp., Berlin (Springer) Racey, A. (1994): Paleoenvironmental significance of larger foraminiferal biofabrics from the Middle East Seeb Limestone Formation of Oman: implications for petroleum exploration. – In: Al-Husseini, M.I. (ed.): Geo’94. The Middle East petroleum geosciences, volume II, Selected Middle East Papers from the Middle East Geoscience Conference, Bahrain. – 793-810, Manama, Bahrain (Gulf PetroLink) Schwarzacher, W. (1963): Orientation of crinoids by current action. – J. Sed. Petrol., 33, 580-586 Seilacher, A. (1973): Biostratonomy: the sedimentology of biologically standardised particles. – In: Ginsburg, R.N. (ed.): Evolving concepts in sedimentology. – 159-177, Baltimore (Johns Hopkins University Press) Whisonant, R.C. (1987): Paleocurrent and petrographic analysis of imbricate intraclasts in shallow-marine carbonates, Upper Cambrian, Southwestern Virginia. – J. Sed. Petrol., 57, 983-994 Further reading: K099; K096 and K097 (Bedding and lamination fabrics) Bioturbation fabrics Bottjer, D.J., Droser, M.L. (1994): The history of Phanerozoic bioturbation. – In: Donovan, S.K. (ed.): Palaeobiology of trace fossils. – 155-176, 5 Figs., Chichester (Wiley) Droser, M.L., Bottjer, D.J. (1986): A semiquantitative field classification of ichnofabric. – J. Sed. Petrol., 56, 558-559 Further reading: K098 Fenestral fabrics Bain, R.J., Kindler, P. (1994): Irregular fenestrae in Bahamian eolianites: a rainstorm-induced origin. – J. Sed. Research, A 64, 140-146 Bathurst, R.G.C. (1959): The cavernous structure of some Mississippian Stromatactis reefs in Lancashire, England. – Journal of Geology, 67, 506-521 Bathurst, R.G.C. (1980): Stromatactis - origin related to submarine-cemented crusts in Paleozoic mud mounds. – Geology, 8, 131-134 Bathurst, R.G.C. (1982): Genesis of stromatactis cavities between submarine crusts in Paleozoic carbonate mud buildups. – J. Geol. Soc. London, 139, 165-18 Bourque, P.-A., Boulvain, F. (1993): A model for the origin and petrogenesis of Red Stromatactis Limestone of Palaeozoic mud mounds. – J. Sed. Petrol., 63, 607-619 Bourque, P.-A., Gignac, H.(1983): Sponge-constructed Stromatactis mud-mounds, Silurian of Gaspé, Quebec. – J. Sed. Petrol., 53, 521-532 Bourque, P.-A., Raymond, L. (1994): Diagenetic alteration of early marine cements of Upper Silurian stromatactis. – Sedimentology, 41, 255-269 Deelman, J.C. (1972): On mechanisms causing birdseye structures. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1972, 582-595 Dupont, E. (1881): Sur l’origine des calcaires Dévoniens en Belgique. – Bull. Acad. Roy. Sci. Belgique, (3) 2, 264280 Dupont, E.(1882): Les iles coralliennes de Roly et de Philippeville. – Bull. Mus. Roy. Hist. Nat. Bruxelles, 1, 89-160 Emery, K.O. (1945): Entrapment of air in beach sand. – J. Sed. Petrol., 15, 39-49
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Fischer, A.G. (1964): The Lofer cyclothems of the Alpine Triassic. – Kansas Geol. Surv. Bull., 169, 107-149 Flajs, G., Hüssner, H. (1993): A microbial model for the Lower Devonian Stromatactis mud mounds of the Montagne Noire (France). – Facies, 29, 179-194 Grover, G., Read, J.F. (1978): Fenestral and associated vadose diagenetic fabrics of tidal flat carbonates, Middle Ordovician New Market Limestone, southwestern Virginia. – J. Sed. Petrol., 48, 453-473 Ham, W.E. (1954): Algal origin of the ‘Birdseyes’ limestone in the McLish Formation. – Oklahoma Acad. Sci. Proceed., 33, 200-203 Heckel, P.H. (1972): Possible inorganic origin for Stromatactis in calcilutite mounds in the Tully Limestone, Devonian of New York. – J. Sed. Petrol., 42, 7-18 Matyszkiewicz, J. (1993): Genesis of Stromatactis in an Upper Jurassic carbonate buildup (Mlynka, Cracow region, Southern Poland): Internal reworking and erosion of organic growth cavities. – Facies, 28, 87-96 Mazzullo, S.J., Birdwell, B.A. (1989): Syngenetic formation of grainstones and pisolites from fenestral carbonates in peritidal settings. – J. Sed. Petrol., 59, 605-611 Monty, C.L.V. (1995): The rise and nature of carbonate mudmounds: an introductionary actualistic approach. – In: Monty, C.L.V., Bosence, D.W.J., Bridges, P.H., Pratt, B.R. (eds.): Carbonate mud-mounds: Their origin and evolution. – Spec. Publ. Int. Ass. Sed., 23, 11-48 Pratt, B.R. (1982): Stromatolitic framework of carbonate mud mounds. – J. Sed. Petrol., 52, 1203-1227 Pratt, B.R. (1995): The origin, biota and evolution of deepwater mud mounds. – In: Monty, C.L.V., Bosence, D.W.J., Bridges, P.H., Pratt, B.R. (eds.): Carbonate mud-mounds: their origin and evolution. – Spec. Publ. Int. Ass. Sed., 23, 49-123 Semeniuk, V. (1971): Subaerial leaching in the limestones of the Bowan Park Group (Ordovician) of Central Western New South Wales. – J. Sed. Petrol., 41, 939-950 Shinn, E.A. (1968a): Burrowing in recent lime sediments of Florida and the Bahamas. – J. Palaeont., 42, 879-894, Pls. 109-112 Shinn, E.A. (1968b): Practical significance of birdseyes structures in carbonate rocks. – J. Sed. Petrol., 38, 215-223 Shinn, E.A. (1983): Birdseyes, fenestrae, shrinkage pores, and loferites: A reevaluation. – J. Sed. Petrol., 53, 619628, 5 Figs., Tulsa Tebbutt, G.E., Conley, C.D., Boyd, D.W. (1965): Lithogenesis of a carbonate rock fabric. – Contrib. Geol., 4, 1-13 Tsien, H.H. (1985): Origin of stromatactis – a replacement of colonial microbial accretion. – In: Toomey, D.F., Nitecki, M.H. (eds.): Paleoalgology. – 274-289, Berlin (Springer) Wallace, M.W. (1987): The role of internal erosion and sedimentation in the formation of stromatactis mudstones associated lithologies. – J. Sed. Petrol., 57, 695-700 Wolf, K.H. (1965): Littoral environments, indicated by openspace structures in algal limestones. – Palaeogeogr. Palaeoclimat., Paleoecol., 1, 183-223 Further reading: K101 Nodular Fabrics Abed, A.M., Schneider, W. (1982): The Cenomanian nodular limestone member of Jordan – from subtidal to supratidal environments. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1982, 513-522 Bathurst, R.G. (1987): Diagenetically enhanced bedding in
Basics: Nodular Fabrics
argillaceous platform limestones: stratified cementation and selective compaction. – Sedimentology, 34, 749-778 Böhm, F. (1992): Mikrofazies und Ablagerungsmilieu des Lias und Dogger der Nordöstlichen Kalkalpen. – Erlanger geologische Abhandlungen, 121, 57-217 Burkhard, M. (1990): Ductile deformation mechanisms in micritic limestones naturally deformed at low temperatures (150-350 °C). – In: Knipe, R.J., Rutter, E.H. (eds.): Deformation mechanisms, rheology and tectonics. – Spec. Publ., Geol. Soc. London, 241-257 Bourbon, M. (1982): La genèse des calcaires noduleux en Biaconnais (Haute-Alpes): une conséquence de l’instabilité tectonique de ce domaine au Malm inférieur. – Mém. Géol. Univ. Dijon, 7, 129-138 Cecca, F., Fourcade, E., Azema, J. (1992): The disappearance of the ‘Ammonitico Rosso’. – Palaeogeography. Palaeoclimatology., Paleoecology., 99, 55-70 Dvorak, J. (1972): Shallow-water character of nodular limestones and their paleogeographic interpretation. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1972, 509-511 Elmi, S., Ameur, M. (1984): Quelques environments des facies noduleux mésogéen. – Geol. Romana, 23, 13-22 Fürsich, F.T. (1973): Thalassionides and the origin of nodular limestone in the Corallian Bed (Upper Jurassic) of southern England. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1973, 136-156 Gründel, J., Rösler, H.J. (1963): Zur Entstehung der oberdevonischen Kalkknollengesteine Thüringens. – Geologie, 16, 1009-1038 Hollmann, R. (1962): Subsolutions-Fragmente (Zur Biostratinomie der Ammonoidea im Malm des Monte Baldo/Norditalien). – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 119, 22-82 Jenkyns, H.C. (1974): Origin of the red nodular limestones (Ammonitico Rosso, Knollenkalke) on the Mediterranean Jurassic: a diagenetic model. – In: Hsü, K.J., Jenkyns, H.C. (eds.): Pelagic sediments: on land and under the sea. – Spec. Publ. Int. Ass. Sedimentologists, 1, 249-271 Kukal, Z. (1975): On the origin of nodular limestones. – Casopis pro mineralogii a geologii, 20, 359-368 Jeans, C.U. (1980): Early submarine lithification in the Red Chalk of Eastern England. A bacterial control model and its implications. – Proc. Yorkshire geol. Soc., 43, 81-157 Kennedy, W.J., Garrison, R.E. (1975): Morphology and genesis of nodular chalks and hardgrounds in the Upper Cretaceous of southern England. – Sedimentology, 22, 311386 Langbein, R., Meinel, G. (1985): Zur Petrologie des Thüringer Tentakulitenknollenkalkes (Devon). – Hallesches Jahrbuch der Geowissenschaften, 10, 55-69 Matyszkiewicz, J. (1994): Remarks on the deposition and diagenesis of pseudonodular limestones in the Cracow area (Oxfordian, southern Poland).– Berliner gewissenschaftliche Abhandlungen, E13, 419-439 Möller, N.K., Kvingan, K. (1988): The genesis of nodular limestones in the Ordovician and Silurian of the Oslo region (Norway). – Sedimentology, 35, 405-420 Müller, J., Fabricius, F. (1974): Magnesian-calcite nodules in the Ionian deep sea: an actualistic model for the formation of some nodular limestones. – In: Hsü, K.J., Jenkyns, H.C. (eds.): Pelagic sediments: on land and under the sea. – Spec. Publ. Int. Ass. Sedimentologists, 1, 235-247 Mullins, H.T., Neumann, A.C., Wilber, R.J., Boardman, M.R.
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Discontinuity Surfaces
(1980): Nodular carbonate sediment on Bahamian slopes: possible precursors to nodular limestones. – J. Sed. Petrol., 50, 117-131 Seyfried, H. (1978): Der subbetische Jura von Murcia (Südost-Spanien). – Geologisches Jahrbuch, Reihe B, 29, 3201 Tucker, M.E. (1974): Sedimentology of Paleozoic pelagic limestones: the Devonian Griotte (Southern France) and Cephalopodenkalk (Germany). – In: Hsü, K.J., Jenkyns, H.C. (eds.): Pelagic sediments: on land and under the sea. – Spec. Publ. Int. Ass. Sedimentologists, 1, 71-92 Wald, S., Kurze, M., Wienholz, R. (1983): Ausbildung und Genese oberdevonischer Kalkknollengesteine im Süden der DDR. – Zeitschrift der geologischen Wissenschaften, 11, 27-39 Wendt, J., Aigner, T., Neugebauer, J. (1984): Cephalopod limestone deposition on a shallow pelagic ridge: the Tafilalt Platform (upper Devonian, eastern Anti-Atlas, Morocco). – Sedimentology, 31, 601-625 Further reading: K104
5.2 Discontinuity Surfaces: From Microfacies to Sequence Stratigraphy Sedimentation is inherently a discontinuous process. A gradual change in environmental conditions may be accompanied by a continuous reaction of the depositional system, but any abrupt change or the passing of a threshold lead to a discontinuity in sedimentation. Discontinuities are commonly marked by a surface in stratigraphic sections and/or by distinct level of facies changes. All surfaces indicating a break in sedimentation and/ or non-deposition in submarine and subaerial environments are called discontinuity surfaces (Heim 1924, 1934). The use of this term is purely descriptive in being independent of the process causing the formation of the surface and the duration of the break. Consequently, all stratigraphic gaps are included in this broad definition, ranging from millimeter- and centimeterscale breaks in sedimentation revealed by microfacies analysis, to small breaks on the scale of limestone beds commonly marked as bedding planes, to major stratigraphic unconformities which may be accompanied by prolonged subaerial exposure. Recognizing discontinuities and unconformities should be a major goal in microfacies studies. Unconformities represent the key element in sequence stratigraphy. Because sequence boundaries are defined as ‘observable discordances in a given stratigraphic section that show evidence of erosion or non-deposition with obvious stratal terminations...’ (Mitchum in Vail et al., 1977), microfacies criteria pointing to breaks in sedimentation (e.g. discontinuity surfaces, condensed deposits, hardgrounds, karst features et al.) are of spe-
cific interest in refining carbonate sequence stratigraphy and the discussion of time involved in stratigraphic breaks. Small-scale and short-lived discontinuities often reflect large-scale variations of relative sea level. Terminology Various names are in use for defining sedimentary breaks and associated substrate types (Box 5.4).
5.2.1 Classification of Discontinuities Practical classifications of the common discontinuity types occurring in carbonate sequences were proposed by Clari et al. (1995) for analyzing shallow- and deepwater carbonates and Hillgaertner (1998) for studying carbonate platforms. The latter classification differentiates four discontinuity surfaces taking into account the importance of allogenic forcing of environmental changes: (1) Exposure surfaces resulting from subaerial exposure, commonly indicated by an overprinting of subtidal carbonates due to a general lowering of sea level. (2) Condensation surfaces comprising all discontinuities related to stratigraphic condensation in subtidal environments, and predominantly dependent on controls by relative sea-level changes and changes in accommodation. (3) Erosion surfaces including discontinuities that show evidence of subaqueous erosion. (4) Change surfaces describing all surfaces indicating changes in facies and/or texture. In most cases small-scale erosion or facies changes are related to locally restricted depositional processes. But allogenic forcing associated with changes in climate and/or relative sea level can trigger changes in wave base and hydrodynamic energy producing erosion surfaces or surfaces characterized by an abrupt change from carbonate to siliciclast-dominated sedimentation.
5.2.2 Major Criteria of Discontinuities Common criteria used to characterize subaerial and/or submarine discontinuities in carbonate rocks include geometry, lateral extent, surface morphology, biological activity, mineralization, facies contrast, early diagenetic contrast, and biostratigraphy (Hillgaertner 1998). The evaluation of these criteria requires the study of the rocks below and above the discontinuity, a thorough analysis of the discontinuity surface in geological
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Discontinuity Surfaces
Box 5.4. Glossary of terms used in the context of the discussion of discontinuity surfaces. Condensation: Sedimentation lag characterized by stagnant to low sedimentation rates over long periods of time (Heim 1924) ranging from thousands to several millions of years. Condensed section: Thin marine stratigraphic units consisting of pelagic and hemipelagic sediments characterized by very low sedimentation rates (Loutit et al. 1988). Corrosion surface: A bedding surface in carbonate rocks that exhibits evidence of lithification, followed by partial solution, of the underlying bed prior to the deposition of the overlying bed (Weiss 1954). Most criteria of corrosion surfaces correspond to those of hardgrounds, but the term underlines the greater importance of etching and dissolution. Dereption: Physical action of bottom currents leading to submarine erosion and transport of seafloor material (Heim 1934). Diastem: A relatively short interruption in sedimentation involving only a brief interval of time with little or no erosion before deposition is resumed (Barrell 1917). The depositional break represents less time than a paraconformity. Disconformity: Disconformities are stratigraphic bounding surfaces which possess a gap in time documentation despite the fact that the beds below and above the discontinuity are parallel. The discontinuity can result from erosion, abrasion, denudation, deflation, dissolution, corrosion, or non-deposition. Discontinuity: A zone at the sea bottom, usually a few cm thick, the sediment of which is lithified to form a hardened surface, often encrusted, discolored, case-hardened, bored and solution ridden (Heim 1934; Clari et al. 1995). It implies a gap in sedimentation and may be preserved stratigraphically as an unconformity. Today the term is not restricted to submarine settings, but also includes subaerial unconformities (Hillgaertner 1998). Emersion: Subaerial exposure of marine areas due to sea-level fall or/and uplift. Firmground: A stiff, unconsolidated sedimentary substrate that can be easily broken and shows no signs of compaction (Ekdale 1985). Often representing the initial stages of hardground formation (Ruffell and Wach 1998). Firmgrounds typically contain a network of burrows (Bromley 1990). Hardground: Discontinuous surfaces of synsedimentary lithification, having existed as hard sea floor prior to the deposition of the overlying sediment. Implies a gap in sedimentation and may be preserved stratigraphically as an unconformity. Hiatus: A break or interruption in the continuity of the stratigraphic record and a lapse in time such as the time interval not represented by rocks or an unconformity. Hiatus concretions: Early diagenetic concretions that were bored and encrusted subsequent to winnowing of adjacent sediments and exhumation. Hiatus concretions represent discontinuous hardgrounds, and indicate short- and longlived breaks in sedimentation (Voigt 1968). Omission: Non-deposition of sediment related to a significant and long-lasting break in sedimentation (Heim 1924), leading to the formation of hardgrounds. Paraconformity: Obscure unconformity in which no or only subtile erosional features are discernible and/or in which the contact over large areas appears to be a bedding plane (Dunbar and Rodgers 1957; Longman et al. 1983). Rockground: A sea ground along which well-lithified sediments are exposed. The term was originally proposed in contrast to hardground for ‘a rocky sea floor, whose lithification was not synsedimentary, representing older layers of rock which through erosion and submersion formed the sea floor’ (Fürsich 1979), but is now used as a term describing surfaces which regardless of their origin were hardened before deposition of their immediately overlying sediment (Clari et al. 1995). Stratal discontinuities: Physical surfaces that either separate strata with different angularity or separate parallel strata where a significant hiatus is present (Doglioni et al. 1990). Subsolution: Carbonate dissolution at the bottom of the oceans (Heim 1924), commonly at great depths and in coolwater regions. Dissolution processes are interrelated with cementation, organic erosion and winnowing of sediment during times of non-deposition (Hollmann 1964). Unconformity: A surface of erosion or non-deposition that separates younger strata from older rocks and represents a significant hiatus (Mitchum 1977). In the context of sequence stratigraphy the definition is restricted to a surface separating younger from older strata, along which there is evidence of subaerial truncation (and, in some areas, correlative submarine erosion) or subaerial exposure, with a significant hiatus indicated (Wagoner 1988).
sections, a combined approach of field observations and sedimentological plus paleontological data including microfacies. Geometry: Many discontinuities exhibit an angular relationship between underlying and overlying strata, caused by the erosional truncation of sedimentary and diagenetic features or entire stratigraphic units. Discontinuity surfaces with non-angular, conformable re-
lationships need microfacies studies, geochemical and/ or biostratigraphical evidence for their interpretation. Lateral extent: Single surfaces with wide lateral continuity (>1 km) indicate environmental changes of at least regional importance. Low lateral continuity (several meters only) is commonly related to locally limited depositional processes. Physical correlation of single open-marine discontinuity surfaces over a dis-
Discontinuities
tance of tens of kilometers or more may point to major oceanographic controls (Di Stefano et al. 2002). Surface morphology: Flat, sharply-cut surfaces in shallow-marine environments are commonly caused by high-energy abrasion of lithified sediment. Irregular wavy and crinkled surfaces may indicate minor erosion or the former existence of microbial mats. Irregular surfaces also can be caused by differential compaction of strata differing in composition and texture. Rough surfaces can originate from burrowing and bioerosion. Stylolitization superimposed on preexisting discontinuity surfaces may lead to jagged surfaces. Biological activity: The distribution of burrowing, boring and encrusting organisms is crucial for understanding changes in sedimentation rates and resulting substrate types. The intensity of bioturbation and the types of trace fossils below and above the discontinuity surface assist greatly in differentiating pre-omission, omission, and post-omission phases (Bromley 1975). Lithification of marine discontinuity surfaces is indicated by boring organisms that cut sharply through fabrics and different microfacies types of underlying rocks. Encrusting microfossils attached to hardgrounds indicate retardation or breaks in sediment input. Mineralization: Discontinuity surfaces showing in situ crusts of iron and manganese oxides and phosphates as well as authigenic glauconite (Pl. 22/2) indicate breaks in sedimentation, commonly in deep subtidal and bathyal environments. In contrast, aluminium-iron oxide crusts are found to be the result of paleosol formation (terra rossa) and alteration during subsequent marine flooding (Wright 1994). Penetrative staining of strata underlying discontinuity surfaces by iron oxides can indicate oxidation due to pedogenesis and karstification. Facies contrast: A sharp change in microfacies across a surface must not necessarily represent an important break in sedimentation. However, a drastic environmental change can be inferred from a superposition contradicting Walther’s law (Clari et al. 1995), e.g. the superposition of limestones exhibiting microfacies associations typical of platform and reef carbonates by limestones showing evidence of pelagic sedimentation. Changes in grain sizes (reflecting changes in water energy) and in the lithologic composition (e.g. siliciclastics versus carbonates) may provide other clues to the discontinuity character of surfaces. Diagenetic contrasts: Diagenetic evidence of discontinuities is given by sharp differences in the diagenetic style at the discontinuity, e.g. vadose diagenesis or early diagenetic changes due to meteoric waters in the underlying and contrasting marine phreatic diagen-
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esis in the overlying strata. Cement stratigraphy, different compaction features in the overlying and underlying strata as well as stable isotopes and trace elements offer other clues to the existence of exposure surfaces or breaks in the diagenetic style. Isotope and trace element data assist greatly in recognizing submarine lithification of hardgrounds, identifying unconformities and the genetically interpretating subaerial and submarine discontinuities (Marshall and Ashton 1980; Videtich and Matthews 1980). Biostratigraphy: Generally biostratigraphic and chronostratigraphic data are the only means for estimating the time involved in discontinuities. In basinal sections with rather monotonous sedimentation patterns, fossils can even be only means of identifying stratigraphical discontinuities. The main limitation, however, is time resolution, which is commonly too low for the majority of discontinuity surfaces occurring in carbonate sequences. Most discontinuities in shallow-marine platform carbonates are below biostratigraphic resolution. Better results have been obtained in the context of studying deep-marine carbonates.
5.2.3 Microfacies Criteria and Significance of Exposure Surfaces Detecting exposure surfaces in limestones by means of microfacies includes the following criteria: • Conspicuous differences in facies types reflected by differences in rock color, lithologic composition, texture and fossil content below and above the exposure surface, • Meteoric-vadose diagenesis at bounding surfaces, characterized by vadose cements in the underlying rocks and contrasting marine phreatic cements in the overlying rocks. Vadose zones extend downwards in the strata for a few tens of centimeters only. The surface can be accentuated by stylolitization. Stable isotope trends across exposure surfaces exhibit significant shifts (Sect. 13.2.4), • Dedolomitized horizons: The alteration of dolomite to calcite takes place in several diagenetic settings (see Sect. 7.8.3) but is a common feature of near-surface vadose zones (Kenny 1992). Open dolomolds may indicate subaerial exposure and dissolution by karst waters (Braun and Friedman 1970), • Microkarst caused by dissolution of CO2-enriched meteoric waters is indicated by sharp-cut and intensely stained surfaces displaying a microrelief with millimeter- to centimeter-sized cavities (Sect. 15.2.1). The ab-
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Hardgrounds
sence of larger paleokarst solution cavities is explained by the formation of epikarst underneath a soil cover during seasonally humid climates (Wright 1994; Mylroie and Carew 1995) and limited duration of subaerial exposure (d’Argenio et al. 1997). Other karst features are the gravitational infill of cavities by stained and reworked lithoclasts and clayey calcareous silt in association with micrite crusts and stalactite cements, • Pedogenic features (Pl. 128/4, Pl. 141/9, 10): Dissolution processes and rhizoturbation during exposure as well as erosion and reworking of the soil during the subsequent marine flooding, can produce a brecciated to pebbly appearance of exposure surfaces. Other criteria are microscopic root molds, soil pisoids and calcrete crusts (Sect. 15.1.1). The absence of a vertical zonation of pedogenic fabrics at these surfaces points to generally short-lived exposures. All these methods have limits and pitfalls and must be combined with a detailed analysis of sequence stratigraphy, biostratigraphy, facies and diagenetic history in a regional context (Saller et al. 1994; Fouke et al. 1996).
• Caverns, breccia and vugs can form in deep subsurface because of dissolution by basinal fluids or the controls of shales overlying carbonates.
Practical significance of exposure surfaces and subaerial unconformities: Reservoir rocks: Identifying and predicting subaerial exposure surfaces is essential for understanding the modification of porosity beneath sequence-bounding unconformities and parasequence disconformities on carbonate platforms, the reservoir degradation below subaerial unconformities, and the development of permeability patterns (Read and Horbury 1993; Saller et al. 1994; Budd et al. 1995; Wagner et al. 1995). Controls of subaerial unconformities on porosity and permeability include the following points: • Carbonate diagenesis during subaerial exposure rearranges the pore network, but does not necessarily increase total porosity. In many cases porosity is reduced under unconformity, but remarkable exceptions exist, • Permeability is more strongly changed by subaerial discontinuities than porosity, and may increase or decrease, • Pore systems evolve with the time of subaerial exposure. Shorter periods (10 000–400 000 yr) are often associated with higher porosity than long intervals of subaerial exposure (1–20 million yr), • Many carbonates subjected to subaerial exposure have little or no porosity in the deep surface due to compaction, cementation and stratal collapse reducing burial porosity. But unconformity-related diagenesis may also enhance reservoir porosity by creating pore systems that are resistant to deeper burial compaction,
Stratigraphically reduced marine carbonates originate from very low sedimentation and alternating deposition and erosion, leading to the formation of condensed sequences. Modern oceanic equivalents have been identified on seamounts. Breaks in the sedimentation caused by non-deposition result in the formation of hardgrounds both in shallow-marine settings and deep-marine oceanic environments. In many ancient open-marine carbonates hardgrounds and condensed sequences are intimately associated.
Ore deposits: Emersion surface-hosted ore deposits associated with paleokarst are known from the Paleozoic and Mesozoic (Parnell et al. 1990; Scott et al. 1993; Ruffell et al. 1998). Pb-Zn-Fe sulfide mineralizations are common below and/or along and above subaerial exposure surfaces of carbonate platforms. One such example, interesting in economic terms, is the ‘Metallifero’ in the southwestern part of Sardinia, a Cambrian carbonate complex, unconformably overlain by Ordovician conglomerates (Fanni et al. 1981; Bechstädt and Boni 1994). Other examples are emersion-bound mineralizations at Devonian-Carboniferous and Triassic disconformities in the Alps (Assereto et al. 1976; Hentschel and Kern 1992).
5.2.4 Microfacies Criteria and Significance of Condensation Surfaces and Hardgrounds
5.2.4.1 Hardgrounds Hardgrounds are centimeter-sized discontinuous surfaces of synsedimentary lithification, having existed as hardened sea floor prior to the deposition of the overlying sediment. They are related to a combination of non-deposition or low sedimentation rates, and condensation. Hardgrounds denote a submarine sediment surface that became lithified in the ambient depositional environment before the next sediment layer was deposited (Demicco and Hardie 1994). Hardground surfaces are the result of submarine cementation by aragonite and magnesian calcite precipitation directly from seawater circulating through the uppermost few centimeters or tens of centimeters of a porous sandy sea bottom. The basic requirement is a longer-lasting stability of the sediment-water interface. The time involved
Ancient Hardgrounds
in the formation of hardgrounds is unclear, Shinn (1969) reported antique Greek pottery incorporated into hardground of the Persian Gulf (a few thousand years). Hardgrounds are common but not confined to deeper-marine settings. They also occur in the terrestrial onshore and estuarine zone (tropical lateritic hardpans) throughout supra- and intertidal (beachrocks) and various marine regions. Marine settings include shallow subtidal environments (Pl. 127/4), off-platform pelagic settings, shallow slopes (Mullins et al. 1980), and deep basins. Holocene hardgrounds have been described from the Bahamas (Dravies 1979), the Persian Gulf (Shinn 1969), the Red Sea, and the Mediterranean Sea (Aghib et al. 1991; Bromley and Allouc 1992). The majority lies less than a few tens of meters under water. In the Bahamas and in the Persian Gulf the development of hardgrounds is probably related to surface water saturation relative to calcium carbonate. In these shallow-water environments, cessation of sedimentation and development of hardgrounds can form in the space of a few months. Some of the hardgrounds, found in oceanic settings, occur in great depths, sometimes on oceanic ridges and guyots (Pl. 23/6-8), between several hundreds of meters to more than 3000 meters (Fischer and Garrison 1967; Bernoulli and McKenzie 1981; Malfait and Van Andel 1980).
Recognizing ancient hardgrounds Fig. 5.9 summarizes some of the most common features of hardgrounds. Important criteria are: Contacts: Sharp upper contact and diffuse lower contact. Grains and cement crystals may be truncated along the upper contacts (Pl. 23/2). Upper surface sometimes has a thin micrite rind. Some hardground surfaces are covered by micritic microbial crusts. The overlying sediment may infill cavities (Pl. 22/5), borings or undercut recesses in the hardground. Surfaces: The upper contact is characterized by smooth or irregular surfaces. Many surfaces exhibit irregular protuberances separated by pits and grooves (Pl. 23/3). Smooth planar surfaces are formed by abrasion and are more common in shallow-marine highenergy environments. Angular surfaces related to solution processes are more common in deep-marine shelf and basinal settings. The upper surface of shelf hardgrounds may be buckled into tepee-like expansion ridges or may show signs of corrosion. The lack of encrustations on these surfaces indicates relatively short omission phases and high-energy conditions. Mineralization: Uppermost layers may be mineralized in the form of crusts or as impregnations (Pl. 23/
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Fig. 5.9. Common criteria of carbonate hardgrounds. Modified after James and Choquette (1983).
3, 6). The composition, form and intensity of mineralization depend on the environment: Shelf hardgrounds are usually mineralized by glauconite (occurring within grains, burrows - Pl. 22/4, borings and cracks), calcium phosphate salts (forming phosphate cements as well as phosphatic stromatolites - Pomoni-Papaioannou 1994) and iron hydroxides (goethite, limonite), whereas pelagic hardgrounds in addition also contain a high percentage of manganese oxides (Fürsich 1979; Lindström 1979; Marshall and Ashton 1980). In some hardgrounds, calcium carbonate has been replaced by baryte (Lindström 1979). Ferruginous crusts covering hardgrounds frequently consist of irregular laminae of iron hydroxide, occasionally forming a cauliflower structure (Pl. 23/5), alternating with laminae of calcareous skeletons of encrusting organisms (e.g. foraminifera - Pl. 23/3, 4, serpulids). Current interpretations favor a strong microbial control on the formation of laminated Fe-Mn crusts. In-place dissolution within microbial carbonates with thin limonitic laminae is explained by a process in which iron crusts acted as a cathode, sea water as an electrolyte, and the underlying carbonate as an anode (Videla 1989; Reitner et al. 1995). The existence of limonite crusts on hardground surfaces is often regarded as indication of relatively long omission phases and low-energy hydrodynamic conditions. Biogenic encrustations: The upper surfaces of hardgrounds are frequently encrusted by organisms who require a hard substrate to settle. Encrusting bryozoans, brachiopods and oysters occur also on the lower surface of hardgrounds. Differences in encrustation indicate the polarization of the fauna according to light
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intensity, rates of sedimentation and turbulence levels. Common encrusters restricted to hard substrates including hardgrounds are encrusting calcareous algae, foraminifera - Pl. 23/4, bryozoans, corals, serpulid polychaetes, cirripeds, and some brachiopods. Note, that some encrusting organisms occurring on hard substrates can also settle on firm substrates (e.g. some oysters, crinoids and brachiopods). Bryozoans have proved to be of particular value for differentiating Cretaceous hardgrounds (Voigt 1968). Some authors consider the existence of only thin biogenic crusts as an indication of relatively short omission periods. Borings: Macro- and microborings (Pl. 52) are common along the upper surface of many hardgrounds (Pl. 22/4). Borings provide sedimentological information about the consistency of the ancient sea floor as well as paleoecological information about environmental and depositional conditions (Ekdale 1985; Bromley
Discontinuities and Unconformities
1994; Ekdale and Bromley 2001). Borers restricted to hardgrounds and other hard substrates (e.g. rock cobbles, shells, coral colonies) include some bivalves, phoronids, sponges, algae and fungi. Several animals can penetrate both firm and cemented substrates (e.g. bivalves and polychaete worms). Burrowing: The strata underlying and overlying the hardground are characterized by specific associations of trace fossils indicating the changes in substrate types. Changes from firmground to hardground stages are reflected by the temporal relationship of bioturbation, boring and encrustation. Clasts of cemented limestones: Lithoclasts similar in composition to the hardground occur above the upper surface, along the discontinuities or as infill of depressions. The clasts have sharp edges and are often bored, encrusted or coated with Fe/Mn crusts (Pl. 23/ 8). These clasts (e.g. ‘subsolution lithoclasts’, Sect.
Plate 22 Discontinuities and Unconformities Recognizing discontinuities and unconformities is a major goal in microfacies studies, because unconformities represent a key element in sequence stratigraphy and subaerial disconformities have a major impact on the porosity and permeability of carbonate rocks. Stratigraphic discontinuities manifest breaks in sedimentation at unconformities in which bedding planes above and below the break are essentially parallel (‘disconformity’). Interruptions can result from various processes including erosion, significant reduction in sedimentation rates, non-deposition, bioerosion, or dissolution below or above carbonate compensation depths. Disconformity surfaces are often connected with variously colored condensed beds and hardgrounds. The latter denote a submarine sediment surface that became lithified in the ambient depositional environment before the next sediment layer was deposited. Hardgrounds are the result of submarine cementation by aragonite or Mg-calcite precipitation directly from seawater circulating through the uppermost part of the sea floor. They are formed in various marine settings (e.g. supratidal and intertidal, shallow subtidal, shallow and deep slopes, deep open-marine environments) but also originate in subaerial settings (e.g. on emersion surfaces). Because of the abundant hiati in deposition and rock-stratigraphic record, very thin condensed beds can represent very large time spans (100 000s to millions of years). 1 Hardground (arrows) separating deep-marine quartz-bearing wackestones with filaments and a wackestone characterized by phosphoritized bioclasts and a glauconite-bearing micrite matrix. The discontinuity surface is marked by Fe/Mn mineralization. Deep subtidal shelf limestones. Jurassic (Callovian-Oxfordian): near Pamplona, Spain. 2 Phosphoritized ammonites, indicating solution and replacement of the aragonitic shell by phosphate during times of starved sedimentation rates. Most glauconite grains within the micritic matrix are replaced by Fe-calcite. Deep subtidal shelf limestones. Jurassic (early Oxfordian): Ricla southwest of Zaragoza, Spain. 3 Hardground in a condensed shelf limestone with ferruginous ooids (O). The concentric layers consist of goethite, the nucleus of broken ooids, lithoclasts and quartz grains. Note the encrusting foraminifera (F). Jurassic (Early Oxfordian): near Pamplona, Spain. 4 Hardground. Distinct borings within an erosional relief were filled with glauconitized peloids. Deep shelf. Jurassic (Oxfordian): near Pamplona, Spain. 5 Microfacies about 50 cm below a distinct disconformity separating two rock units. The thin section exhibits a cavity within a foraminiferal limestone. The cavity walls crosscut the matrix and some of the foraminifera (Lepidocyclina, arrows) indicating that the cavity was formed by dissolution below the discontinuity plane, after lithification of the rock. The cavity is filled with laminated mudstone and skeletal wackestone. The uppermost layer contains imbricated Lepidocyclina tests. Tertiary (Brac Formation and Cayman Formation, Oligocene/Miocene): Cayman Brac Island, British West Indies. –> 1–4: Behnke 1981, 5: Jones and Hunter 1994
Plate 22: Discontinuities and Unconformities
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4.2.8) represent the result of a complex intraformational reworking of both biological and erosional origin (Pl. 139/7, 10). Hiatus concretions: Reworked, encrusted and bored ‘hiatus concretions’ occurring in carbonate and noncarbonate sediments reflect multiple phases of deposition, burial and excavation and provide insight into substrate conditions and the diagenetic history of hardgrounds (Kennedy and Klinger 1972; Fürsich et al. 1992). Small-scale sedimentary intercalations: Sediments with marine fossils are interlayered with cements in cavities beneath the upper surface (Pl. 22/5). Alternation of hardground surfaces and cement: Many hardgrounds are characterized by multiple surfaces spaced millimeters to tens of millimeters apart, and a high volume of carbonate cement causing laterally continuous cement crusts and layers (Pl. 23/3). Carbonate cements: Most hardgrounds yield LowMg calcite cements occurring as blocky mosaics, syntaxial rim cement and palisade-like fibrous cement, and radial crystal aggregates. Microspar (including ant-egg spar) is abundant. Often calcite crystals are intergrown with glauconite, goethite and pyrite. Microfacies: Microfacies types of underlying and overlying strata may differ in texture and biotic content, or may not. Hardground communities: Clues to lithification stages In addition to cobbles, boulders and invertebrate skeletons hardgrounds are a type of substrate which can be colonized by a wealth of organisms. Here they can thrive as encrusters or as burrowers and borers (Brett 1988). Because the rate of early cementation of substrates imposes considerable constraints upon ecological successions of the burrowing, boring and encrusting biota, the occurrence and abundance of these biota can be used to evaluate lithification stages of hardgrounds (Goldring and Kazmierczak 1974; Gruszczynski 1979). Processes which might trigger ‘non-deposition’ and the formation of hardgrounds Submarine hardgrounds mark hiatuses in submarine deposition, they are omission or condensation surfaces. Controls on non-deposition include • cessation of carbonate mud import through emergence or drowning of platforms supplying it, • extreme dissolution of carbonates and silica at the sea bottom, • intermittent scouring at the sea bottom caused by changing bottom currents, and • depletion of nutrients.
Hardgrounds
Long-lived formation of hardgrounds requires major changes in environmental conditions including • shallowing to the point that seawater circulation becomes too restricted for significant biogenic carbonate production caused by eustatic drop or tectonic uplift, or filling of shallow-water basins, • deepening beyond the effective phototrophic zone due to sea-level rise or tectonic subsidence, or • significant increase in salinity due to barriers. Short-lived formation of hardgrounds can be attributed to • local changes in sediment distribution patterns, • short-term climatic variations affecting the frequency and travel paths of storms. In this case, the subtidal sediments may be dominated by lithoclasts derived from firmgrounds. Box 5.5. Selected case studies of ancient hardgrounds. A comprehensive bibliography of studies related to ancient and modern hardgrounds and other hard surfaces can be found in the Internet (Wilson 1999: Bibliography of lithologic substrate studies). Tertiary: Bromley and Goldring 1992; Carter et al. 1982; Glenn and Kronen 1993; Gupta 1995; Lewis and Ekdale 1992; Malaroda 1962; Mikulas and Pek 1995; Pedley and Bennett 1985 Cretaceous: Bromley 1967, 1968, 1975; Bryan 1992; Carson and Crowley 1993; Kennedy and Garrison 1975; Lewis and Ekdale 1992; Malaroda 1962; Massari and Medizza 1973; Molenaar and Zijlstra 1997; Pomoni-Papaioannou 1994; Pomoni-Papaioannou and Solakius 1991; Rose 1970; Ruffell and Wach 1998; Voigt 1959, 1974, Zitt and Nekvasilova 1996 Jurassic: Behnke 1981; Böhm et al. 1999; Di Stefano et al. 2002; Fels 1995; Fürsich 1971,1979; Fürsich et al. 1992; Garcia-Hernandez et al. 1986; Gruszcynski 1979, 1986; Hallam 1969; Kazmierczak and Pszczolkowski 1968; Krawinkel and Seyfried 1996; Lehner 1992; Lualdi 1986; Marshall and Ashton 1980; Mensink and Mertmann 1984; Palmer and Fürsich 1974; Seyfried 1978; Wendt 1970; Wings 2000 Triassic: Bertling 1999; Kazmierczak and Pszczolkowski 1968; Schlager 1974; Wendt 1970 Carboniferous: Dawson and Carozzi 1983; Loope 1994 Devonian: Baird 1981; Bourrouilh et al. 1998; Tucker 1973; Wendt 1988 Silurian: Bridges 1975; Cherns 1980; Kobluk et al. 1977 Ordovician: Brett and Brookfield 1984; Brett and Liddell 1978; Dronov and Federov 1996; Droser and Bottjer 1989; Ekdale and Bromley 2001; Jaanusson 1961; Lindström 1963, 1979; Nordlund 1989; Palmer and Palmer 1977; Pickerill and Narbonne 1995; Saadre 1993; Wilkinson et al. 1982; Wilson 1985 Cambrian: Brett et al. 1983; Chow and James 1992; Hessland 1955; Palmer 1982
Condensation
Development of hardground formation over time Box 5.5 indicates higher numbers of studies of Ordovician, Jurassic and Cretaceous hardgrounds than other Phanerozoic hardgrounds. This pattern reflects not only the interest in specific time intervals but also the distributional pattern over time (Brett 1988; Wilson and Palmer 1990, 1992). Hardgrounds were particularly abundant in Middle and Late Ordovician, Jurassic and Late Cretaceous. Spectacular Ordovician examples are the beautifully colored Orthoceras limestones of Scandinavia and the Baltics (Lindström 1979). The general community structure of Ordovician hardgrounds shows relative little change during the Paleozoic. Mesozoic encrusting communities differ strongly in their composition from middle Paleozoic counterparts. The general diversity and guild structures observed in modern hardgrounds were firmely established by the Jurassic. Late Cretaceous chalk hardgrounds occur as well as cemented hard beds within a succession of porous, slightly cemented soft chalk. The upper surfaces of these hardgrounds are strongly bored and encrusted, and often show signs of abrasion and corrosion. Glauconitization and phosphatization are common features. The duration of the depositional hiatus associated with the formation of hardgrounds has been estimated at only several hundreds to thousands of years for the chalk, as compared to about 200 000 years for Ordovician hardgrounds. Surprisingly little is known about Late Paleozoic and Tertiary hardgrounds.
5.2.4.2 Condensation Surfaces and Condensed Sections Condensed sections are characterized by strata which are much thinner than their coevally formed stratigraphic equivalents, due to greatly reduced sedimentation rates. Condensed deposits accumulate over a large time span and should have a recognizable and chronological biostratigraphic record. It must be kept in mind that condensation reflects different factors – the decrease of the sedimentation rate (stratigraphic condensation: Heim 1924), the decrease of the accumulation rate (sedimentary condensation), and a mixture of different age fossils (taphonomic condensation: Brett and Baird 1986). Controls on condensed sections: Major processes involved in condensation include mechanical concentration, chemical concentration and biological concentration (Fels and Seyfried 1993):
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Mechanical concentration is achieved by filtering (sieving), screening, bypassing, or scouring sediment. Condensation by sieving is well-known from neptunian dikes (Wendt 1976; Lehner 1992) and from carbonate ramps in times of sea-level rise when the clastic input is reduced (Dromart 1989; Schlager 1989; Leinfelder et al. 1993). Condensation by screening is common across strongly rifted passive continental margins (Bernoulli and Jenkyns 1974; Seyfried 1980; Bice and Stewart 1990). Condensation by bypassing occurs on carbonate ramps during sea-level lowstands. Condensation by winnowing is a common feature on clastic shelves where burial-exhumation-reworking processes are effective (Fürsich 1979; Hallam 1988). A combination of winnowing and filtering processes during rising sea level may lead to shell concentrations on clastic shelves and bonebeds on carbonate platforms (Kidwell 1985, 1986, 1987; Nummedal and Swift 1987; Fürsich et al. 1991). Concentration of lithified bottoms by scouring requires extense current activity (Wilson and Palmer 1992). Chemical condensation can be caused by dissolution below the carbonate compensation depth, as strongly favored by Jurassic cephalopod limestones from the Tethys (Hollmann 1964; Garrison and Fischer 1969). But precisely these examples were questioned with respect to the water depths and the processes involved in the production of the condensed deposits (Schlager 1974). Many authors now favor relatively shallow depositional environments for the Tethyan condensed sections (e.g. Wendt and Aigner 1985; Martire 1992). Relative sea-level fluctuations and a combination of winnowing and bioerosion of differently lithified substrates could be responsible for the condensation features attributed predominantly to dissolution effects at the sea bottom until now. Condensation through biological concentration is caused by bioerosion on hardgrounds as shown by the common record of microborings at pelagic Triassic and Jurassic discontinuity surfaces (Seyfried 1981; Schmidt 1990). Microfacies criteria pointing to condensation and reduced or omitted sedimentation in pelagic environments: • Centimeter- to meter-sized rock units consisting of bedded, variously colored limestones and limestone/ marl sequences. • Alternation of intensively burrowed beds and beds exhibiting current transport and quiet-water deposition. • Variously spaced intercalation of hardgrounds ex-
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Discontinuities
Plate 23 Discontinuities: Criteria of Submarine Hardgrounds and Condensed Carbonate Rocks A hardground is a usually only centimeter-thick zone at the sea bottom, which is lithified to form a hardened surface. It implies a gap or retardation of sedimentation and may be preserved as an unconformity. Common criteria of submarine hardgrounds are sharp upper and diffuse lower contacts, mineralizations (iron hydroxides, manganese oxides, phosphates, glauconite) of upper surfaces in the form of crusts or microstromatolites (–> 2, 4, 6; Massari 1981; Préat et al. 1999), ferruginous impregnations of grains, cements or within peloids and oncoids, encrustations by microbes, foraminifera, serpulids, brachiopods, bivalves et al., borings and burrows in endured substrates (–> 1), e.g. omission surfaces, condensed fauna, lithoclasts resulting from intraformational reworking of hardgrounds, reworked hiatus concretions, interlayers of marine sediment and cements, dissolution features, e.g. steinkerns of primarily aragonitic shells. Hardgrounds originate in ‘deep-marine’ pelagic environments (slopes, basins, ridges) as shown in the examples on this plate, but are also abundant on the sea bottoms of shelves and ramps. 1 Break in sedimentation. A fine-grained biomicrite (BM) is covered by dark red argillaceous microbioclastic micrite, penetrating into a microrelief. Some skeletal grains enter into the lower region as well. Bioturbation (right upper corner) is shown by the whorled alignment of shell fragments. Deposition of the dark sediment probably took place subsequent to borings by bioeroding organisms (cf. Ekdale and Bromley 2001) into hardgrounds or firmgrounds. Differentiation of hardgrounds and firmground borings requires criteria such as the shape of endolithic structures (better seen in the field than in non-oriented thin sections as in this example) and the type of the boundary between the host sediment and the borings (sharp, truncating fossils or discontinuities surfaces, or indistinct). A break in sedimentation is indicated by accumulations of glauconite grains. These ‘Orthoceras limestones’ represent a condensed ca. 50 m thick pelagic sediment, deposited under relatively cool-water marine conditions, on the Baltic Shield, extending one-half million square kilometers (Lindstrom 1963). The rocks are predominantly micritic, fossiliferous and thinly bedded. Complex ichnofabrics indicate that the sea floor ranged from softground to firmground to hardground. Early Ordovician (Limbata limestone, Arenigian): Öland Island, southern Sweden. 2 Hardground: The echinoderm fragments in the bottom half of the picture have been dissolved along a discontinuity and subsequently truncated. Hardground criteria: Limonitic cauliflower structures (CS, ‘deep-water stromatolite’) above the discontinuity surface. Decimeter-sized, dome-shaped microbial microstromatolites growing on discontinuity surfaces, lithoclasts or shells, are common features of deep-marine pelagic limestones (Massari 1981). Early Jurassic (Adnet limestone): Adnet near Salzburg, Austria. The Adnet limestone is an example of widely distributed Tethyan ‘deep-water limestones’ generally summarized as ‘Ammonitico rosso’ because of the common occurrence of ammonites and the red color of the rock (see Pl. 139). This facies represents the post-drowning facies of ‘pelagic carbonate ramps’. The Adnet limestone originated on drowned late Triassic carbonate platforms in upper and lower slope settings. 3 Hardground: Fe/Mn mineralizations upon and within mm-sized columns constructed by agglutinated sessile foraminifera (Placopsilina d’Orbigny, arrows) growing on lithified microbial substrates and forming minute ‘reef’ structures (Wendt 1969). Encrusting agglutinated foraminifera are abundant in modern manganese crusts and nodules. Late Triassic (Hallstatt limestone, Carnian): Feuerkogel near Bad Aussee, Styria, Austria. The Hallstatt limestone originated on the top of fault-bounded ridges in water depths of several hundreds of meters. 4 Detail of ‘deep-water stromatolites’ with abundant sessile foraminifera (arrows). Late Triassic (Carnian): Feuerkogel near Bad Aussee, Styria, Austria. 5 Limonitic (brownish or yellowish) ‘cauliflower structure’ (originating by migration of colloidal Fe-solutions, CS) into a cavity below a discontinuity surface (arrows). Note differences in sediment composition: Gray biomicrite with foraminifera and shells (GB), and red fine-bioclastic biomicrite with ostracods and other skeletal debris (RB) also occurring in the cavity below the discontinuity (arrows). E – echinoid spine. Late Triassic (Norian): Hallstatt, Upper Austria. 6 Pelagic hardground. Note the thin upward growing stromatolitic Fe-Mn crust (arrow). Post-drowning micrites above a sunken atoll. Late Cretaceous?: Isakov Guyot, northwestern Pacific. 7 Hardground at the top of a drowned carbonate platform. Microbial Fe-Mn microstromatolite (arrow) growing perpendicular to the wall of a dissolved component within a completely phosphatized rock. Middle Cretaceous (Late Albian): Charly Johnson Guyot, northwestern Pacific. 8 Phosphatization: Resedimented hemipelagic mudstone clasts with Fe-Mn crusts in a phosphorized matrix. Clasts were transported by submarine currents prior to the formation of the crusts. Cretaceous: Heezen Guyot, northwestern Pacific. –> 6, 7: Grötsch and Flügel 1992
Plate 23: Hardgrounds and Condensed Carbonate Rocks
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Condensation and Hardgrounds
Fig. 5.10. Based on thin-section observations of Early and Middle Jurassic limestones. Note variously preserved ammonite shells, various-sized Fe/Mn nodules and sutured discontinuity surfaces above the hardground. The latter exhibits typical micrite-filled microborings covered by microbial microstromatolites that are often impregnated with Fe and/or manganese oxides. The crusts alternate with micrite layers. A common feature of these crusts in Triassic and Jurassic limestones are attached foraminifera. After Wendt (1970).
hibiting specific trace fossil association on omission surfaces (Bertling 1999). • Reworked hardgrounds and occurrence of poorly sorted lithoclasts. • Bored and encrusted, early diagenetic, clay-hosted concretions (Hesselbo and Palmer 1992). • Ammonite pavements, occurring together with manganese and iron crusts. Primarily aragonitic cephalopod shells are preserved as calcite spar, micrite or goethite pseudomorphs. High percentages of broken or partly dissolved shells (Pl. 90/3). • Highly variable microfacies types occurring only within a distance of a few centimeters and ranging from mudstones, wackestones, packstones and grainstones to floatstones and rudstones. • Common microbiota: Attached foraminifera, serpulids, small gastropods, thin-shelled bivalves, and ostracods. High concentrations of planktic and benthic fossils (sometimes constituting mixed assemblages (Gehring 1986) • Authigenic minerals (glauconite, phosphorite, siderite etc.). • Ferruginous ‘microstromatolite’ crusts (Fig. 5.10), nodules and Fe-ooids (Pl. 22/3): Goethite crusts (massive or consisting of up to 50 µm thick laminae, alternating with calcitic layers and forming flat to domeshaped digitate structures. Centimeter-sized goethite oncoids (Sect. 4.2.4.1). Nuclei are lithoclasts from adjacent hardgrounds or rockgrounds. Laminae contain large amounts of sessile foraminifera. Oncoids occur isolated or densely packed within a single layer. • Manganese crusts and scattered manganese nodules (Banerjee and Iyer 1991). • Phosphatic ‘microstromatolites’, composed of stratiform and columnar sheets consisting of alternating dark (phosphatic, with microbes) and light laminae (nonphosphatic, sediment). Phosphoritized calcareous shells
(Pl. 22/2). Current induced condensation accompanied by phosphogenesis is often associated with platform drowning (Föllmi 1989). • Accumulations of organic matter. The history of many condensed units often exhibits a three-stage sequence: (1) Reduced sedimentation stage, recorded by shell concentrations (e.g. ammonite pavements), bored and encrusted bio- and lithoclasts, reworked iron crusts, sometimes indication of current transport, (2) Omission stage, characterized by irregular, bored and encrusted hardgrounds as well as scoured surfaces exhibiting truncation of pre-omission burrows. Reworking breccias and large cavities may be common, (3) Fe and Fe/Mn crust stage. Time involved in deep-marine condensation horizons and hardgrounds Geological condensation refers to a long-lasting break in the rock-stratigraphic record commonly associated with a hiatus (Fig. 5.11). Very thin rock units may represent very large timespans ranging from hundreds of thousands to several millions of years. Excellent case studies dealing with the processes and the time involved in the formation of condensed carbonate sequences in the Jurassic of Spain give some idea of the immense time spans associated with deep-marine condensation (Seyfried 1978, 1980, 1981; Fels and Seyfried 1993; Fels 1995). The authors described an instructive example showing that at least 10 million years may be concentrated in a one-meter thick condensed carbonate succession. Because the carbonate microfacies types were deposited relatively rapidly, most of the time should be hidden within the goethite crusts, hardgrounds and the erosional discontinuities.
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Submarine Condensation
Fig. 5.11. Submarine condensation and lost or hidden time record. The upper figure shows a wire-cut section of a sequence of strongly condensed Liassic and Middle Jurassic limestones in western Sicily. A distinct marine hardground horizon (2) occurs at the top of Carixian-Domerian crinoidal packstones (1). The crinoidal sand was deposited on a stepped escarpment. The hardground (shown in detail in the lower figure) forms a distinct paleorelief and is characterized by a distinct erosional surface, covered by a dark crust (3) of variable thickness. The crust consists of manganese oxides, goethite and calcite. The origin of the hardground is interpreted as result of dissolution related to the widespread early Toarcian anoxic event in the Tethyan Sea. The overlying limestone succession represents the typical Ammonitico Rosso facies of the Mediterranean region (see Pl. 139). The succession consists of several units differing in texture, microfacies and biostratigraphical age: Fillings of depressions at the top of the hardground (4) yield bio/lithoclastic wackestones with reworked Fe/Mn crusts. The overlying Rosso Ammonitico Inferiore unit (RAI) consists of red nodular cephalopod limestones (RAIa, 5), pink wavy cross-bedded limestones (RAIb, 6) and red pseudonodular limestones (RAIc, 7). Note the current ripple structure of unit RAIb indicating episodic deposition of bioclastic sands with pelagic bivalve shells (Bositra) from shifting currents of microturbidites on the otherwise starved sediment surface. Unit RAIc is strongly bioturbated and contains small-scale FeMn coated condensation surfaces as well as laminated goethitic Fe-Mn rich marls (lgm) exhibiting signs of early deformation (8) subsequent to compaction. Ammonites found in the limestones of the units permit the recognition of various biozones which indicate the existence of the Toarcian (2, 3), Bajocian (4), Bathonian (6) and Callovian (7). Some of the missing biozones are known from other nearby localities, but some biozones could not be verified. The time represented by these biozones may be hidden in the condensation horizons, but just imagine, the total duration of the end-Torcian to middle Callovian time ranges is about 20 million of years! Modified from Di Stefano et al. (2002).
Assuming a comparable growth rate of goethite crusts and modern deep-sea manganese nodules, a timespan of about 9.5 million years is estimated for the duration of the discontinuity events. Different growth rates (60 mm/103 and 200 mm/ 103 years) were estimated for Fe/Mn crusts on Jurassic hardgrounds using sessile foraminifera growing upon ferruginous laminae (Wendt 1970). A time range between 25 000 and 125 000 years was estimated for the duration of the breaks associated with the formation of Ordovician hardground of the Scandinavian-Baltic shelf (Lindström 1979).
5.2.5 Discontinuities and Sequence Stratigraphy Unconformities and condensed sections are essential for the identifying depositional sequences and their boundaries. The observation and interpretation of small-scale discontinuities assist in high-resolution sequence analysis, of shelf carbonates (Krawinkel and Seyfried 1996) and pelagic carbonates (Martire 1992).
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Conventional sequence stratigraphy relates the large-scale stratal architecture of sediment bodies to changes in sea level. Stratal architecture alone, however can be a rather ambiguous guide to sea-level history. A combination of microfacies analysis, the study of diagenetic patterns and stratal geometries is required to single out the role of changes in accommodation and carbonate production for the development of platforms in comparison with sea-level controls. An excellent example was provided by Evarts et al. (1995) while studying early Cretaceous platform margins. Frequency and distributional patterns of environmentally diagnostic grain types and discontinuities in sections covering different parts of the platform and the platform-basin transition allow lowstand and highstand sedimentation phases to be differentiated, resulting in a significant refinement of the sequence stratigraphic model. Other advantages in using microfacies data and particularly the types and distribution of discontinuities have been demonstrated by Hillgaertner (1998). The author compared the distribution of the major discontinuity types in carbonate platform sections with depositional types characterized by microfacies and biota. This resulted in the recognition of ‘discontinuity zones’ that reflect specific positions within the sequence stratigraphic framework. Most zones with prevailing exposure surfaces correspond to large-scale sequence boundaries. In contrast, discontinuity zones with dominating condensation surfaces fall between the major sequence boundaries. This implies that the occurrence of specific small-scale discontinuities reflects large-scale second- or third-order sea-level changes. The generalization that condensed sections found in deep-marine carbonates have been developed exclusively during transgressions and that condensation intervals should be interpreted as the result of transgressive maxima is questionable. Gomez and Fernandez Lopez (1994) drew attention to the fact that the relative distribution of unconformities and condensed sections should not be used as unequivocal diagnostic criteria for particular paleogeographical settings (e.g. basinal),or as indicators of maximum regional transgressions, because condensation processes also occur in the shallowest parts of carbonate platforms and during maximum regression periods. Further implications of microfacies studies in the context of sequence stratigraphic investigations are discussed in Sect. 16.1.2. Basics: Discontinuity surfaces in carbonate rocks Bertling, M. (1999): Taphonomy of trace fossils at omission surfaces. – Palaeogeogr. Palaeoclimat., Paleoecol., 149, 27-39 Brett, C.E. (1988): Paleoecology and evolution of marine hard substrate communities: an overview. – Palaios, 3, 374-378
Basics: Discontinuities
Budd, D.A., Saller, A.H. and Harris, P.M. (eds., 1995): Unconformities and porosity in carbonate strata. – Mem. Amer. Ass. Petrol. Geol., 63, 313 pp. Clari, P.A., Dela Pierre, F., Martire, I. (1995): Discontinuities in carbonate successions: identification, interpretation and classification of some Italian examples. – Sed. Geol., 100, 97-112 Evarts, A.W.E., Staffleu, J., Schlager, W., Fouke, B.W., Zwart, E.W. (1995): Stratal pattern, sediment composition, and sequence stratigraphy at the margin of the Vercors carbonate platform (Lower Cretacous, SE France). – J. Sed. Research, B65, 119-131 Fels, A. (1995): Prozesse und Produkte geologischer Kondensation im Jura der westlichen Tethys. – Profil, 9, 363-472 Fürsich, F.T., Oschmann, W., Singh, B., Jaitly, A.K. (1992): Hardgrounds, reworked concretion levels and condensed horizons in the Jurassic of western India: their significance for basin analysis. – J. Geol. Soc. London, 149, 313-331 Goldring, R., Kazmierczak, J. (1974): Ecological successions in intraformational hardground formation. – Palaeontology, 17, 942-962 Gomez, J.J., Fernandez-Lopez, S. (1994): Condensation processes in shallow platforms. – Sed. Geol., 92, 147-159 Heim, A. (1934): Stratigraphische Kondensation. – Eclogae geol. Helvetiae, 27, 372-383 Hillgaertner, H. (1998): Discontinuity surfaces on a shallow carbonate plattform (Berriasian, Valangian, France and Switzerland). – J. Sed. Research, B68, 1093-1108 Krawinkel, H., Seyfried, H. (1996): Sedimentologic, paleo ecologic, taphonomic and ichnologic criteria for high-resolution sequence analysis: a practical guide for the identification and interpretation in shelf deposits of discontinuities in shelf deposits. – Sed. Geol., 101, 79-110 Lindström, M. (1979): Diagenesis of Lower Ordovician hardgrounds in Sweden. – Geologica et Palaeontologica, 13, 9-30 Palmer, J.T. (1982): Cambrian to Cretaceous changes in hardground communities. – Lethaia, 15, 309-323 Pedley, H.M., Bennett, S.M. (1985): Phosphorites, hardgrounds and syndepositional subsidence: a paleoenvironmental model from the Miocene of the Maltese Island. – Sed. Geol., 45, 1-34 Voigt, E. (1959): Die ökologische Bedeutung der Hartgründe (hardgrounds) in der oberen Kreide. – Paläontologische Zeitschrift, 33, 129-147 Wendt, J. (1988): Condensed carbonate sedimentation in the late Devonian of the eastern Anti-Atlas (Morocco). – Eclogae geol. Helvetiae, 81, 155-173 Further reading: K102, K174
5.3 Syn- and Postdepositional Features: Fissures, Veins, Breccias Carbonate rocks are often affected by destructive processes leading to the opening and filling of variously sized fissures, voids and veins, and the formation of limestone breccias and conglomerates. Some of these processes occur syndepositionally, others take place a long time after final lithification.
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Neptunian Dikes
5.3.1 Sediment-Filled Fissures: Neptunian Dikes and Fissure Fills Many limestones yield fissures filled with carbonate or other sediment (sedimentary dikes) and with carbonate cements. Fissure walls are commonly subparallel, planar or irregularly undulated. The fissures cut the host rock obliquely or vertically across bedding planes, or parallel to the bedding. The width of vertical or obliques fissures varies between a few centimeters and several meters, they may be several hundreds of meters up. Horizontal sediment-filled fissures can be followed a few tens of centimeters up to several hundreds of meters. Many fissures occur beneath erosional or stratigraphic unconformities. Fissures may extend several tens of meters below unconformity surfaces. The formation and infilling of these fissures can take place both in submarine and subaerial environments. Submarine fissures are common in platform carbonates, particularly at platform margins, but also occur in slope and basinal settings. Terminology Several terms describe submarine and subaerial voids and fissures filled with sediment (see Box 5.6). Definitions of these terms are not consistent, because of differences in the descriptive and/or genetic aspects. Common terms are fissure and dike. Fissure stresses the existence of fractures within rocks, whereas dike relies on the filling with sedimentary material and chemically precipitated cements. The conditions of sediment deposition are conventionally used to define the major types of sediment-filled fissures. Neptunian dikes are defined as fissures within rocks exposed on the sea bottom that have been filled with submarine sediments. Most definitions include discordant cutting of bedding, infill of younger (but also stratigraphically coeval) sediments and infilling from above. If we accept these criteria, then neptunian dikes represent a specific type of sedimentary dike. The most workable definition of neptunian dikes is probably that of Fischer (1964): ‘Discordant dikes of limestones within limestones’ parallel to the bedding planes. This definition coincides with that of ‘S-fissures’. The names S- and Q-fissures (German: ‘S- und Q-Spalten’, Q = quer, oblique) describe the geometrical relationship of fissures with the host rock (Wendt 1971). Q-fissures cut the bedding (Pl. 24/1, Fig. 5.12), and S-fissures run parallel to it (Pl. 24/2, Fig. 5.13). The use of the term fissure fill for continental sediments filling subaerial cavities and fissures (e.g. paleo-
Box 5.6. Glossary of terms used in the context of the study of sediment-filled fissures. Based on Bates and Jackson (1980), Visser (1980), Smart et al. (1994). Clastic dike: A sedimentary dike, transecting bedded sediments at various angles, and consisting of clastic material derived from underlying or overlying beds. Cavity fill: Subaqueous fills deposited in brackish or phreatic freshwater environments (Swart et al. 1988). Dike (British spelling: dyke): A discordant or concordant fissure filling. Fissure: A fracture or a crack in rocks along which there is a distinct separation of walls. Fissure fill: Terrestrial deposit infilling voids and fissures in older rocks created, e.g. by karstification. The term is the terrestrial equivalent to neptunian dikes referring to submarine infillings. Injection dike: A sedimentary dike cutting vertically through massive or normally bedded strata filled by material through hydrostatically controlled pressure of injection from below or above or from the side. Neptunian dike: A sedimentary dike formed by infilling of submarine sediment in fissures and cavities of rocks exposed on the sea floor. The term was originally coined by Pavlov (1896) for dikes, occurring in Early Cretaceous clays and filled with Oligocene sandstone. The British term ‘neptunian’, now commonly used for carbonate rocks, refers to the submarine origin of the fissures and their submarine infillings. Neptunian sill: Fissures that essentially follow bedding planes and are filled with submarine sediments (Playford et al. 1984). Sedimentary dike: A tabular mass of sedimentary material that cuts across the structure of bedding of preexisting rock. It is formed by the filling of a crack or fissure from below, above or laterally.
karst, see Sect. 15.2.1) may be meaningful but sometimes troublesome when mixed infills of terrestrial and marine sediment occur.
5.3.1.1 Origin, Development and Filling of Sedimentary Fissures The initiation, development and sedimentary filling of fissures can take place in a variety of environments and is influenced by various processes such as karst solution in subaerial and vadose settings, solution in phreatic and mixing zones between freshwater and saline waters, and processes occurring in shallow and deep marine settings. Initiation and development of fissures Initial fissures can be depositional, synsedimentary or tectonic in origin. Primary depositional fissures are
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of some significance only in reef carbonates and forereef breccias. These fissures correspond to cavities between large reef-building organisms or reef blocks. The cavities are largely filled at an early stage by sediment derived from the erosion of the reef itself (Land and Goreau 1970). Synsedimentary fissures develop by downslope sliding and slumping of early-cemented slope deposits on uncemented interbeds as a result of earthquake shak-
Tectonic Fissures
ing (Martire 1996). Shallow-marine lagoonal carbonates can be affected by desiccation and shrinkage leading to the formation of dikes (Fischer 1964). Tectonic fissures are often related to block faulting and rifting. Regional tectonics associated with active block faulting and extensional movements causes the formation of faults and gravity sliding of large sedimentary blocks along unstable slopes due to oversteepening (Füchtbauer and Richter 1983; Winterer and Sarti
Plate 24 Sediment-Filled Fissures: Neptunian Dikes and Terrestrial Infills Vertical, horizontal or oblique millimeter- to meter-sized fissures filled with carbonate or other sediment or with carbonate cements can be formed in submarine and subaerial environments. Neptunian dikes are ‘discordant dikes of limestones within limestones’ filled with submarine sediments. Fissures are common in platform carbonates, particularly in drowned platforms (–> 4) and at platform margins, but also occur in slope and pelagic limestones (–> 1, 2, 6). Submarine formation of fissures is triggered by synsedimentary tectonics and rifting or differential compaction. Subaerial fissure formation occurs during karstification (–> 5) and vadose dissolution. Internal fillings in neptunian dikes consist of mud, pelagic and benthic fossils and lithoclasts eroded from fissure walls. Microbial crusts can occur. Common cements are early-marine radiaxial-fibrous calcite and later granular calcite covering the walls of fissures. Common fillings of paleokarst fissures are erosional carbonate lithoclasts and relicts of overlying deposits, silt, and pedogenic sediment, carbonate cements are of meteoric-phreatic and burial origin. Analysis of neptunian dikes requires one to study the geological framework, compare the microfacies of fissure sediments and host rocks, evaluate fossils (taphonomy, biotopes, age) and examine the diagenetic criteria. Common microfacies types of submarine fillings are red and gray mudstones and wackestones/packstones containing foraminifera, mollusk and echinoderm coquinas and lithoclasts. The timing of fissure fillings can be determined by microfossils, particularly foraminifera. Differences in texture reflect open or closed fissure systems. Alternations of cement tapestries and sediment infill reflect repeated rupture of fissures. Neptunian dikes indicate breaks in sedimentation and changes in sea level and subsidence. Low sedimentation leads to condensed sequences with condensed faunas, covering a substantial time period. 1 Neptunian dike (outlined in white) within red pelagic limestones with ‘filaments’ (pelagic bivalves). The host rock is a filament wacke/packstone, the filling is a fossiliferous mudstone. Note the distinct difference between the microfacies of the host rock and that of the infilling sediment. This difference and the irregular wavy boundary of the fissure (white line) indicate a modification of the fissure system by submarine dissolution and a certain time break. Deep-water limestone. Late Triassic (Carnian): Hydra Island, Greece. 2 Internal filling of a horizontal neptunian dike. The alternations of grain-rich and grain-free laminae and aligned shells point to the existence of bottom currents within the dikes. Jurassic (Rosso Ammonitico Veronese): Altopiano di Asiago, Italy. 3 Repeated rupture and fissure formation. Two generations of fissures (outlined with white ink) in a red pelagic wackestone. An older system is cut by younger fissures filled with radiolarian packstone. Late Triassic (Norian): Epidavros, Greece. 4 Submarine fissures in drowned platform carbonates. The Late Triassic host rock (peloid grainstone, PG) is cut by repeatedly ruptured fissures exhibiting several phases of Liassic fillings: 1: Red echinoderm wackestone, 2: Red mudstone, 3: Calcite vein cutting the red mudstone, 4: Gray crinoid packstone. The arrow points to the youngest fissure. Triassic/ Jurassic: Kirchbruch, Adnet, Austria. 5 Fissure fill. Terrestrial fissure within Jurassic shallow-marine bindstones. The karst fissure has been filled with late Cretaceous angular to subrounded limestone clasts. Note the plane boundary of the fissure and the cracks filled with burial calcite cement. Late Jurassic (Ernstbrunn Limestone, Tithonian): Ernstbrunn, north of Vienna, Austria. The Ernstbrunn limestone is part of a Late Jurassic-Early Cretaceous carbonate platform which was affected by intensive karstification prior to a Late Cretaceous transgression (Hofmann et al. 1999). 6 Pelagic wackestone (PW) with filaments, radiolarians and Globochaete, cut by synsedimentary subvertical Q-fissures, infilled by fine-grained sediment and juvenile mollusks (M). Note bivalve (Halobia) coquina, C) at the bottom and thin calcite tapestries at the walls of the fissures. Late Triassic (Hallstatt limestone, Norian): Bad Ischl, Upper Austria. –> 1, 3: Dürkoop et al. 1986, 2: Martire 1996
Plate 24: Neptunian Dikes and Terrestrial Infill
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Fissure Infillings
gence, commonly exhibit changes in shape and dimensions due to solution, bioerosion and karstification. Filling In general, fissure infill occurs much less rapidly in terrestrial than in marine environments. Marine fissure fillings are primarily controlled by the relative rates of autochthonous and allochthonous sedimentation, by the size of the fissures and by the interconnection of the fissure with the external environment. Laminations and crossbedding of internal sediments within neptunian dikes can indicate relatively high flow rates within fissure systems (Pl. 24/2; Figs. 5.12, 5.13). Fig. 5.12. Early Jurassic neptunian dike (right) within a Late Triassic reef limestone (left). The organic structures of the white Upper Rhaetian reef limestones (Lamellata wähneri Flügel and Sy, a hydrozoan - L, and a cyanobacterial crust C) are cut by a vertical fissure (Q-fissure) filled with red Liassic echinoderm packstone. The planar boundary between the host rock and the dike indicates a tectonic origin of the fissure (related to initial rifting of a drowned platform). The peripheral calcite tapestry consisting of radiaxial cement points to a time break between the opening of the fissure and the infilling of the Liassic sediment. Note the solution seams on both sides of the calcite tapestry (arrows). Rofan, Sonnwend Mountains, Tyrol, Austria. Scale is 5 mm.
1995). Fissures can also be produced by differential compaction of platform and basinal carbonates (Saller 1996). Rapid extensional tectonic activities and high hydrostatic pressure can, perhaps, also cause injection of sediments into voids and fissures (Castellarin 1982). Hydrothermal effects related to thermochemical effects (Sotak et al. 1993; Belka 1998) might also result in the formation of neptunian dikes. Opening of fissures in platform carbonates during rifting phases and infilling of pelagic sediment during and/or after drowning of the platforms have been repeatedly reported from the Mesozoic of the Alpine-Mediterranean region. Many dikes are associated with regional extensional features, as shown by the coincidence of the preferred orientations of neptunian dikes and tectonic structures (Blendinger 1986). Neptunian dikes are not restricted to platforms and platform margins but also occur within pelagic carbonates deposited on tilted fault blocks (Martire 1996). Opening of bedding planes can occur as a result of both tectonic and synsedimentary induced fracturing. Fracture orientation parallel or oblique to platform margins can be indicative of synsedimentary or tectonic origins of fissures, respectively. Many neptunian dikes show little evidence of dissolutional and mechanical enlargement after the initial opening. Conversely, fissures associated with emer-
Neptunian dikes are filled with sediment and fossils derived from variable sources, and carbonate cements formed at different times. Internal sediments comprise allochthonous grains and mud as well as autochthonous elements originating within the fissures. Allochthonous constituents are predominantly fossils (commonly pelagic fossils or organisms living on the sea-bottom nearby) forming densely packed, stratigraphically condensed accumulations of cephalopods, thin-shelled bivalves (e.g. Posidonia, Halobia, ‘filaments’) and gastropods (‘Fossillagerstätten’) as demonstrated by Triassic and Jurassic S- and Q-fissures in Austria and Sicily (Krystyn et al. 1971). Many of these fossils have been swept into and trapped in the fissures but some, characterized by significantly small growth forms, represent probably cavity dwellers (Wendt 1976). Common cavity-dwelling organisms are ostracods (Aubrecht 1997), foraminifera, and some brachiopods. Autochthonous sediments consists of material eroded from fissure walls and fossils formerly living at the edges of or within neptunian dikes (GonzalezDonoso et al. 1983; Playford 1984). Autochthonous biota comprise microbial crusts as well as benthic foraminifera, ostracods, serpulids and brachiopods. Microbes are able to form distinctive morphological structures (fissure-dwelling ‘endostromatolites’: Monty 1982, see Sect. 9.1). Terrestrial infill within subaerially formed fissures differs from the internal sediments of neptunian dikes in the lack of marine fossils (Pl. 24/5). Internal sediments of paleokarst fissures consist commonly of fine limestone or dolomite lithoclasts eroded from fissure walls, erosional relicts of overlying deposits, calcareous silt, siliciclastics, and redeposited pedogenic sediment (MacDonald et al. 1994; Pl. 24/5, Pl. 129/1). Carbonate cements of paleokarst fissures are often of meteoric-phreatic origin (see Sect. 15.2.1).
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Neptunian Dikes
5.3.1.2 Microfacies Analysis of Neptunian Dikes Field data of neptunian dikes include rock color, orientation, geometry, and size of the fissures as well as the spatial relationships between fissures and breccias (Füchtbauer and Richter 1983). The color of the neptunian dikes can be distinctly different from that of the host rock (e.g. red Liassic fissure infillings in gray or white Triassic platform and reef carbonates of the Alpine-Mediterranean region, Figs. 5.12 and 5.13), but many synsedimentary fissures are filled with carbonates displaying the same color as the host rocks. Sharp boundaries between the fissure infillings and the host rock indicate a tectonic origin of the fissures (Fig. 5.12), wavy and undulated boundaries can be caused by subaerial or submarine dissolution and erosion (Pl. 24/1). Of major importance is the comparison of microfacies types of fissure sediments and host rocks, the evaluation of fossils with regard to taphonomic and ecologic criteria as well as original biotopes and age, and the investigation of carbonate cements using thin sections, cathodoluminescence petrography, and isotope data. Microfacies data which should be evaluated in thin sections are: Composition and texture of internal sediment: Common microfacies of submarine fillings are red and gray mudstones and bioclastic wackestones and packstones with foraminifera, mollusk shells (ammonites, gastropods, accumulations of pelagic pelecypods, Pl. 24/6) and echinoderms (accumulations of crinoids, Pl. 24/4, Figs. 5.12 and 5.13). In addition, brachiopods and various pelagic biota, e.g. radiolarians and globochaete algae may occur (Pl. 24/3). Lithoclasts, derived from fissure walls, consisting of reworked calcite cements, or resulting from the erosion of host rocks and stratigraphically overlying rocks, may be common. Lithoclasts can represent reworked fragments of hardgrounds (Seyfried 1981; Wendt 1976). Differences in the sorting of particles reflect fissure systems that are more or less open and closed (Winterer et al. 1991). S-parallel fissure infills are commonly finebedded (Fig. 5.13). The individual beds can differ in color shades, fossil content and stratigraphic age of the fossils included. Depending on the environment the ‘internal sediments’ are of shallow-marine to pelagic origin. Microfacies of neptunian dikes were described by Gonzalez-Donoso et al. (1983), Vachard et al. (1987) and Vera et al. (1988). Cements: Fissures that were inaccessible to infilling sediments are lined and partly or totally filled with
calcite cement (Fig. 5.13). Common cements are earlymarine radiaxial-fibrous calcite and younger granular calcite. Age-diagnostic micro- and macrofossils: Fissure infillings record continuous or interrupted sedimentation over very long time spans. These time spans can often be evaluated with the help of conodonts, foraminifera, and ammonites and by using ‘microfacies zones’ (see Sect. 10.3). Repeated rupture and filling: Alternations of cement and sediment infill indicate repeated rupture and filling of the fissures (Pl. 24/4, Wiedenmayer 1963). As many as twenty generations of fillings have been observed. Excellent examples are neptunian dikes cutting through Upper Triassic lagoonal limestones and filled with several generations of Liassic sediments at the Loser in the Styrian Northern Alps.
5.3.1.3 Case Studies of Neptunian Dikes in Carbonates Neptunian dikes are particularly associated with carbonate buildups because • lithified carbonate responds in a brittle manner to stress and develops fractures at an early stage of lithification, • steep carbonate slopes tend to develop fractures due to unloading and mass movements along margins of active and drowned platforms, and • the solubility of carbonate favors percolation of undersaturated waters, assisting in void and fissure formation.
Box 5.7. Selected case studies of neptunian dikes. Tertiary: Premoli Silva et al. 1998; Vachard et al. 1987 Cretaceous: Andrieux 1967; Company et al. 1982, Garcia-Hernandez et al. 1989, Greber 1988 Jurassic:Böhm et al. 1999; Bouillin and Bellono 1990; Fabricius 1968; Fels and Seyfried 1993; Jurgan 1969; Lantos and Mallarino 2000; Lehner 1991; Martire 1996; Molina et al. 1985; Schöll and Wendt 1971; Seyfried 1978, 1980; Vera et al. 1984; Wendt 1969; Wiedenmayer 1963; Winterer and Sarti 1994 Triassic: Blendinger 1986; Cozzi 2000; Fabricius 1966; Fischer 1964; Füchtbauer and Richter 1983; Krystyn et al. 1971; Schlager 1969; Schöll and Wendt 1971; Wendt 1973; Wiedenmayer 1963 Permian: Pray and Stanton 1992; Stanton and Pray 1993 Devonian: Belka 1989; Gischler 1992; Krebs 1968; Macdonald et al; 1994, Playford et al. 1989; Szulczewski 1973; Tucker 1973; Wendt and Belka 1991; Wendt et al. 1984
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Examples of Neptunian Dikes
Impressive examples of long-ranging graduated fissure fillings have been reported from Devonian, Triassic and Jurassic platform and reef carbonates as well as from slope and basinal carbonates (Box 5.7). The ‘Blue Holes’ of the Bahamas: a modern example of neptunian dikes Underwater caves formed on synsedimentary fractures paralleling the margin of the Bahama Bank and enlarged by dissolution activity in a mixing zone of fresh and saline waters may represent a modern analogue of neptunian dikes (Smart et al. 1988; Mylroie et al. 1995). These so-called Blue Holes got their name from the seemingly blue water of the surface water. The underwater cavities were formed as a result of erosional and dissolution processes controlled by Pleistocene high and low stand phases of the sea level. Some of the fractures of the Bahama Bank are associated with large submerged cave systems. Steep-sided planar walls, synchronous deposition of erosional and encrusting biota within the cavities, as well as the size and geometry of the cavities indicate that the Blue Holes are a good present-day example of active neptunian dikes, fissures and cavern infill. Neptunian dikes in Devonian reef and platform carbonates, Western Australia Neptunian dikes are frequent along the Frasnian platform margins of the Canning Basin (Playford et al. 1984). They become rarer towards the platform interior in the reef-flat facies and down marginal slopes in reef slope facies. Interestingly, dikes are especially well developed in Famennian back-reef limestones formed subparallel to the Frasnian reef front. Individual dikes may become 6 km long, 80 m deep and 20 m wide. But most of them are less than 100 m long, 20 m deep and about 5 cm wide. S-parallel dikes (‘neptunian sills’) are very abundant in marginal-slope deposits. Both fissure types show a complex history of fissure development, filling, lithification and refracturing. Early fillings include encrusting organisms, fossil debris, ooids, lime sand and mud, terrigenous sand, cavity peloids and ‘spar balls‘, radiaxial calcite and internal breccias. Synsedimentary dikes in the Late Permian reef complex of Texas and New Mexico Vertical dikes near and parallel to the shelf-edge, composed of encrusting carbonate laminae, as well as dikes with abundant marine biota (algae and/or invertebrates) are known from the Capitan limestone (Stanton and Pray 1993). Fossil-rich dikes are usually 0.3-0.5 m wide and exhibit progressive opening, fill-
Fig. 5.13. Horizontal Liassic fissure (S-fissure) in Late Triassic peloidal grainstones (bottom, PG). Note that the picture does not show the entire width of the fissure. The Upper Rhaetian grainstone originated in a back-reef setting. The neptunian dike exhibits two-fold fillings: (1) alternating infillings of fine-grained crinoidal debris and carbonate mud, followed by (2) occlusion of a younger fissure within the lithified echinoderm packstone by several generations of calcite cement. The lack of calcite tapestries at the boundary (indicated by two short black lines) between the host rock and the fissure and current-induced textures within the fissure (alignment of shells - black arrows, laminations in micrite layers white arrow) points to a rapid infilling of the dike. Both, the Triassic host rock and the neptunian dike were transected by a network of thin microcracks (3), filled by gray calcareous silt possibly of meteoric origin and related to late tectonics. The Early Jurassic age of the echinoderm packstone is indicated by diagnostic foraminifera (In - Involutina liassica Jones). Rofan, Sonnwend Mountains, Tyrol, Austria. Scale is 5 mm.
ing by intact and fragmented skeletal grains and/or cement, and reopening which creates vertical layering. Most bioclasts in the dikes were loose seafloor particles, but some (whole and paired shells and large Tubiphytes) probably were fissure dwellers. The scarcity of micritic
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Sediment-Filled Fissures
internal sediment in fossil-rich dikes suggests little free mud on the sea floor, bolstering the cemented seafloor concept of the Capitan limestone. Tectonic Jurassic dikes in Triassic limestones of the Northern Calcareous Alps, Austria The Sonnwend Mountains in Tyrol represent an excellent example of Jurassic fissure infillings of different stratigraphic age in Late Triassic reef limestones (Fabricius 1962; Wendt 1969; Pl. 24/4, Figs. 5.12 and 5.13). The reef limestones are overlain by a condensed Jurassic sequence of red bedded limestones. Identical lithofacies types occur in vertical and oblique fissures starting at the Triassic/Jurassic boundary and transecting the upper part of the reef limestone. In contrast, lithofacies types and fossils in the sedimentary infillings of the horizontal ‘S-dikes’ within the reef limestones, running parallel to bedding planes, differ strongly from that of the Q-dikes. These conspicuous differences have caused considerable problems in deciphering the structure of the mountains prior to the recognition of the red apparent intercalations as sedimentary fissures.
5.3.1.4 Significance of Sediment-Filled Fissures Fissures and neptunian dikes are of major importance in basin analysis: • Internal sediments can preserve biota and sediment for periods where other records are lacking. Low sedimentation rates within fissures lead to condensed sequences covering substantial time periods. • If sediments are not recognized as the internal sediments of neptunian dikes severe stratigraphical and structural misinterpretations may be the result. • Neptunian dikes and fissure fills are sensitive sealevel indicators. Changes in the composition of internal sediments in dikes and fissures can reflect changes between marine nd subaerial settings, e.g. the beginning of transgressions or long-lasting karstification phases (Gonzalez-Donoso et al. 1983). • The existence of multiple crosscutting neptunian dikes can be used to reconstruct recurrent syndepositional extensional tectonic activities along drowned platform margins and as paleostress indicators (Lehner 1991, Cozzi 2000). Indications of tensional disruption other than neptunian dikes are fractures, normal faults, shatter breccias and laterally discordant intraformational breccias dissecting platform-margin carbonates.
Basics: Neptunian dikes and fissure fills Cozzi, A. (2000): Synsedimentary tensional features in Upper-Triassic shallow-water carbonates of the Carnian Prealpes (northern Italy) and their importance as paleostress indicators. – Basin Research, 12, 133-146 Fischer, A.G. (1964): The Lofer cyclothems of the Alpine Triassic. – In: Merriam, D.F. (ed.): Symposium on cyclic sedimentation. – Kansas Geol. Surv. Bull., 169, 107-149 Füchtbauer, H., Richter, D.K. (1983): Relations between submarine fissures, internal breccias and mass flow during Triassic and earlier rifting periods. – Geologische Rundschau, 72, 53-66 Garcia-Hernandez, M., Rey, J., Vera, J.A. (1989): Disques neptunicos de edad Cretacica en la Sierra de Quipar (Subbetico externo, Prov. Murcia). – Rev. Soc. Geol. España, 2, 85-93 Krystyn, L., Schäffer, G., Schlager, W. (1971): Über Fossillagerstätten in den triadischen Hallstätter Kalken der Ostalpen. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 117, 284-304 Lehner, B.L. (1991): Neptunian dykes along a drowned carbonate platform margin: an indication for recurrent extensional tectonic activity? – Terra Nova, 3, 593-606 Martire, L. (1996): Stratigraphy, facies and synsedimentary tectonics in the Jurassic Rosso Ammonitico Veronese (Altopiano di Asiago, NE Italy). – Facies, 35, 209-236 Playford, P.E., Kerans, C., Hurley, N.F. (1984): Neptunian dikes and sills in Devonian reef complexes of the Canning Basin, Western Australia. – Amer. Ass. Petrol. Geol. Bull., 68, p. 517 Schöll, W.U., Wendt, J. (1971): Obertriadische und jurassische Spaltenfüllungen im Steinernen Meer (Nördliche Kalkalpen). – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 139, 82-98 Smart, P.L., Palmer, R.J., Whitaker, F., Wright, V.P. (1988): Neptunian dikes and fissure fills: an overview of some modern examples. – In: James, N.P., Choquette, P.W. (eds.): Paleokarst. – 149-163, Berlin (Springer) Vera, J.A., Ruiz-Ortiz, P.A., Garcia-Hernandez, M., Molina, J.M. (1988): Paleokarst and related pelagic sediments in the Jurassic of the Subbetic zone, southern Spain. – In: James, N.P., Choquette, P.W. (eds.): Paleokarst. – 364384, Berlin (Springer) Wendt, J. (1971): Genese und Fauna submariner sedimentärer Spaltenfüllungen im mediterranen Jura. – Palaeontographica, Abteilung A, 136, 122-192, Stuttgart Winterer, E.L., Sarti, M. (1994): Neptunian dykes and associated features in southern Spain: mechanics of formation and tectonic implications. – Sedimentology, 41, 1109-1132 Further reading: K103
5.3.2 Microfractures and Veins (Calcite Veins) Many carbonate rocks display millimeter- to centimeter-sized, mineral-filled (often calcite-filled) microfractures (veins and veinlets, Box 5.8). Fractures are discrete breaks within a rock mass and comprise microfractures, joints and faults. The existence of ductile shear zones, brittle shear zones and brittle-ductile shear
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zones may be deduced from calcite-filled cracks in limestones. The spatial distribution of geological fractures is of particular interest with regard to reservoir potential, hydrology or slope stability. Microfractures are only a part of largescaled fracture systems, and can therefore, be genetically understood only in the context of regional structural analysis. However, small-scaled fractures should not be overlooked in microfacies studies because of their importance in deciphering tectonic activities, fluid migrations, diagenetic history, reservoir potential and the mechanical properties of carbonate rocks.
5.3.2.1 Origin and Classification of Calcite Veins Fracture systems in carbonate rocks form in various stages of sedimentation (e.g. oversteepening of semilithified sediments, network of thin veinlets representing water escape structures: Macdonald et al. 1994) and diagenesis (e.g. compaction load, freshwater dissolution, deep-burial diagenesis: Geiser and Sansone 1981), but most calcite veins are due to brittle failure and tectonic fracturing of lithified carbonate rocks. Tectonic fracturing is caused by • stress and shear displacement (Narr and Burruss 1984; Logan 1984; Ramsay and Huber 1987; Roehl and Weinbrandt 1988; Vrolijk and Sheppard 1991), • extensional movements, or • natural hydraulic fracturing (Sibson et al. 1975; Gretener 1976). Abnormally high natural pore fluid pressures form during (a) rapid sediment burial in basins of high sedimentation rate, (b) tectonic deformation causing closure of pore spaces at one site and high pore pressure due to migrating fluids at some other site, or (c) aquathermal pressure or mineral-phase changes as a result of diagenesis or low-grade regional metamorphosis and contact metamorphism (e.g. smectite-illite transformation). These fractures are characterized by the dominant vertical orientation and parallel nature of the fractures (Longman 1985). Vertical orientation, parallel nature and close vertical spacing of fractures, and the absence of shear offsets can be indicative of hydrofracture origin. Indications of fractures formed by extension are the lack of shear offsets on the fractures and the presence of only a single fracture set. The history and timing of shearinduced fractures in carbonate rocks is elucidated by laboratory experiments providing information on the controls of fracture development by stress, strain, strain
Calcite Veins
Box 5.8. Terms refering to tectonically induced fissures. Fractures: Discrete breaks in a rock mass where cohesion was lost due to brittle failure by stress. Fractures include faults, across which shear displacements occur, joints which have an aperture but show no significant shear displacements at hand specimen scale, and filled fractures, such as veins and dikes (Ramsey and Huber 1987, Gillespie et al. 1993). Microcracks: An opening that occurs in rocks and has one or two dimensions smaller than the third dimension (Simmons and Richter 1976). Generally millimeters to centimeter in size. Microfractures: mm- to cm-sized structures developed by brittle failure. Veinlets: Small veins. Veins: Fractures filled with coarse crystalline calcite or various epigenetic minerals. Vein widths range from millimeters to over a meter, vein walls are subparallel, planar or irregular undulating. Contacts with host rock are usually sharp but embayed margins due to solution and/or replacement of host rock are also common (Logan and Semeniuk 1976).
rate, fluid pressure, temperature and lithology (Paterson 1978). Crack pathes observed in carbonate rocks appear to be controlled by texture and composition (e.g. irregular cracks preferentially within the matrix, not through the grains: Hugman and Friedman 1979; Olsson 1974). Fracture abundance often shows an inverse correlation to the amount of argillaceous material. Relations also exist between the carbonate texture type and the frequency of calcite veins. Lime mudstones and wackestones seem to be more strongly affected by shear-induced fracturing than grainstones. Classification of calcite veins Attempts to classify calcite veins rely on fracture types, geometry and morphology of veinlets, mineralogy and cement texture of filling as well as the timing and age relationships of veins. Four types of tensional fracture veins in outcrop and hand specimens have been distinguished (Logan and Semeniuk 1976): (1) parallel sheet (a single subparallel vein set), (2) rhomboidal and (3) rectangular (conjugate veins cutting the host rock idens appearing as rhomboid to rectangular in cross section), and (4) breccioid (caused by irregular and numerous conjugate veins). Misik (1966) distinguished separate veinlets formed by tension from those formed by compression using filling and relationships with microstylolites. He drew attention to a specific veinlet type characterized by a set of thin parallel veinlets separated by micrite and believed to have been formed by repeated cracking and recrystallization (Misik 1998).
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Microfactures
5.3.2.2 Descriptive Criteria of Calcite-Filled Microfractures The following criteria are useful in describing and discussing of calcite veins: • Spatial relations of fractures and grains (Simmons and Richter 1976; Kranz 1983): (a) more crumbly fractures (breaking around solid particles, intergranular and intercrystalline cracks), or grain boundary cracks associated with grain boundaries, (b) sharp-edged fractures (breaking through particles and intragranular and intracrystalline cracks extending from a grain boundary crossing into one or more other grains, Pl. 25/1). Sedimentary texture and composition seem to be the essential controls on the main fracture types of carbonate rocks. • Occurrence of single microfractures or fracture sets (composite crack-seal veins, Pl. 25/1) characterized by the consistency of orientation. • Morphology of fractures (straight, arched, branching, stepped, displaced, Pl. 25/4, 6). • Existence of offsets on microfractures (indicating shear, Pl. 25/3). • Shape of upper and lower terminations (sharp, pinching out, irregular). • Orientation of calcite veins with regard to bedding planes (vertical, oblique, parallel). • Occurrence of crosscutting fractures (Pl. 25/1, 6). • Distribution, spacing and density of microfractures and larger fractures (Gillespie et al. 1993). Microcrack densities are expressed either as the number of cracks per unit area or per grain. Crack densities increase with increasing stress, the ratio of grain boundary cracks to intragranular cracks decreases at higher applied loads. These generalized density patterns, however, observed predominantly in non-carbonate rocks, are not always relevant for carbonate rocks because cracking can be surpressed at higher pressures to the benefit of twinning. • Width and length of microfractures. • Filling and sealing. Fracture filling should be described with regard to (a) mineralogy, (b) inclusion bands (Ramsay and Huber 1987), (c) cement types (Muchez et al. 1994), (d) timing of filling generations, (e) mode of vein-growth by various crack-seal mechanisms (e.g. crack-seal fiber growth due to syntaxial overgrowth of grains in microcrack walls and oriented or random nucleation and initial growth of crystals: Cox and Etheridge 1983; Ramsay 1980; solution cleavage, Alvarez et al. 1978), and (f) dynamic effects (e.g. twinning, curving of twin lamellae, undulose extinction). • Spatial relations of microfractures with sedimentary grains and structures (Pl. 25/2).
• Association with stylolites (Nelson 1981; Beach 1982; Pl. 25/4). • Volumetric importance of calcite veins related to the total rock. • Differences in the occurrence of veins in various beds. • Relation of vein abundance and bed thickness (fracture density can have an inverse logarithmic relation to bed thickness, regardless of structural setting (McQuillan 1973). • Relative age of joints and calcite-filled fractures. Conventional analysis of geological fracture data considers the orientation and distribution of fractures measured on sample line/lines oriented normal to the fracture strike and shown by stereonet projections. Methods are described by McQuillan (1973), Huang and Angelier (1989) and Gillespie et al. (1993). Microfracture density can be expressed by (a) the number of fractures intersecting a line in thin section, hand specimen or core, or (b) the number of vertically and obliquely oriented calcite veins related to a defined thin-section area. Geochemical and mineralogical criteria Microfacies thin sections allow only a first evaluation of calcite vein types, distribution and frequency. Detailed analysis of fracture fillings demands that fluid inclusions, isotopic composition and other geochemical criteria as well as SEM and TEM be investigated (Marshall 1982; Dietrich and McKenzie 1983; Ramsay and Huber 1987; Fairbairn and Ferguson 1992; Migaszewski et al. 1998; Suchy et al. 2000).
5.3.2.3 Significance of Microfractures in Carbonate Rocks Microfractures have great potential for evaluating the postsedimentary history of carbonate rocks. The following points are of particular interest: • Larger and smaller fractures in subsurface rocks are known to impart the permeability necessary for the economic recovery of hydrocarbons from reservoir rocks, and some fractures provide a major part of reservoir storage capacity (Drummond 1964; Narr and Currie 1982; Kulander et al. 1990). Fracturing of carbonate rocks enhances permeability and porosity, both of which are important in the reservoir potential. Many large hydrocarbon reservoirs are connected with fractures (e.g. Longman 1985; Roehl and Weinbrandt 1985; McQuillan 1985; Irish and Kempthorne 1987).
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• Details of small-scale fracturing assist in understanding relationships between tectonic structures and fracture distribution and their origin within a regional geological frame (Flügel and Kirchmayer 1962; Logan and Semeniuk 1976; McQuillan 1985). • Mineralogical and geochemical data of microfissure fillings associated with larger joints and major faults add to the explanation of the course of mineralization and ore formation (Fairnbairn and Ferguson 1992). • Abundance of calcite-filled veins influences technological properties (hardness, strength, breakage, slake durability and weathering) of carbonate rocks.
Microfractures and Calcite Veins
Basics: Microfractures and ‘calcite veins’ in carbonate rocks Al-Aasm, I.S., Coniglio, M., Desrochers, A. (1995): Formation of complex fibrous calcite veins in Upper Triassic strata of the Wrangellia Terrain, British Columbia, Canada. – Sed. Geol., 100, 81-95 Antonellini, M., Mollema, P.N. (2000): A natural analog for a fractured and faulted reservoir in dolomite: Triassic Sella Group, Northern Italy. – Amer. Ass. Petrol. Geol. Bull., 84, 314-344 Boukadi, N., El Ayeb, S., Kjarbachi, S. (2000): Analyse quantitative de la fractuation des calcaires yprésiens en Tunisie: l’exemple de Jebel Ousselat. – Bull. Soc. géol. France, 171, 309-317
± Fractures can raise the permeability of hydrocarbon reservoir rocks. Mineralogical and geochemical data of microfractures assist in explaining ore formation. The abundance and type of calcite-filled veins influences the technological properties of carbonate rocks (e.g. strength, breakage, weathering). Many limestones exhibit millimeter- to decimeter-sized calcite- or sediment-filled microfractures, commonly called ‘calcite veins’. These structures are part of large-scaled fracture systems affecting the rock together with joints and faults. Although microfractures can be understood genetically only in the context of a regional structural analysis, the relationships between fractures, faults and folds, ‘calcite veins’ should not be overlooked in microfacies studies because of their importance in deciphering the postsedimentary history of the rocks, including the tectonic activities, burial diagenesis, fluid migrations, reservoir potential and mechanical properties of carbonate rocks. Useful descriptive criteria of microfractures are (a) spatial relations of fractures and grains (–> 1), (b) occurrence of isolated microfractures or fracture sets, (c) morphology (e.g. straight, arched, branching, –> 6), (d) offsets (–> 3), (e) shape of terminations (–> 4, 6), (f) filling and crack-sealing (e.g. mineralogy, cement type, inclusions, growth pattern, dynamic structures) and sealing, (g) orientation with regard to bedding planes, (h) crosscutting (–> 1, 6), (i) distribution, spacing and density, width and length, volumetric importance, association with stylolites (–> 4), and (j) the relation of vein abundance and bed thickness. Fracture systems in carbonate rocks can form during various stages of sedimentation (e.g. oversteepening of semilithified sediment, water escape structures, –> 5) and diagenesis (compaction load, freshwater dissolution), but most calcite veins are due to brittle failure and tectonic fracturing of lithified carbonates caused by stress and shear displacement (–> 2, 3), extensional movements or natural hydraulic fracturing. Attempts to classify calcite veins rely on fracture types, geometry and morphology of veinlets, mineralogy and cement texture of filling (–> 2) as well as the timing and age relationships of the microcracks. 1 Microfracture sets. Radiolaria wackestone. Most larger radiolarians are crossed by calcite-filled cracks. Early Cretaceous (Aptychus beds): Schrecksee, Allgäu, Germany. 2 Shear fractures. Fusulinid foraminifera (Neoschwagerina), displaced along calcite cement-filled shear systems. Note the difference in cement fillings. Middle Permian: Straza quarry, Bled, Slovenia. 3 Microcracks with offsets. Bioclastic mudstone. Late Triassic (Kössen beds): Schrecksee, Allgäu, Germany. 4 Burrowed lime mudstone, crossed by crosscutting calcite veins associated with stylolite seams (arrows). Note the interconnected burrow network (dark appearance) created by sediment-inhabiting organisms. Burrowing took place in aerobe environment. Stylolites are often associated with small-scale calcite-filled fracture sets, which can be used as paleostress indicators. Early Jurassic (Liasbasiskalk): Schrecksee, Allgäu, Germany. 5 Early diagenetic fractures. Vertical microcracks caused by dewatering of the sediment. The rock is characterized by wavy to crinkled fine laminae composed of couplets of mudstone layers (dark) and layers with quartz grains (light). These couplets are known from peritidal settings and muddy lagoonal environments. Late Triassic (Kössen beds): Schrecksee, Germany. 6 Bedded lagoonal fenestral bindstone. Note the differences in shape, thickness and terminations of calcite veins. Middle Triassic (Wetterstein limestone): Schrecksee, Allgäu, Germany.
Plate 25: Microfracures and Calite Veins
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Fairbairn, R.A., Ferguson, J. (1992): The characterisation of calcite-filled fractures from the Northern Pennine orefield. – Proc. Yorkshire Geol. Soc., 49, 117-123 Geiser, P.A., Sansone, S. (1981): Joints, microfractures and the formation of solution cleavage in limestone. – Geology, 9, 280-285 Gillespie, P.A., Howard, C.B., Walsh, J.J., Watterson, J. (1993): Measurement and characterisation of spatial distribution of fractures. – Tectonophysics, 226, 113-141 Irish, J.P.R., Kempthorne, R.H. (1987): The study of natural fractures in a reef complex, Norman Wells Oil field, Canada. – In: Tillman, R.W., Weber, K.J. (eds.): Reservoir sedimentology. – Soc. Econ. Paleont. Min. Spec. Publ., 40, 153-167 Kranz, R.L. (1983): Microcracks in rocks: a review. – Tectonophysics, 100, 449-480 Logan, B.W., Semeniuk, V. (1976): Dynamic metamorphism, process and products in Devonian carbonate rocks: Canning Basin, Western Australia. – Western Australian Geological Society, Spec. Publ., 6, 138 pp. Misik, M. (1966): Tentative classification of veinlets in limestones. – Geol. Sborn. Slov. Akad. Vied, 17, 337-344 Misik, M. (1998): Peculiar types of thin veins in the Mesozoic carbonates and silicates of the Western Carpathians. – Geologica Carpathica, 49, 271-298 Narr, W., Burruss, R.C. (1984): Origin of reservoir fractures in Little Knife field, North Dakota. – Amer. Ass. Petrol. Geol. Bull., 68, 1087-1100 Ramsey, J.G., Huber, M.I. (1987): The techniques of modern structural geology. Vol. 2: Folds and fractures. – 700 pp., London (Academic Press) Suchy, V., Hejlen, W., Muchez, Ph., Dobes, P., Hladikova, J., Jackova, I., Safanda, J., Zeman, A. (2000): Geochemical study of calcite veins in the Silurian and Devonian of the Barrandian Basin (Czech Republic): evidence for widespread post-Variscan fluid flow in the central part of the Bohemian Massiv. – Sed. Geol., 131, 201-219 Troalen, J.-P. (1993): Une méthode indirecte de quantification de l’évolution de la microfissuration thermique dans des roches carbonatées. – Engineering Geology, 33, 189199 Vennin, E., Moreno-Eiris, E., Perejon, A., Alvaro, J.J. (2001): Fracturación sinsedimentaria y diagenesis precoz en las bioconstrucciones del Cámbrico inferior di Alconera (Ossa-Morena). – Rev. Soc. Geol. España, 14, 75-88 Further reading: K071
5.3.3 Carbonate Breccias and Conglomerates Carbonate breccias and conglomerates are sedimentary rocks characterized by limestone or dolomite fragments embedded in a groundmass consisting of fine-grained matrix or/and cements. Breccias are characterized by abundant angular clasts (usually > 50%), conglomerates by abundant well-rounded clasts of pebble- or gravel size. The word breccia originated in Italy where it has long been used for the crushed rubble used in walls.
Breccias
The Latin word conglomeratus means heaped, rolled and pressed together. The study of breccias and conglomerates requires integrating field observations and laboratory data. Microfacies analysis provides a powerful means of deciphering the provenance of the clasts and understanding the relationships between depositional and tectonic controls on recycled sediments.
5.3.3.1 Terminology A variety of terms is applied to rocks composed of carbonate clasts. Most terms recognize the rock to be genetically interpretated. Breccias related to seismic events (see Sect. 12.4) and impact breccias caused by meteorites will be discussed later. Impact breccias are generally characterized by strongly fractured clasts and a large amount of groundmass consisting of fragmented particles. The rock is twisted, sheared and heavily veined. Cement filled vugs are common (Martinez et al. 1994; Warme and Kuehner 1998). Other terms focus on the dimensions of the clasts, which vary widely: Microbreccias: containing mm- and cm-sized carbonate clasts are often described within the framework of common limestone classifications using the terms intra- or lithoclast, and indicating size ranges by the use of ‘calcirudite’, ‘rudstone’ or ‘floatstone’ (e.g. ‘polymict intracalcirudite’). The boundary between lithoclast limestones and breccias is artificial, some authors propose a boundary of about 2 mm. Some microbreccias are extremely diverse with regard to the microfacies types of the carbonate clasts. Auley-Charles (1967) for example, reported a Miocene microbreccia consisting of more than 30 limestone types. Microbreccias are common constituents of slope deposits (Pl. 114/1, 3) or represent distal debris-flow deposits (Pl. 26/4). Megabreccia are coarse matrix-supported breccias containing individual angular blocks up to a meter to more than 100 meters in size. The breccias result from sedimentary and tectonic processes and develop downslope by large thrusts connected with gravitational slidings, e.g. mega-debris flow deposits (Cook et al. 1972; Mountjoy et al. 1972; Leigh and Hartley (1992). They are formed on a broad range of slope angles and contains coarse-grained as well finer-grained breccias. The latter represent a specific Standard Microfacies Type (Pl. 114/1, 3). Explanations for the formation of megabreccias are the release of pore-water overpressure of confined aquifer horizons beneath the sea floor (particularly during falls of sea level) or a catastrophic collapse of high-angle oversteepened carbonate plat-
Describing Breccias
form margins (Spence and Tucker 1997). Diagnostic criteria of megabreccias are (a) the association of megabreccias with various types of mass-flow deposits, (b) mixtures of extrabasinal clasts, intrabasinal clasts and tabular intrabasinal blocks with dimensions of tens of meters to several hundreds of meters, and (c) the common disorientation of geopetal fabrics within blocks (Pl. 26/7) or from one block to another block. Spectacular examples are known from toe-of-slope deposits (e.g. Gullo and Vitale 1987; Flügel et al. 1991; Haas 1999) and in accretionary settings (Sano and Tamada 1994). Decameter-sized constituents of megabreccias should not be mistaken for autochthonous formations. The timing of breccia or conglomerate formation can be indicated by the adjective ‘intraformational’. Intraformational breccias are formed by synsedimentary processes and consist of clasts believed to have formed within the depositional basin. The clasts originate from penecontemporaneous reworking and redeposition fairly close to the site of formation. Breakingup takes place subaerially (e.g. drying of mud on tidal flats) or in a submarine setting by tidal currents (Pl. 132/4), storm waves (Pl. 132/2), or gravity flows (Walcott 1894; Assaruri and Langbein 1987).
5.3.3.2 How to Describe Carbonate Breccias? Basic criteria relevant to the microfacies analysis of carbonates breccias include: Field relationships Geometry, thickness and lateral extension: Geometry as well as local or regional extension of breccia bodies are crucial in understanding breccia genesis. Shallow-marine breccia and carbonate conglomerates are characterized by: individual layers usually a few decimeters thick, discontinuous layers ranging from meter-sized lenses to a lateral extension of many tens of kilometers (Grotzinger 1986). Mass-flow breccias can extend laterally several hundreds of meters to tens of kilometers. Stratification, grading and sedimentary structures: Breccias are massive or stratified. Stratification should be described in terms of distinctness, dimensions and sedimentary structures (cross-stratification, high-angle edgewise structures, common in modern beach deposits, occurring in lenses and pockets). Edgewise conglomerates have various origins, they occur as erosional scours in supratidal flats, as discontinuous layers near linear mounds, and on beach ridges along eroded shorelines. Breccia grading can be normal or inverse.
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Boundaries and relationships to adjacent rocks: The genetic interpretation of breccias needs data concerning the basal and upper boundaries and the relations to the underlying and adjacent rocks. Lower contacts can be erosional and sharp, gradual or passing into fractures and fissures (e.g. internal breccias). Breccias formed by sliding and slumping often exhibit multiple brecciation phases and reverse gradation of the clasts. Differences in the color of clasts and groundmass: The color of the clasts and the fine-grained matrix is important in recognizing extra- or intrabasinal origins of the material. Clasts Lithologic composition: The composition of clasts is differentiated with regard to (a) homogeneity or differences in lithology (monomict and polymict breccias, pure carbonate breccias, lithologically mixed breccias), and (b) the quantitative proportions of different lithologies. Lithology is commonly described in terms of field observations only (e.g. limestone, dolomite or chert clasts) but should be evaluated by thin sections. Commonly, very small clasts are attributed to the matrix during field work, but thin sections might provide evidence that precisely these clasts are the only relicts of stratigraphically older rocks that have suffered repeated erosion and redeposition. Quantitative analysis needs statistical sampling based on line-, ribbon-, area- or point-counting (Sect. 6.2.1). Quantification must consider not only the frequency of lithologic types but also the frequency of clasts which catch the eye by their particular rock color, rounding or size. Compositional clast analysis offers key information on (a) extra- or intrabasinal sources, (b) the environmental controls in the provenance area, (c) multiple redeposition, (d) depositional or non-depositional formation of the breccias and (e) the site of breccia formation (terrestrial, marginal-marine, marine). Fabric: The fabric of breccias is characterized by matrix-support or clast-support. Matrix-supported breccias exhibit clasts floating within fine-grained matrix (Pl. 26/2) or a calcarenitic groundmass. Clast-supported breccias consist of clasts supporting each other (Pl. 26/ 3). This differentiation resembles the distinction between mud-support and grain-support used to characterize depositional structures of limestones (Dunham 1962), but should not be regarded as strictly genetic, because matrix- and clast-support fabrics are caused by depositional as well as non-depositional processes (e.g. solution, brecciation, tectonic shear). Size and sorting: The size of carbonate clasts is highly variable. Clasts can range from arenite to boulder size. Shallow-marine breccias and mass-flow brec-
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cias formed in upper slope environments often exhibit a bimodal grain size composition, with larger clasts and interstitial well-sorted sand-sized particles (peloids, ooids, skeletal grains, Pl. 26/7). Size and sorting are essential criteria in discriminating genetic conglomerate types. Shape and surface of carbonate clasts: The form may vary from sharply angular to more or less rounded, the shape can be isometric, platy, or tabular. Many round
Carbonate Breccias
conglomerate pebbles exhibit biogenic encrustations and borings indicating interruptions and resting times during transport processes (Pl. 26/7). Roundness: Per definition breccia clasts are angular, but carbonate breccias display all transitions possible from predominantly angular to variously rounded clasts, which relate the latter samples to conglomerate categories (see Pl. 26/3). The roundness of breccia clasts is determined with the help of comparison charts that
Plate 26 Carbonate Breccias Carbonate breccias and conglomerates, made of angular or rounded rock fragments embedded within a matrix and/or cemented by various minerals, are common in marine and non-marine sequences. Integration of field and laboratory data is necessary for classifying and interpreting these rocks. In polymict breccia, clasts and matrix are not genetically related, and result from transport and depositional processes (e.g. submarine landslide, fluvial transport and deposition, subaerial and subaqueous debris flows (–> 3), slumping and volcanic phenomenons (–> 2). Mechanisms triggering mobilization are gravity, earthquake and water loading that induce overpressure, storms (producing flat-pebble conglomerates), tides, oversteepening of slopes (–> 7), gas hydrate destabilization, and glaciation processes. In autoclastic or intraformational breccias, clasts and matrix are genetically related. In terrestrial or meteoric-vadose environments, these breccias result from in situ deformation or reworking processes (dehydration, karstification). 1 Polymict stylobreccia consisting of tightly packed limestone clasts. The scarce marly matrix is a residue product of pressure solution. The breccia represents a massflow deposit formed by submarine gravity flows and modified by subsequent pressure-solution. Rounding of clasts is caused by burial solution, not by transport. Arrows point to solution seams. Note the different microfacies of the clasts: The dark clasts are pelagic bioclastic wackestone with filaments, ostracods and radiolaria. Light clasts are bioclastic packstones. Early Jurassic (Adnet Limestone): Adnet near Salzburg, Austria. 2 Matrix-supported agglomerate (volcanic breccia) formed by submarine volcanic explosion followed by admixture and massflow transport of volcanic clasts (V) and rounded carbonate extraclasts (E). Middle Triassic (Ladinian): Mt. Agnello, Southern Alps, Italy. 3 Clast-supported monomict limestone breccia. The clasts exhibit different roundness/sphericity values: subangular, subrounded and rounded clasts (fields B, E and I in the Krumbein chart, cf. Fig. 4.30). The clasts are pelagic bioclastic wackestones and packstones. Calcite-filled microfractures are limited to the clasts indicating fracturing prior to brecciation. Interclast voids are filled with several generations of marine fibrous and blocky calcite cements. The breccia is explained as mass-flow breccia resulting from submarine slides of semiconsolidated pelagic sediments on a steep slope of a drowned Triassic reef, triggered by tectonic activities (Hudson and Jenkyns 1969; Böhm et al. 1995). Because of the attractive appearance of the red limestone clasts within the white sparry calcite, the breccia has been widely used as decorative stone. Early Jurassic (‘Scheck breccia’, Piensbachian/Toarcien): Adnet near Salzburg, Austria. 4 Polymict clast-supported microbreccia (sandy lithoclast packstone) consisting of angular (gray) dolomite clasts, (dark) limestones clasts and (white) quartz grains. Note the absence of matrix. Debris-flow deposit. Early Cretaceous: Mosaic material of a Roman building, Kraiburg, southern Bavaria, Germany. 5 Monomict tectonic fault breccia. Note the angularity and the complex and strong fracturing of individual clasts as well as the whole rock. Microfractures differ in width, filling and orientation. Some calcite veins clearly postdate brecciation. The matrix consists of small broken and sheared rock fragments bound together by sparry calcite. Late Triassic (Dachstein Limestone, Norian): Northern Alps, Austria. 6 Polymict breccia consisting of extraclasts (left: shallow-marine algal wackestone with dasyclads - AW, bottom right: microbioclastic packstone - P) and alveolinid foraminifera. Interpretation: Massflow deposit. Redeposition of shallowramp bioclasts and lithoclasts in deeper ramp areas caused by regressive events (Kulbrok 1995). Early Tertiary (Eocene): St. Paul Monastery, Wadi Araba, eastern Egyptian Desert. 7 Polymict massflow breccia. The material was eroded from reefs formed on the upper slope of a platform margin. The sample is part of a megabreccia consisting of meter-sized reef- and slope-derived blocks.. Deposition took place in a proximal position on the foreslope. Most clasts are boundstones except for the crinoid packstone (P) representing slope facies. Note biogenic encrustations (arrows) indicating breaks during the transport on the slope. GP: Upside-down geopetal structure (arrow). The interstitial grainstone ‘matrix’ (GM) represents platform-derived material infilled subsequent to the deposition of the clasts. Fracturing postdates brecciation. Middle Triassic (Ladinian): Seiser Alm, Dolomites, Italy.
Plate 26: Carbonate Breccias
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consider both roundness and shape (Sect. 4.3.1, Pl. 26/ 3), or as the percentage of clasts with at least one rounded edge (roundness number: Richter and Füchtbauer 1981). Breccias formed in place commonly consist of high proportions of angular or only partly subrounded clasts (e.g. internal breccias, karst breccias, pedogenic breccias, Pl. 128/2, Pl. 129/1). Roundness can increase with transport (e.g. in mass-flow breccias) and by multiple reworking of breccia beds (Buggisch and Flügel 1980). However note, that rounded carbonate rock fragments are not necessarily products of abrasion during transport but can also be caused by partial dissolution (Pl. 26/1) or replacement processes after deposition. In addition, organic encrustations will modify the shape of clasts and transform angular rock fragments into round boulders. Fitting: Fitting (Richter and Füchtbauer 1981) describes the grain-to-grain relations of clasts – the degree and percentage to which clast boundaries match those of adjacent clasts (Pl. 26/1, 5). In measurements, each clast is counted only once. Differences in fitting are powerful criteria in distinguishing genetically different breccia types (Fig. 5.14). Orientation: The orientation of clasts may be absent or parallel to bedding or stress, and increasingly or decreasingly parallel to the transport direction. Flatlying orientation in shallow-marine carbonate breccias is sometimes caused by compaction. Current-controlled deposition is manifested by the imbrication of clasts and partial infilling of interstices between the clasts with sand after the deposition of the gravel framework. However, it should be taken into account that ‘pseudo-imbrication’ also occur at the foresets of large-scale crossstratification. Fracturing or veining of clasts occurs prior to brecciation, during the process of brecciation, or subsequent to brecciation (Pl. 26/3, 5, 7). Fractures and veins within clasts of many mass-flow breccia and reef breccia predate fracturing and veining whereas fracturing during brecciation is common in caliche and solution breccia. Internal breccias and shear breccias exhibit abundant and superposed generations of post-brecciation fractures and veins (Pl. 26/5). Internal breccias often show inherited calcite-filled fractures. Microfractures should be described by abundance, orientation, size and filling (see Sect. 5.3.2.2). Clast/matrix boundary: Most boundaries are sharp (e.g. Pl. 26/3, 5) except for diagenetic pseudobreccias and solution-evaporite-collapse breccias, which may exhibit irregular replacement margins. Microfacies: Sampling for microfacies studies of clasts must consider the size, roundness and shape, packing of clasts as well as conspicuous sedimentary
Groundmass of Breccias
Fig. 5.14. Relations between fitting of clasts and the amount of groundmass (matrix + cement). After Füchtbauer and Richter (1983).
structures (e.g. stratification). Carbonate clasts of breccias and conglomerates are classified according to depositional texture, composition and fossils. The frequency and size of clasts give some indication of the proximity or distality of source areas. The relative proportions of microfacies types should also be estimated or measured (e.g. counting all clasts within a defined thinsection area). A comparison of standard microfacies types (see Sect. 14.3) of breccia clasts with regional or generalized facies models may provide clues to the provenance area. Fossiliferous clasts must be tested with an eye to age determinations. Groundmass Abundance of the matrix and clast/matrix ratio: Matrix can be abundant, scarce or even missing due to diagenetic solution (e.g. stylobreccia). Differentiation of breccia types and genetic interpretation requires the quantitative data of the amount of matrix and the clast/ matrix ratio. These data are determined in the field (sometimes also in large rock samples and thin sections) using the counting methods mentioned in conjunction with the compositional clast analysis, or estimating the frequency of clasts. Matrix composition: The matrix of carbonate breccias consists of sediment or rock detritus formed by the destruction and attrition of rock material. The descriptive criteria of the sedimentary matrix are color, mineralogy, uniformity, texture (e.g. micrite, calcilutite or calcarenite matrix), and small-scaled sedimentary structures (e.g. grading, lamination). Finegrained matrix is represented by (a) sediment deposited between clasts coeval with the deposition of the
Types of Breccias
clasts Pl. 26/7), (b) sediment infilled (‘sucked in’) after the deposition of the clasts, (c) the residual matrix of beds which have been affected by solution (stylobreccias, Pl. 26/1), or (d) fine erosional detritus consisting of rock debris (Pl. 26/5) The major aim of the study of matrix composition is to find out if clasts and matrix are different with regard to the environment and timing of breccia formation. Cement minerals: Common chemical precipitates of carbonate breccias are calcite cements. The mineralogical composition, cement types and spatial distribution of pore-filling carbonate cements are important criteria for understanding the course of the diagenetic and tectonic history of breccias. Fossils The fossils occurring within the matrix and in the clasts are the main criteria in (a) reconstructing the environment and facies patterns in the source area of depositional breccias and conglomerates (Pl. 27), and (b) evaluating age and timing of breccia and conglomerate formation. In basin analysis, the microfacies data of isolated pebbles and of breccia clasts assists significantly in recognizing erosion events connected with the breakdown of carbonate platforms (e.g. Flügel and Kraus 1988) and of allochthonous sedimentation on slopes and in deeper basins (e.g. Herbig 1984). Abundant stratigraphically useful microfossils seen in microfacies thin sections of Phanerozoic conglomerates and breccias include foraminifera, calcareous algae and many microproblematica (see Chap. 10). The combined use of thin-section biota and microfossils recognized in solution residues of limestone pebbles (e.g. conodonts) is highly recommended (e.g. Pohler and James 1989). Because fossils can be reworked several times and set free from surrounding rock material, difficulties can arise in evaluating ‘isolated’ microfossils found within the matrix. Taphonomic criteria (preservation) should help overcome these pitfalls (Pl. 9/7).
5.3.3.3 Carbonate Breccia Types: Origin, Classification, Criteria Carbonate breccias are of different origins. Their classification is essentially genetic and differs conceptually from those of limestones. Classification systems are based either on descriptive criteria or inferred from brecciation processes (Spence and Tucker 1997).
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The following summary of common breccia types is based on the suggestions of Norton (1917), Blount and Moore (1969), Keith and Friedman (1977) and Richter and Füchtbauer (1981). Major genetic categories include depositional, non-depositional, tectonic and diagenetic breccias (Fig. 5.16). Depositional, nondepositional and tectonic breccias as well as conglomerates differ in composition and the role of transport processes (Fig. 5.15). Depositional breccias Mass-flow breccia Field observations: Bedded, primary sedimentary structures, stratigraphically concordant, stratabound. Variable thickness, large lateral extension. Lens-shaped bodies, massive laterally extended beds. Variable grading (inverse and normal). Upper and lower boundary sharp or undulated, lower boundary often erosive or channelized. Clasts: Lithologic composition: often polymict (Fig. 5.21). Fabric usually clast-supported, no primary fitting but common boundaries due to pressure solution (Pl. 26/1). Close to open packing. Variable shape, all degrees of roundness, subangular and rounded clasts common (Pl. 26/1, 3), Pl. 27/1). Variable size, poor to moderate sorting (Pl. 114/1, 3). Fracturing or veining of clasts (if present) predates breccia formation. Clast/ matrix boundary sharp. Microfacies variability very
Fig. 5.15. Relations between the composition of breccias and conglomerates and the relative motion of clasts. Note the differences between non-depositional breccias (shrinkage and solution/collapse breccias) and tectonic breccias (internal, fissure-fill and shear breccias) as compared with depositional breccias. Conglomerates, including synsedimentary soft-clast conglomerates, appear to be distinctly separated from breccias, but may act as precursor stages of breccia formation. Modified from Richter and Füchtbauer (1981).
234 Depositional breccias resulting from the deposition of eroded carbonate material)
Types of Breccias Mass-flow breccia: Breccia originating from the downslope transport of shallow-marine and slope sediments moving under the force of gravity. Includes breccias formed by slumps and slides, debris flows, grain flows, and turbidity flows. Submarine rockfall breccia: Mass-flow breccia formed by the accumulation of coarse, angular rock fragments derived by falling from a cliff, escarpment or steep rocky slope. Peritidal and shallow-marine breccia: Breccia formed by synsedimentary deposition of eroded peritidal, shallow subtidal as well as subaerial carbonates, often related to storm events. Deposition takes place in inter- and supratidal settings, and at the beach. Forereef breccia: Breccia deposited on the seaward slope of high-energy reefs. Consisting of eroded reef material and remains of organisms living in the reef or on the foreslope.
Non-depositional breccias (resulting from in-place dissolution)
Caliche breccia: Breccia formed by in-situ brecciation in arid and semiarid climates, controlled by soil-forming processes, and connected with extensive weathering, erosion, solution and shrinkage. Solution-evaporite-collapse breccia: Breccia formed by collapse of beds subsequent to the removal of soluble material within some beds (e.g. evaporites).
Tectonic breccias (resulting from internal dislocation of carbonate rocks)
Fissure fill breccia: Breccia formed within submarine neptunian dikes or subaerial fissure infills and karst fissures. Internal breccia: Breccia formed by rupture and fracturing of carbonates near the depositional surface. These breccias are products of dilation of slightly lithified limestones caused by tectonics (e.g. hydraulic fracturing, earthquakes). Internal breccias occur in platform and slope carbonates which were brecciated shortly after deposition and before final lithification. Shear breccia: Breccia caused by brittle deformation associated with thrust and sliding displacement.
Diagenetic breccias (resulting from early diagenetic processes)
Pseudobreccia: Mottled limestones and dolomites with breccia-like textures caused by patchy recrystallization and cementation, possibly controlled by the distribution of organic compounds. Stylobreccia: Breccia in which fragments are bound by stylolites. Caused by fracturing of carbonate rocks, accompanied by pressure solution between the fragments of the breccia.
Fig. 5.16. Main types of carbonate breccias. Because many breccias are products of multiple processes, the boundaries between some categories are transitional. The formation of fissure fill breccias, for example, depends on the existence of a tectonic framework, but the composition of the breccias is depositionally controlled.
high to moderate, and may include the entire spectrum of Dunham’s texture types (Fig. 5.21). The clasts represent a mixture of material derived from various shallow-water and slope environments and sometimes also from basinal environments. Groundmass: Fine-grained sedimentary matrix (e.g. pelagic micrite) or matrix brought to the basin together with the clasts. Depending on the depositional pattern, the matrix volume is less,than equal to or greater than that of the clasts. Inter-clast carbonate cements are relatively rare (Pl. 26/3). Fossils: Common and abundant in clasts and within the primary sedimentary matrix (Pl. 9/7), relatively rare in the allochthonous matrix. Case studies: Bektas et al. 1995; Bernoulli et al. 1987; Braithwaite and Heath 1992; Cook 1979; Engel 1970; Flügel et al. 1993; Gawlick 1996; Günther and
Wachsmuth 1969; Leigh and Hartley 1992; Payros et al. 1999; Peryt 1992; Remane 1970; Ruberti 1993; Steiger 1981; Stock 1994; Wächter 1987. Submarine rockfall breccia Field observations: Massive or thick-bedded. Intercalated in bedded limestones and marl/limestone sequences. Thickness ranges between several meters and some tens of meters. Lateral extension can reach more than 100 m. The clasts are often eroded from coastal escarpments or steep submarine slopes. Clasts: Lithological composition often polymict. Limestone clasts may be derived from different lithoor biostratigraphic units. Fabric clast-supported. Close packing. Variable shape, angular and subangular clasts common. Rounded fragments are not caused by abrasion during transport but rather to partial dissolution.
Peritidal Breccias
Variable size, mixtures of m-sized and cm-sized clasts are common. Poor sorting. Clast/matrix boundary sharp. Microfacies variability of clasts high. Groundmass: Matrix common to abundant, consisting of micrite or fine-grained rock debris. Several phases of matrix infilling in fissures within the semilithified breccia blocks are common. Fossils: Common in clasts and as isolated, often only poorly preserved fragments. Case studies: Prtoljab and Glovacju Jernej 1994; Semeniuk and Johnson 1985; Sanders 1997; Thiedig 1975. Peritidal and shallow-marine breccia Field observations: Individual layers usually a few decimeter thick. Layers discontinuous, ranging from meter-sized lenses to a lateral extension of many tens of kilometers. Flat-lying clast orientation, sometimes caused by currents, sometimes caused by compaction. True imbrication patterns are indicated by partial infilling of interstices of gravel clasts with sand after the deposition of the gravel framework. Pseudo-imbrication may occur at the foresets of large-scale cross-stratification. Cross-stratification and high-angle edgewise structures are common. Clasts: Lithological composition monomict, but strong microfacies differentiation. Fabric clast-supported to matrix-supported. Low fitting. Bimodal grain size, with interstitial well-sorted sand-sized grains (peloids, ooids, skeletal grains). Variable shape, many clasts are highly rounded and form conglomerates. These conglomerates comprise flat-pebble conglomerates caused by erosional rip-up of intertidal and supratidal laminites during storm events, and subtidal conglomerates. The latter originate as storm or swell lags in shore-face environments, as lateral accretions at the base of tidal channels (Pl. 132/4), or from a ripup of underlying older tidal flat sediments during rises in sea level (Pl. 132/3). The last case results in a distinct polygenic composition of the conglomerate. Microfacies of clasts highly variable. Constituents include desiccated and often laminated mud clasts derived from modern intertidal and supratidal settings (desiccation breccias), clasts derived from subaerial or submarine cemented hardground (e.g. lithoclasts consisting of reworked supratidal crusts or of blackened soil crusts), and grainstone intraclasts. Clast/matrix boundary usually sharp. Groundmass: Fine-grained carbonate matrix, often dolomitized. Fossils: Rare in supra- and intertidal clasts, but common in subtidal clasts. Matrix with ostracods, foraminifera and mollusk shells.
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Case studies: Aigner 1985; Craft and Bridge 1987; Einsele and Seilacher 1982; Shinn 1986; Waters et al. 1989; Demicco and Hardie 1994. Forereef breccia Field observations: Meter- to tens of meter thick units consisting of coarsely bedded or massive breccias. Lateral extension tens to hundreds of meters. Considerable differences between proximal and distal parts of the breccia with regard to size and sorting of clasts, clast : matrix ratio, and the occurrence of inter-clast cements. Clasts: Lithological composition monomict but strong microfacies differentiation. Fabric clast-supported or mud-supported depending on the position on the foreslope. Size strongly variable within millimeter to decameter range. Sorting poor to moderate. Shape variable, angular to apparently well-rounded. The latter clasts may have become round by repeated biogenic encrustations. Low or no fitting. Fracturing or veining of clasts prior to deposition. Clast/matrix boundary sharp. Microfacies variability high because clasts are derived from different parts of the reef complex including reef flat, reef margin and various slope environments. In addition, back-reef sands can be deposited together with forereef clasts. Common microfacies types of the clasts comprise boundstones, bio- and lithoclastic wackestones and floatstones, packstones as well as grainstones and rudstones. Groundmass: Fine-grained, usually poorly sorted matrix (wackestone, packstone, grainstone), sometimes graded. Submarine carbonate cements in interclast voids, commonly several cement phases (Pl. 27/2). Differential submarine cementation can result in downslope displacement of carbonate clasts and sands. Fossils: Reef builders and reef dweller are abundant and often highly diverse within the clasts. Biogenic encrustations and borings are common on clast surfaces. Isolated fossils occur in the interclast groundmass. Cryptic habitats in cavities between reef blocks are characterized by specific microfossil associations. Case studies: Brachert and Dullo 1990, 1994; Enos and Moore 1983; Hopkins 1977; Krebs 1969; Melim and Scholle 1995; Stentoft 1994; Gischler 1995; Vecsei 2001. Non-depositional breccias Caliche breccia Field observations: Related to exposure profiles. Often superimposed on solution-collapse breccias. Bedding dependent on primary clast distribution. Clasts: Strong compositional similarity of limestone clasts and the limestone underlying beds. Polymict
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clasts, representing a mixture of limestone clasts, clay minerals, iron oxides, quartz silt and isolated bioclasts. Solution relics (‘lithorelics’) of the underlying rocks floating within the matrix. Clasts of variable size within millimeter and centimeter range. Greatly variable shape, rarely rounded, but often ‘subrounded’ due to dissolution. Extremely poor sorting. Fabric clast-supported. Fitting often high. Sharp clast/matrix boundary. Circumgranular cracks around the clasts. Clay-filled fractures (crypto- and microkarst). Clast/matrix boundary sharp. Veining common, fracturing uncommon. Microfacies: Variability of clasts high including lime mudstones, floatstones and rudstones.
Non-Depositional Breccias
Groundmass: Fine-grained calcitic or dolomitic matrix abundant. Often clayish commonly reddish matrix, corresponding to the insoluble oxidized residue remaining after the dissolution of underlying limestone. Meteoric-vadose and meteoric calcite cements. Fossils: Rare. Case studies: Semeniuk 1986; Wright and Tucker 1991. Solution-evaporite-collapse breccia Field observations: Thick, discontinuous breccia bodies. Unstratified or only indistinctly stratified. Some gravity segregation. Usually proximal to evaporite pri-
Plate 27 Carbonate Breccia: Provenance Analysis of Clasts Microfacies analysis of the clasts of carbonate breccias has great potential for basin analysis. Compositional studies of the clasts assist in evaluating intensity and timing of erosion processes and in reconstructing ‘lost’ carbonate platforms whose material has been deposited in slope and basin positions (Sect. 16.3). Because the destruction of shallow-marine platforms depends on changes in water depths caused by sea-level fluctuations, clast analysis is a promising method in sequence stratigraphy. The samples shown in this Pl.27/1 and in Fig. 5.21 come from the ‘Tarvis breccia’, a breccia horizon in Early Permian sediments of the Southern Alps which can be followed laterally for a distance of several tens of kilometers. The breccia was studied with regard to the lithologic composition and textural characteristics of the clasts and the matrix, and with respect to the microfacies types and microfossils (Buggisch and Flügel 1980; Flügel 1980). The microfacies types of the clasts reflect the erosion of a sequence of shallow-marine platform carbonates formed in restricted and open shelf lagoons. Platform-margin reef limestones are of subordinate importance. Most clasts were deposited as mass-flow breccias (Fig. 5.21) and rockfall breccias in mixed carbonate-siliciclastic settings. Some carbonate clasts are also associated with volcanic material in lacustrine settings (–> 1). Differences in fitting, size and rounding of the clasts, and the change from carbonate and marly matrix types in the lower parts of the breccia sequence to predominantly siliciclastic matrix types in the upper part indicate a regressive shift from a shallow-marine to a coastal and continental depositional environment accompanied by syntectonic events. 1 Lacustrine breccia. The microfacies of the lithoclasts reflects the erosion of shallow-water marine platform carbonates. The bioclastic grainstone clast (left) with dasyclad green algae (Connexia carniapulchra, Mizzia sp., Gyroprella sp.) and fusulinids (Pseudofusulina) was derived from carbonates formed in a protected high-energy environment within a shelflagoon. The high primary interparticle porosity was reduced by submarine radiaxial cements. The bioclastic packstone (right) yields poorly sorted skeletal grains (echinoderms, foraminifera) which underwent selective dolomitization prior to the erosion and deposition of the lithoclast. Fusulinid foraminifera (Pseudoschwagerina - center and Pseudofusulina right) attest to an Early Permian age. Both clasts were deposited together with volcanic pebbles (VP) within a lacustrine mud represented by a dark homogeneous micrite with calcite-filled dewatering structures (arrows). During burial the breccia was affected by strong pressure-solution resulting in distinct penetration of clasts (CP). Early Permian Tarvis Breccia: Sexten Dolomites, Southern Alps, Italy. 2 Clast-supported forereef breccia. Lithoclast rudstone. Poorly sorted angular, subangular and subrounded clasts of lithified reef carbonates as well as some isolated reef organisms are cemented by submarine fibrous and radiaxial calcite cements (RC), characteristic for upper slope deposits. Roundness/sphericity corresponds to the fields B, D, J and K of the Krumbein visual chart (Fig. 4.30). Reef clasts exhibit diverse microfacies types, indicating erosion in different reef crest areas. Most limestone clasts were eroded from central parts of the reef, others (bottom right) represent sediments originally deposited within cryptic environments of reef cavities (indicated by fine-grained peloidal texture and diagnostic foraminifera). Larger fossils are redeposited sponges (Cheilosporites -S). Erosion of sediments originally deposited in low-energy protected environments (deeper reef crest or slope) is indicated by the specific biotic association of sphinctozoid sponges (S), spongiomorphid sponges (SP), coral fragments (C) and solenoporacean red algae (SO). Note the microbial coating of frame-building organisms (arrows, cf. Pl. 50/1). Late Triassic (Dachstein Reef Limestone, Norian): Gosaukamm, Austria.
Plate 27: Provenance of Clasts in Breccias
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mary strata or associated with vanished evaporites (see Sect. 12.1.7 for criteria). Clasts: Often polymict. Commonly dolomite or dedolomite clasts. Weathering of dolomite clasts results in the formation of a cellular structure (Rauhwacken). Unsorted, variously sized, usually angular and subangular clasts. Usually clast-supported fabric. Variable fitting. Clast-matrix boundary sharp except at the replacement margins of some clasts. Often densely packed. Very common pre-breccia veining, fracturing common. Clast microfacies types include lime and dolomitic mudstones with evidence of former evaporites, and replacement dolomites and dedolomites. Groundmass: Fine-grained matrix corresponds to a microbreccia resulting from internal brecciation of coarser clasts, or to microcrystalline calcite. Fossils: Absent or very rare. Case studies: Assereto and Kendall 1977; Beales and Oldershaw 1969; Middleton 1961; Morrow 1982; Karakitsios and Pomoni-Papaioanou 1998; PomoniPapaioanou and Dornsiepen 1987; Stanton 1966; Visser 1986. Tectonic breccias Movement along a fault surface at high levels within the crust generally results in the fracturing and breaking of wall rock fragments and crystals due to cataclasis. Products of this process are fault breccias (Pl. 26/5), microbreccias (with largest fragments < 1.0 mm) and fine-grained gouges (consisting of fragments diminuated to 0.1–100 µm). Many tectonic breccias are calcite cemented. The terminology of fault-related breccias was discussed by Wise et al. (1984). Fissures filled with breccias are often related to synsedimentary tectonics. Other tectonically controlled breccias are internal breccias and shear breccias formed. Marine fissure-fill breccias Field observations: Fissure-fill breccias occur in karst cavities (see Sect. 15.2.1) and in vertical and oblique submarine dikes. Commonly in strata underlying strata with internal breccia. Clasts: Monomict. Angular or subrounded. Clasts usually small, cm-sized. Unsorted. Clast/matrix boundary sharp. Poor fitting. Microfacies: Lithoclastic floatstones. Groundmass: High matrix percentage consisting of recrystallized fine-grained calcite. Fossils: Broken and worn mollusk shells and echinoderms may be associated with breccia clasts. Case studies: Schlager 1969; Richter and Füchtbauer 1981.
Internal Breccias
Internal breccias Field observations: Strata-bound, but laterally discontinuous. Many internal breccias consist of shallowmarine platform carbonate rocks overlain by basinal sediments (e.g. red pelagic limestones). A characteristic sequence (Fig. 5.17) includes from base to top (a) shallow-water limestones, locally with calcite-cemented joint, (b) shallow-water with fissures, filled with a breccia of the same lithology and exhibiting a red carbonate matrix, (c) internal breccia consisting of shallowwater limestones with a red carbonate matrix filled-in from above, (d) mass-flow breccia consisting of the same shallow-marine limestones and the same red matrix as in (c), and (e) red pelagic limestones. Common in shelf-to-slope transition settings. Clasts: Usually monomict. Angular and subangular clasts. Shape irregular to isometric. High to moderate fitting. Low relative dislocation of clasts. Fracturing of clasts characterized by several superimposed generations. The breccia includes fragments of calcite veins as well as of older internal breccias. Variable microfacies types including all common types of platform carbonates. Groundmass: Matrix sucked in from above, or cement. Variously colored, often red. Fossils: Common within clasts and also in the matrix. Other terms used for breccias: Autoclastic breccia (Fairbridge 1978), dilatation breccia (Roehl 1981). Case studies: Richter and Füchtbauer 1981; Füchtbauer and Richter 1983; Herbig 1985; Wächter 1987. Shear breccia Field observations: Poorly stratified and discontinuous. Shear faces crossing the breccia and crushed clasts.
Fig. 5.17. Schema of the development of internal breccias. Fractured shallow-marine limestones are overlain by internal breccia, followed by mass-flow breccia and pelagic limestones. Modified from Füchtbauer and Richter (1983).
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Diagenetic Breccias
‘Clasts’ and ‘groundmass’: The clast-like mottled areas are of different size, have irregular and transitional outlines and appear poorly sorted. The calcite crystals of the mottles are larger than those of the inter-mottle areas. The proportions of ‘clasts’ and ‘matrix’ differ widely. No fitting and no veined ‘clasts’. The ‘groundmass’ is more frequently affected by late diagenetic dolomitization than the ‘clasts’. This breccia type may appear either ‘clast’ or ‘matrix’ supported. Case studies: Bathurst 1959; Blount and Moore 1969; Solomon 1989. Note that the term ‘pseudo-breccia’ is also applied to breccias consisting of in-place clasts, separated by a network of joints and interpreted as being the result of earthquake loading. Similar breccias can also be caused by wave-induced processes including pseudo-cracking, crack collapse and heterogeneous liquefaction (Bouchette et al. 2001).
Fig. 5.18. Internal breccia. The bivalve coquina consisting of Halobia shells is brecciated and fractured. Note the corresponding boundaries of many ‘clasts’ (white arrows) and calcite-filled cracks. Late Triassic (Hallstatt limestone, Norian): Obersalzberg, Berchtesgaden, Germany. Scale is 1 mm.
Clasts: Usually monomict. Closely packed and dissected by shear planes. Low fitting. Relative dislocation of clasts high. Different generations of fractured sometimes difficult to separate. Flat to lens-shaped clasts. Clasts closely packed. Groundmass: Groundmass consists of cataclastic clasts or cement. Fossils: If present poorly preserved. Other terms used for shear breccias: Tectonic breccias (Blunt and Moore 1969). Case studies: Roehl 1981. Diagenetic breccia Pseudobreccia caused by patchy recrystallization Field observations: Fine-grained bedded dolomites or limestones displaying a distinct irregularly distributed mottled texture, which mimics the appearance of breccia. The mottles differ in darker color from the areas in-between. Known predominantly from wackestones and packstones.
Stylobreccia caused by variable pressure solution Field observations: Massive or indistinctly bedded limestones, often displaying a ‘nodular’ fabric. Clasts: The ‘clasts’ are ‘idens’ (see Sect. 7.5.2) and comprise resedimented lithoclasts and fossils affected by pressure solution (Pl. 26/1). They are separated by thin residual clay seams or microstylolites. Many ‘clasts’ appear subrounded or subangular. Clast sizes highly variable. Fabric appears clast-supported Microfacies types highly diverse. A specific case are styloconglomerates formed by selective pressure solution of lime mudstones (Pl. 47/6). Groundmass: More or less absent or restricted to small amounts of residual micrite. Fossils: Common. Case studies: Logan and Semeniuk 1976; Steiger 1981.
5.3.3.4 Carbonate Conglomerates Most authors distinguish conglomerates sensu strictu (formed by extraclasts) and synsedimentary soft conglomerates (formed by intraclasts). The latter comprise flat-pebble conglomerate produced by erosional ripup of intertidal and supratidal laminites during storm events. Descriptive parameters of conglomerates are fabric, stratification and grading, size and sorting of the pebbles, lithologic uniformity or variability, and the amount of matrix. The composition of the groundmass is variable (carbonate, siliciclastics, matrix and/or cement). Orthoconglomerates are characterized by clast-
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Setting of Carbonate Breccias
Fig. 5.19. Polymict orthoconglomerate consisting of well-rounded limestone pebbles of different age, quartz pebbles (Q) and completely rounded fossils (worn and recrystallized colonial coral: top left). Note the strong difference in the microfacies and the age of the limestone pebbles: Intraclastic grainstone (IG, Triassic), peloidal packstone (PP, Late Cretaceous), argillaceous mudstone (M, Eocene) with some ostracods. The light gray pebbles are dolomite (D, Triassic). The microfacies of the conglomerate indicates reworking of Northern Alpine rocks, the excellent rounding of the limestones points to a short fluvial transport distance of the material into the Molasse Basin. Interparticle pores are filled with granular calcite. Penetration of adjacent clasts (arrows) indicates strong compaction. Early Tertiary (Oberaudorf Formation, Late Eocene): Florianiberg near Oberaudorf, Bavaria, Germany. Scale is 1 mm.
support or by clasts floating within a sandy matrix. The clasts of paraconglomerates float within a muddy matrix. The main criteria used for the quantification of conglomerates are clast size and rock types. The formation of conglomerate is essentially controlled by transport processes. The larger the pebbles of a conglomerate, the higher is the transport energy. Conglomerates become better sorted during transport and the proportion of matrix decreases.
Limestone conglomerates originate in nonmarine and shallow- and deep-marine environments, and are common in fluvial and beach settings (Fig. 5.19). Conglomerates composed entirely of limestone pebbles record conditions of rapid erosion and burial, e.g. along a fault scarp (e.g. erosion of coastal limestone cliffs). Transgression of an advancing sea in marginal marine environments can produce limestone conglomerates consisting of beach gravels. Shallow-marine carbon-
Fig. 5.20. Environmental setting of depositional carbonate breccias. A: Cross section from land – to platform and platform margin with reefs – to slope and deep swells – down to deep-water basin. B: Cross section across a distally steepened ramp.
Example of a Polymict Breccia
241
Fig. 5.21. Polymict depositional carbonate breccia. This mass-flow breccia is part of the Early Permian ‘Tarvis breccia’ in the Southern Alps (see Pl. 27/1). The microfacies of the angular and subangular limestone clasts testifies to the erosion and destruction of different parts of a carbonate platform and platform margins. Textural criteria of clasts: Shape angular and variously rounded, high roundness number, low fitting, clast/matrix boundary sharp, ratio clasts : matrix 3 : 1. Color differences in the clasts correspond partly to differences in microfacies. Microfacies of clasts: The numbers refer to microfacies types: The environmental interpretation is based on reference sections exhibiting the spectrum and the spatial distribution of microfacies types of shelf and shelf margin platform carbonates (Flügel 1980) and on environment-diagnostic associations of calcareous algae (Flügel and Flügel-Kahler 1980). 1 - bioclastic grainstone with coated skeletal grains, worn fusulinid shells and rounded intraclasts (high-energy shelf lagoon), 2 - bioclastic packstone with dasyclad green algae and smaller foraminifera (protected lagoon), 3 - bioclastic poorly sorted packstone with various algae, smaller foraminifera and echinoderm fragments (open-marine shelf lagoon), 4 - bioclastic grainstone with gastropods and dasyclad algae (protected shelf lagoon), 5 - poorly sorted oolitic grainstone with tangential small ooids, various skeletal grains and cortoids (mixed material derived from ooid shoals and open shelf lagoonal environments), 6 - lithoclastic packstone with various worn skeletal grains and rounded intraclasts (high-energy shelf lagoon), 7 oncoid floatstone with large cyanobacterial oncoids (protected subtidal parts of a shelf lagoon), 8 - boundstone with Tubiphytes, encrusting sponges and Archaeolithoporella, see Pl. 98/8 and Pl. 145 (shelf margin reefs, these boundstones differ in rock colors and the proportions of encrusting organisms and grains), 9 - small clast of black lime mudstone (probably of lacustrine origin, compare Pl. 27/1), 10 - bioclastic wackestone with sessile smaller foraminifera (mud mound), 11 - finegrained bioclastic grainstone with abundant green algae (protected lagoon), 12 - coated bioclastic grainstone with algae, foraminifera and other skeletal grains (open-marine shelf lagoon), 13 - fine-grained bioclastic grainstone with skeletal grains and peloids (protected lagoon), 14 - laminated bindstone (tidal flat). Scale is 1 cm. The uniform red matrix (M) consists of poorly sorted lithoclastic wackestones without fossils. The angular lithoclasts represent erosional debris infilled together with carbonate mud into interclast voids. Solution voids at the margin of some clasts and fractures are filled with scalenohedral calcite crystals (CC), subsequent to compaction and pressure solution. White arrows show penetration of clasts, black arrows indicate surficial solution of some clasts.
242
ate conglomerates include (a) storm or swell lags in shore-face environment (Einsele and Seilacher 1982; Nemec and Steel 1984; Aigner 1985; Craft and Bridge 1987), (b) lateral accretion lags at the base of tidal channels (Shinn 1986; Waters et al. 1989), (c) transgressive lags, originating from a rip-up of underlying older tidal flat sediments during rises in sea level (Pl. 9/7).
5.3.3.5 Significance of Carbonate Breccias and Conglomerates Fig. 5.20 shows the setting of breccias and conglomerates. Box 5.9 summarizes the significance of breccias and conglomerates for facies analyses and economic issues. Basics: Carbonate breccias and conglomerates Bathurst, R.G.C. (1959): Diagenesis in Mississippian calcilutites and pseudobreccias. – J. Sed. Petrol., 29, 365-376 Blount, D.N., Moore, C.H. (1969): Depositional and nondepositional carbonate breccias, Chiantla Quadrangle, Guatemala. – Geol. Soc. Amer. Bull., 80, 429-442 Böhm, F., Dommergues, J.L., Meister, C. (1995): Breccias of the Adnet Formation: indicators of a Mid-Liassic tectonic event in the Northern Calcareous Alps (Salzburg/ Austria). – Geol. Rundschau, 84, 272-286 Flügel, E., Kraus, S. (1988): The Lower Permian Sexten breccia (Sexten Dolomites) and the Tarvis breccia (Carnic Alps): Microfacies, depositional environment and paleotectonic implications. – Mem. Soc. Geol. Ital., 34, 67-90 Füchtbauer, H., Richter, D.K. (1983): Carbonate internal breccias: A source of mass flow at early geosynclinal platform margins in Greece. – Soc. Econ. Paleont. Min. Spec. Publ., 33, 207-215 Füchtbauer, H., Richter, D.K. (1983): Relations between submarine fissures, internal breccias and mass flow during Triassic and earlier rifting periods. – Geol. Rundschau, 72, 53-66 Koster, E.H., Steel, R.L. (1984): Sedimentology of gravels and conglomerates. – Canadian Soc. Petrol. Geol. Mem., 10, 441 pp. Laznicka, P. (1988): Breccias and coarse fragmentites. – Dev. Econ. Geol., 25, 832 pp., Amsterdam (Elsevier) Norton, W.H. (1917): A classification of breccias. – J. Geol., 25, 160-194 Richter, D.K., Füchtbauer, H. (1981): Merkmale und Genese von Breccien und ihre Bedeutung im Mesozoikum von Hydra (Griechenland). – Z. deutsch. geol. Ges., 132, 451501 Schaad, W. (1995): Die Entstehung von Rauhwacken durch die Verkarstung von Gips. – Eclogae geol. Helvet., 88, 59-90
Significance of Breccias and Conglomerates
Box 5.9. Significance of carbonate breccias and conglomerates. Paleoenvironmental proxies Since the formation of many breccia types is restricted to specific settings (Fig. 5.20), depositional and nondepositional breccias assist in the paleoenvironmental interpretation of non-marine and marine carbonates. Caliche breccias and karst breccias indicate subaerial exposure, solution-evaporite-collapse breccias may point to sea-level lowering (Obrador et al. 1992). Depositional patterns Breccias can mark underlying sequence boundaries (e.g. Biddle 1984, Garcia-Mondejar 1990) and offer information on the geometry of carbonate ramps, e.g. distal steepening. Reconstruction of provenance areas Clast analyses of breccias and microfacies studies of conglomerates provide excellent data which allow detailed reconstructions of the changing sedimentary input into marine or non-marine basins, the provenance of the material and the climatic and tectonic controls of the processes involved (see Pl. 27, Fig. 5.21 and 16.3). Tectonic events Breccias can indicate the proximity of faults which may otherwise not be directly detected (e.g. Bosellini 1989; Garcia-Mondejar 1990). Economic importance Reservoir rocks: Fractured carbonate breccias are important hydrocarbon reservoirs. Of particular interest are solution breccias (Loucks and Anderson 1985; Miller 1985), debris-flow breccias (Enos 1988) and karst breccias. Mineralizations: Breccia-hosted karst and collapse breccias host Mississippi Valley-type lead-zinc deposits (Klau and Mostler 1983; Bechstädt et al. 1994; Liaghat et al. 2000). Building stones: Many polymict carbonate conglomerates and breccias are used as building and decoration stones (Pl. 26/3), e.g. ‘Nagelfluh’ (an interglacial conglomerate), or the socalled ‘Untersberg marble’ (a Cretaceous mass-flow breccia) in Austria.
Stanton, R.J. (1966): The solution brecciation process. – Geol. Soc. Amer. Bull., 77, 843-848 Wise, D.U., Dunn, D.E., Engelder, J.T., Geiser, P.A., Hatcher, R.D., Kish, S.A., Odom, A.L., Schamel, S. (1984): Fault related rocks: suggestions for terminology. – Geology 12, 391-394 Further reading: K072
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6 Quantitative Microfacies Analysis
The previous chapters of this book introduced various aspects of recognizing qualitative microfacies criteria and their signifcance, whereas this chapter will treat quantitative criteria. The first part describes selected principles of grain-size analysis and demonstrates the potential of grain size in the context of microfacies studies. The second part deals with methods of frequency analysis and their use in microfacies-based interpretations. The last part discusses the evaluation of microfacies by multivariate statistical analyses.
tending from <1/256 mm to >256 mm. A modification of this scale is the φ-scale which allows grain-size data to be expressed in units of equal value. The φ size is the negative logarithm to the base 2 of the grain size measured in millimeters (Krumbein 1936). A grain size of 1 mm corresponds to a φ value of zero. Smaller grain sizes are expressed by positive phi values, larger grain sizes by negative φ values. Pierce and Graus (1981) and McManus (1982) have discussed the use and misuse of the φ-scale. Phi values can be obtained from conversion tables (Page 1955; Griffith 1967) or nomograms (Friedman and Sanders 1978).
6.1 Grain-Size Analysis Grain size is a basic descriptive measure of sediments. Grain-size distribution patterns may be characteristic of sediments deposited in certain environments and can yield information about depositional processes. In turn grain sizes are a major control on porosity and permeability. Various methods are in use ranging from simplistic diagrams of grain-size parameters up to sophisticated analyses of statistically treated data. Measures for characterizing grain-size populations have been reviewed at length in many standard texts (e.g. Boggs 1995) and include a wide spectrum of analytical methods (McManus 1988; Syvitski 1991; Pye 1994). Most descriptive and interpretative concepts are based on studies of modern quartz sands deposited in specific environmental settings. Grain-size criteria of modern carbonate sediments have been examined predominantly for tropical and subtropical inter- and subtidal sands accumulating near beaches, lagoons, on shallowmarine banks and within reef complexes (Box 6.1). Contrary to the abundant grain-size analyses of sandstones, only a few authors have studied grain-size distributions of ancient carbonate rocks. Grain-size scales: The commonly adopted size scales are shown in Fig. 6.1. Sedimentary grains range in size from a few microns to a few meters. Because of this wide range, logarithmic or geometric scales rather than linear scales are more useful for expressing size. The Udden-Wentworth scale is a geometric scale ex-
Grain-size groupings in carbonate rocks: The terms calcilutite, calcarenite and calcirudite, proposed by Grabau (1904), refer to the sizes of carbonate particles. Calcilutite consists of clay- and silt-sized particles (<2 µm to 62 µm); it corresponds to a consolidated carbonate mud. Calcisiltite (Kay 1951) embraces the coarser part of calcilutite (Sect. 4.1.4). Calcarenite is a limestone consisting of sand-sized grains (0.062 to 2 mm); it corresponds to a consolidated calcareous sand. Calcirudite is a limestone consisting of grains which are larger than 2 mm. The term embraces consolidated calcareous gravels and rubbles as well as limestone conglomerates and breccias. Some authors use a lower limit (1 mm). The terms calcarenite and calcirudite are of common use in microfacies studies. Thin sections exhibiting characteristic medium- and coarse-grained calcarenites are shown in Pl. 43/1, Pl. 44/2, and Pl. 45/1. Typical calciruditic limestone types are depicted in Pl. 45/2 and Pl. 45/3. A comparison of these microphotographs shows considerable differences with regard to the number of the sand- or pebble-sized grains, the character of the matrix and the mud- or grain-supported fabric. These differences have been taken into account by a classification used by Carozzi (1988) and his students. In the original definition the upper grain size of calcirudite is not limited and embraces rocks consisting of millimeter- to meter-sized clasts. It seems, however,
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_6, © Springer-Verlag Berlin Heidelberg 2010
Grain Size Analysis
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This chapter focuses primarily on some of the methods used in grain-size analysis and discusses possibilities and problems related to evaluating of grain-size patterns in the context of interpreting modern and ancient sedimentary environments.
Boulders
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6.1.1.1 Measuring Grain Sizes and Describing Size Distributions
very coarse coarse medium
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6.1.1 Grain-Size Analysis: Methods and Aims
Clay
Micrite
very fine
Fig. 6.1. Wentworth-Udden grain-size scale and grain-sizebased terms used for differentiating of carbonate rocks. Note that the upper limit of micrite is often expanded and the size range of calcisiltite is a matter of discussion (Sect. 4.1.1 and 4.1.4).
advisable to restrict the term to the millimeter- to centimeter-sized range, and to call limestones consisting of larger and large clasts breccia or conglomerate. The term calcirudite and the term rudstone used in limestone classifications are not identical (Sect. 8.3.2). Both terms refer to carbonate rocks containing particles > 2 mm, but calcirudite considers only grain size whereas rudstone is defined by grain size and the number of grains larger than 2 mm. Note that calciruditic limestones embrace rudstones as well as floatstones, and that the term calcarenite covers grain sizes represented in grainstones and wackestones of the Dunham classification.
The size of siliciclastic grains is determined by several techniques. Unconsolidated sediments and disintegrated rocks of clay-, silt-, and sand-size ranges are commonly measured by sieving or sedimentation methods (Coakley and Syvitski 1991; Matthews 1991). Modern carbonate sands have usually been measured by sieving. Lithified sandstones and limestones require thin-section measurements for silt- and sand-sized rocks and electron microscopy for clay-sized samples. Image analysis is a highly promising method for rapid and accurate measuring of grain sizes, both of loose grains and indurated sediments (Schäfer 1982; Kennedy and Mazzullo 1991). The interpretation of grain-size patterns requires graphical and mathematical treatment of the data. Common graphic methods for presenting grain-size data are histograms, frequency curves and cumulative curves. Histograms are bar diagrams in which grain size is plotted along the abscissa of the graph and individual weight percent along the ordinate. Frequency curves result from connecting the midpoints of each size class in the histogram with a smooth curve. Cumulative curves are grain-size curves generated by plotting grain size against cumulative weight percent frequency. The curves can be plotted on an arithmetic scale or on a log probability scale in which the arithmetic ordinate is replaced by a log probability ordinate. The latter method allows the important grain-size statistical parameters to be calculated using the φ values read from the cumulative curve for various percentiles (Tab. 6.1). Breaks and discontinuities in cumulative curves produced by plotting of grain sizes on a log probability scale are interpreted as results of mixing of sediment populations related to different transport processes (Visher 1969). Grain-size data can be plotted statistically, deriving parameters that allow average size and sorting of sedimentary populations to be expressed mathematically:
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Sorting
Fig. 6.2. Chart for visually estimating of sorting (based on Pettijohn et al. 1973). The chart was developed for sandstones and differs from the charts in Fig. 6.3 by the overrepresentation of larger grains. After Tucker (1981).
Average grain size is the easiest parameter to determine but is sometimes a rather inaccurate value. The average grain size depends on the grain sizes of the material available and the amount of energy in the transport medium. Average grain sizes are measured by mode, median size or mean size distribution. The mode or modal diameter (Mo) describes the most frequently occurring particle size in a population of grains. The modal size diameter corresponds to the diameter represented by the steepest point (inflection point) on the cumulative curve or the highest point on the frequency curve. The median size (Md) is expressed by the midpoint of the grain size distribution and corresponds to the 50% percentile diameter on the cumulative curve. The mean size (Mz) is determined by the arithmetic average of all the particle sizes in a sample. Sorting is a measure of the grain sizes present in a grain population and describes the magnitude of the spread or scatter of these sizes around the mean size. Sorting of water-laid sediments is produced by currents, waves and splashing. Sorting can be described with the help of visual comparison charts (Figs. 6.2 and 6.3) or, more precisely, by parameters which document the number and the percentage of grain-size fractions within a grain mixture. The φ sorting standard deviation and skewness (Sk) values are determined graphically or by calculation (Tab. 6.1). φ skewness describes the deviation from a normal or lognormal grain size distribution and the degree of asymmetry. Skewness reflects sorting in the tails of a grain-size population. Positively skewed populations have a tail of excess fine particles, negatively skewed populations a tail of excess coarse particles. Terms for sorting that correspond to various values of standard deviations and skewness have been proposed by Folk and Ward (1957) and Friedman (1962); Fig. 6.5. Skewness is given an environment sensitive significance by many authors (e.g. Friedman 1967; Folk and Robles 1964; Valia and Cameron 1977).
Beach sands, for example, tend to have negative skewness, since fine particles have been removed by the persistent wave action. River sands are often positively skewed, since not much silt and clay is removed by the currents. The degree of the peakedness or sharpness of grainsize curves is expressed by kurtosis values, which are commonly calculated along with the other grain-size parameters. Commonly used formulas for calculating grain-size parameters by graphic methods are shown in Tab. 6.1. Grain-size statistical parameters can also be calculated without reference to graphic plots by the mathematical method of moments. The parameters obtained by this method are analogous to those of graphical statistics, but because moment measures employ the entire frequency distribution rather than a few selected percentiles, the parameters are considered to be more representative for the sample than the graphically derived values. For overviews on the various grain-size parameters see Folk (1966) and Syvitski (1991). Table 6.1. Formulas for calculating statistical parameters for grain sizes by graphic methods (Folk and Ward 1957). The Greek letter φ (phi) for grain sizes indicates that the percentiles were measured using the phi and not the mm-scale.
Graphic mean
φ16 + φ50 + φ84
Mz =
3
Inclusive graphic standard deviation φ84 – φ16 σi = + 4
φ95 – φ5 6.6
Inclusive graphic skewness SKi =
(φ84 + φ16 – 2 φ50) 2 (φ84 – φ16)
Graphic kurtosis
+
KG =
(φ95 + φ5 – 2 φ50) 2 (φ95 – φ5)
(φ95 – φ5) 2.44 (φ75 – φ25)
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Interpretation of Grain-Size Data
Fig. 6.3. Comparison charts for particle sorting after Longiaru (1987). Labels indicate the degree of sorting (standard deviation) approximated by the comparator and the descriptive terminology of Folk (1966). Visual estimates of sorting can be obtained by comparison with charts, some of which have been developed particularly for calcarenites (e.g. Harrell 1984). Note that visual comparison charts assume a lognormal distribution of grains, which is not always the case, and that in most charts larger grains are overrepresented except in the charts shown in this figure. The bar at the lower right of each diagram indicates the maximum grain size which might occur in the samples.
6.1.1.2 Approaches to the Environmental Interpretation of Grain-Size Data Hydrodynamically controlled covariations of two or more measures, such as mean size and sorting are commonly available. Following early studies by Friedman (1961, 1967) and Douglas (1968) the use of bivariate scatter diagrams (e.g. skewness vs. standard deviation or mean size vs. standard deviation) has become a common approach in attempts to discriminate modern river,
beach and dune sands (Friedman 1979; Eschner and Kirchner 1984; El-Ella and Coleman 1985; Sutherland and Lee 1994). The Friedman diagrams allow the plots to be differentiated into major environmental fields. Passega (1957, 1964. 1977) suggested that the ratio of the coarsest one percentile, C, to the median diameter, M, and the percentage of grains finer than 31 µm (L) are indicators of the dynamics in different depositional environments. The strongest current should define the largest stable particle size. Different fields in
Grain Size in Thin Sections
the C-M diagram typify deposits from traction currents, from beaches or from various forms of suspension (Rizzini 1968; Fabri et al. 1986). Different environments can be distinguished by examining the patterns produced by plotting a number of samples from the same unit on C-M or L-M diagrams (Fig. 6.6). The Passega diagram allows the type of sediment transport to be differentiated preferentially for sediments with C < 1 mm (Gromoll 1992). The diagram has been successfully used both in the interpretation of modern sediments and ancient sedimentary rocks (Faugères 1974). Other authors have suggested that the depositional environments and different transport dynamics can be interpreted on the basis of the shapes of grain-size cumulative curves plotted on probability paper (Visher 1969; Glaister and Nelson 1974). Several multivariate statistical techniques (factor analysis, discriminant function analysis, log-hyperbolic distribution, entropy analysis) have been successfully applied to the interpretation of grain-size distribution (Klovan 1967; Vincent 1986; Forrest and Clark 1989; Merta 1991). The value of nearly all these techniques for identifying depositional environments has been questioned in parts or as a whole (e.g. Vandenberghe 1975; Tucker and Vacher 1980; Sengupta et al. 1991). The main reasons for these difficulties are related to the variability of depositional conditions within the major environmental settings and the complexity of depositional processes: • Hydrodynamic conditions within the sedimentation area can considerably vary, depending on the size of the sedimentary basin. • Spatial and temporal grain size distribution within a given physiographic setting are highly variable, and demand critical statistical analysis (Piller and Mansour 1990, Medina et al. 1994). • Frequency and cumulative curves reflect grain-size subpopulations, distinguished by rolling, saltation or suspension transport (Smolka 1990). • Primary grain-size distributions can be changed by bioturbation and episodic high-energy events, e.g. strong storms, as well as by variations in sediment influx caused by sea-level fluctuations. In summary, grain-size distributions primarily reflect transport and depositional processes which are not necessarily unique to a particular environment. Nevertheless, in the context of facies analyses, grain-size data assist in environmental interpretation if the data are incorporated in the framework of textural and compositional criteria and all lines of evidence are regarded.
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However, contrary to the information provided by ecology or sedimentary structures, the grain-size evidence provided has the advantage of being universal.
6.1.1.3 Grain-Size Analysis in Thin Sections In sandstones and limestones grain sizes are usually measured in thin sections, peels or microphotographs. Rocks should be sectioned in a consistent direction. Sections parallel to the bedding plane would be more appropriate but most students use vertical or non-oriented thin sections. Peels can be made of much larger surfaces than conventional thin sections, allowing a large amount of grains to be measured (Gutteridge 1985). Grain size analysis of thin sections requires statistical grain selection according to a well-defined sampling plan. Techniques used include point-, line-, area-, and ribbon-counting (Sect. 6.2.1). Point counting (grains selected at grid points) appears to be the most appropriate method for sandstones, ribbon-counting as well as point-counting are commonly used for limestones. The number of grains measured ranges between 150 and 500. Measurement of 300 grains is often regarded as statistically sufficent for sandstones. Calcarenitic and calciruditic carbonates must be handled differently (Sect. 6.2). Measurement errors: Grain sizes derived from thin sections are apparent, not true grain sizes. The probability that a grain will be cut by a random section is proportional to the size of the grain. Diameters measured in thin sections, therefore are usually smaller than the true grain diameter. Random sections lead to an increase of measurements of small grains. Conversion of grain-size data: Basic concepts of grain-size analysis were developed by studying loose sediments or disintegrated sedimentary rocks. Because the commonly used sieve method can not be applied to indurated rocks, a conversion of granulometric data obtained by thin section analysis into sieve data appears necessary if statistical parameters are also to be evaluated for sandstones and carbonate rocks. This conversion has been based on different regression models (Friedman 1958, 1962; Harrell and Erikson 1979; Merta 1991). Müller (1964) developed a nomogram from which the sieve-size distributions can be directly derived. A comparison of sieve and thin-section methods used for siliciclastic rocks indicates that the two techniques must be regarded as mathematically independent methods (Burger and Skala 1979).
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6.1.2 Grain-Size Studies of Modern and Ancient Carbonates The following chapter summarizes the still rather incomplete information about grain-size distributions of recent marine carbonate sands, and discusses the application of grain-size studies to the microfacies analysis of carbonate rocks.
6.1.2.1 Grain-Size Studies of Modern Carbonate Sediments How does sorting of carbonate sediments compare with sorting of siliciclastic sediments? Differences between siliciclastic and carbonate sands: The grain-size distribution of siliciclastic sands depends on the source, weathering, erosion and transport processes. Marine carbonates are predominantly biogenic sediments (Sect. 2.1.2). Therefore, the size of carbonate particles is highly controlled by the original size and decay of the skeletons, and by size-reduction processes which take place during transport. In subtidal carbonate environments, however, the areal effect of larger transport processes is rather restricted and the sediment is produced nearby or at the site of final deposition. Larger transport processes occur if storms and strong currents cause bottom erosion and redeposition of carbonate sands. Many bioclastic carbonates exhibit a bimodal grain-size distribution consisting of silt- and sand-sized grains and mud. Common size range of carbonate grains: The sizes of carbonate grains vary between <0.10 mm and several centimeters. Skeletal grains, oncoids, rhodoids, aggregate grains and extraclasts exhibit wide size ranges, peloids, ooids, pisoids and intraclasts rather narrow ranges: • Skeletal grains: Commonly within the range of < 0.250 mm to about 20 mm; coarse calcisiltite, calcarenite and calcirudite; medium silt, sand, granules and to coarse pebble size. • Peloids: Commonly about 0.10 to 0.20 mm (except for fecal pellets which can be larger); fine to medium-sized calcarenite; fine and medium sand size. • Ooids: < 2 mm, commonly 0.5 to 1 mm; coarse and very coarse calcarenite. • Pisoids: > 2 mm to several tens of millimeters, but sometimes < 2 mm; calcirudite; granules to coarse pebble size.
Controls on Grains Size
• Oncoids: < 2 mm to > 40 millimeter, commonly < 20 mm; calcirudite; granules to very coarse pebble size. • Rhodoids: about 5 to 40 mm, commonly < 100 mm; calcirudite; fine to very coarse pebble size. • Aggregate grains: 0.5 to several millimeters, commonly < 1 mm; calcarenite and calcirudite; coarse sand to medium pebble size. • Intraclasts: < 0.10 to about 10 mm; calcarenite and calcilutite; fine sand to coarse pebble size. • Carbonate extraclasts: < 1 mm to several tens of millimeters; calcarenite and calcirudite. Controls on the sizes of carbonate grains If the grain size of carbonate rock is taken as a guide to energy levels in the environment of deposition, several points must be born in mind: • size and sorting of skeletal grains may reflect only growth sizes as for example, shown by foraminiferal sands in back-reef settings; • different sizes of microfossils can by caused by hydraulic sorting as described from calciturbidites (Herbig and Mamet 1994); • calcareous organisms have different architecture and will therefore disintegrate in different ways; corals and Halimeda, for example, have an ‘inbuilt’ bimodality (Folk and Robles 1964; Hoskin et al. 1983; Fig. 6.4); • skeletal grains have unique hydrodynamic properties and are commonly porous or hollow and contain organic matter (Braithwaite 1973); • the effect of the hydraulic modification on the depositional grain size of skeletal grains depends on the shape, bulk density, size, angularity and structure of bioclastic grains. The variability in the shapes of skeletal grains causes variations in the settling velocities of particles with identical diameters; • skeletal grains, ooids, peloids and intraclasts can differ in their susceptibility to transport depending on the differences in grain shape and density of the grains (Matthews 1991); • sand-sized skeletal grains and intraclasts often exhibit similar grain-size and sorting patterns; • carbonate fecal pellets differ in their hydrodynamic modification from other carbonate grains (Wanless et al. 1981); • ooids are susceptible to long-distance transport by currents: sorting of ooids is a measure for ooids deposited more or less in place or allochthonous transported ooids (Anwar et al. 1984; Sect. 4.2.5); • specific grain sizes can be produced by organisms (fishes, crabs) that crush hard parts in their search for food;
Modern Grain-Size Patterns
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Fig. 6.4. Disintegration of different calcareous organisms results in different grain size fractions. In this example, the degree of disintegration depends on the growth form and microstructure of the skeletons. Other case studies indicate significant controls by hydrodynamic conditions. Calcareous green algae (Halimeda) disintegrate into millimeter- to micron-sized particles (see Fig. 4.3), corals (Acropora) into fine sand- and gravel-sized particles. In high-energy environments, the branching colonies of the corals break into well-sorted sticks (average diameter about 4.5 mm); on abrasion, the sticks break further down into well-sorted grains forming coral debris (0.25 mm). There is little coral material of intermediate size. Halimeda first falls apart into thin leaflike segments and further breaks into pieces averaging about 1 mm. During abrasion segments eventually disintegrate into micron-sized aragonite crystals contributing to the formation of carbonate mud (< 0.004 mm). Note that the example does not demonstrate the quantitative importance of ‘algal micrite’ in low-energy lagoonal sediments.The graphic is based on sieve analyses of high energy beach sediments of Isla Perez, Alacran reef complex (Yucatan). Modified from Folk and Robles (1964). • major storm events, e.g. hurricanes and cyclones, can completely change long-term grain-size distribution patterns in reefs and on platforms within hours.
Grain-size patterns of modern carbonates Bioclastic carbonate sands: Most grain-size studies of recent carbonates concern sand-sized sediments deposited near the beach, in shallow inter- and subtidal environments, and in reef complexes (Box 6.1). Studies of modern depositional environments reveal some general patterns with respect to the size distribution of carbonate grains in coastal and shallow-marine environments and allow their major controls to be discussed. These patterns were first recognized by Ginsburg (1956) in Florida and later confirmed by statistically based analyses of Caribbean reefs and platforms (Alacran reef, Mexico: Folk 1962, 1967; Folk and Robles 1964; Belize: Gischler 1994; St.Croix: Hubbard et al. 1981), the Australian Great Barrier Reef (Flood and Orme 1977), land-locked environments (Piller and Mansour 1990) and reefs in the Red Sea, and in the near-coast environments of the Persian Gulf and the Gulf of Oman (Sarnthein 1971). Study areas have different scales: small – up to about one kilometer: Alacran reef and St.Croix; medium – several kilometers: Belize and Red Sea; large – more than hundreds of kilometers: Great Barrier Reef, Persian Gulf and include
different depositional settings (open-marine platform edges, open and protected lagoons, platform reefs, topographically differentiated shallow bays, subaerial and submerged coastal areas). The gross composition of the sediments is similar and is characterized by the dominance of bioclastic sands composed predominantly of variously disintegrated calcareous algae, corals, mollusks and foraminifera. In addition, fecal pellets and intraclasts are of some importance. Occurrence of coarse-, medium- and fine-grained calcareous sands in shallow-marine environments (1) Coarse- and very coarse-grained calcareous sands (0.5 to 2 mm): Coarse bioclastic sediments composed predominantly of calcareous algae, corals or foraminifera occur in different shallow-water settings including outer windward reef rims and forereef areas, reef flats, situated near patch reefs in inner platform and lagoonal settings, at the boundaries of inner platform basins and beaches. Some of these sands are lag deposits (reef rims). Many but not all coarse-sized sands occur in high-energy environments. (2) Medium to fine-grained sands (0.25 to 0.12 mm): Common in areas affected by tidal currents or in subtidal areas of changing hydrodynamic conditions.
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(3) Fine- and very fine-grained sands (< 0.25 mm): Common in central parts of lagoons, in innerplatform basins, in deeper depressions of topographically differentiated subtidal areas, in small or large protected areas and in areas covered by seagrass or mangroves. Vegetation favors the accumulation of fine-grained well-sorted bioclastic sediment; the sediment consists of small calcareous organisms living on the plants and allochthonous grains trapped within the plants. Sorting Specific sorting values of shallow-marine calcareous sands measured by inclusive standard deviation and skewness are not limited to particular depositional environments but are important in combination with other grain-size parameters. Well to moderately well sorted, coarse calcareous sands are common on beaches and at reef rims. Moderately sorted coarse sands occur on beaches, submarine ridges, near inner platform and lagoonal patch reefs, on inner carbonate ramps and open-marine platforms. Moderately to poorly sorted sands are widely distributed. Coarse sands occur in tidal and shallow subtidal environments. Fine and very fine moderately to poorly sorted fine sands are known from tidal environments, the leeward part and the central parts of reefs and in lagoons of platform reefs. Poorly to very poorly sorted fine-grained sands occur in protected basins and topographic depressions. Summary Grain-size distributions reflect the availability of skeletal grains, i.e. the occurrence of certain organisms in specific subenvironments, the degree of water energy controlled by tides, waves and storms, and the differentiation of hydrodynamic conditions by submarine bottom topography. Grain size reflects only some of the variability of the sedimentary bottom facies. Well sorted to moderately sorted sands appear to be commonly controlled by high-energy conditions, brought about by wave action and/or currents. Poor sorting may be caused by tidal and bottom currents or reflect just the deposition of larger eroded grains (e.g. coral detritus) within finer sediment. In larger reef complexes, the grain-size distribution of the bioclastic sediments is indicative of hydrodynamic processes acting within defined depositional environments. Large-scale grain-size studies based on multivariate analyses (e.g. Flood and Orme 1977) support the classification of carbonate rocks according to textural criteria (grain size, sorting, etc.) which are in-
Summary: Grain Size
terpreted as indicators of different water energy (Plumley et al. 1962; Sect. 12.1.1.2). The combination of grain-size parameters (mean, skewness, kurtosis) separates predominantly those environments which differ strongly in their hydrodynamic controls. These environments are not restricted to specific depositional settings. Box 6.1. Selected grain-size studies of modern and ancient carbonates. * Studies of calcareous sands in reef complexes. Modern environments Shallow-marine banks, platforms and reefs: Anwar et al. 1984; Basan 1972*; Davies and Conley 1977; Early and Goodell 1968; Emery 1963; Flood and Orme 1977*; Folk 1962, 1967; Folk and Robles 1964*; Ginsburg 1956*; Gischler 1994*; Hoskin and Nelson 1971; Hubbard et al. 1981; Jindrich 1969; Kench and McLean 1997; Khalaf et al. 1984; Koopmann et al. 1979; Land 1979*; Lewis 1969*; Lynts 1966; Mallik 1976*, 1979*; Maxwell and Flemming 1961*; Maxwell et al. 1961*; McKee 1959*; Milliman 1973; Montaggioni et al. 1986*; Müller 1964; Müller and Müller 1967; Passega 1964; Pilkey et al. 1967; Piller and Mansour 1990*; Purdy 1963; Rao and Amini 1995; Rizzini 1968; Sarnthein 1971; Scoffin et al. 1980*; Upchurch 1972 Size and hydrodynamics of carbonate grains: Braithwaite 1973; Chave 1960; Chardy et al. 1988; Driscoll and Wettin 1973; Force 1969; Hoskin et al. 1983; Maiklem 1968; Swinchatt 1965; Wanless et al. 1981 Ancient environments Shallow-marine platforms and reefs: Boccaletti et al. 1969 (Jurassic); Carozzi 1958, 1961* (Devonian); Carozzi and Diaby 1981 (Early Carboniferous); Carozzi and Hulse 1963* (Devonian); Carozzi and Lundwall 1959* (Devonian); Carozzi and Reichelderfer 1987 (Early Carboniferous); Carozzi and Zadnik 1959* (Silurian); Carss and Carozzi 1965* (Devonian); Dawson and Carozzi 1986* (Late Carboniferous); Fabricius 1967 (Late Triassic and Early Jurassic); Flügel and FlügelKahler 1963* (Late Triassic); Faugères 1974 (Jurassic); Gindy and Sultan 1978 (Miocene); Hötzl 1966 (Late Jurassic); Jaanusson 1952 (Ordovician); Merta 1991 (Jurassic); Ott 1969 (Late Jurassic); Rad 1974 (Tertiary and Quaternary); Rao and Carozzi 1973 (Mississippian); Rogen et al. 2001 (Cretaceous chalk); Usdowski 1962 (Triassic); Zacher 1973 (Cretaceous) Turbidites: Babek and Kalvoda 2001 (Late Devonian and Early Carboniferous); Carozzi and Baneree 1984 (Devonian); Carozzi and Frost 1966 (Silurian); Eder 1971; Engel 1974 (Early Tertiary); H.W. Flügel and Pölsler 1965 (Late Jurassic); Herbig and Mamet 1994 (Early Carboniferous); Meischner 1964 (Devonian); Steiger 1981 (Late Jurassic)
Guide to Grain-Size Studies
6.1.2.2 Application of Grain-Size Analyses to Carbonate Rocks Box 6.2 summarizes important points which should be considered when performing grain-size analyses of microfacies data. Not mentioned in the Box is the index of clasticity introduced by Carozzi (1958). This index is defined as the mean apparent diameter (expressed in millimeters) of the 10 largest grains present in the thin section. The index can be only used for grains which have obviously undergone transport. Grains used
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are bioclasts, fossils, detrital minerals (e.g. terrigenous quartz) and intrabasinal particles (intraclasts, peloids, ooids, oncoids). Because specific densities of bioclasts vary as a function of intraparticle porosity and because the transport potential of skeletal grains depends on shape and size, not all bioclast types can be used equally in establishing the clasticity index. Carozzi recommends echinoderm fragments, in particular crinoid columnals. In the case of bimodal sediments, both modes are measured. The index is regarded as a relative measure of the energy level in the depositional environment. It is
Box 6.2. Guide to grain-size studies of carbonate rocks. Field work Sampling: Samples should be large enough to show vertical and lateral textural variability (Davies and Conley 1977). Take several samples if variability appears great. Choose beds showing grain-supported fabrics and grains within sand to pebble-size range. The evaluation of grain size using Passega diagrams requires at least 20-30 samples taken from a depositional sequence and representing all textures present. Laboratory work Thin sections and peels: Investigate large thin sections, with a size of at least 5 x 5 cm. Peels are useful if the outlines of the grains are clearly visible. What to measure? In general, all grains should be measured, but skeletal grains, cortoids and synsedimentary reworked particles (intraclasts) are more useful than peloids or large oncoids. How to measure? Follow a strict sampling plan and use the ribbon counting method or the point counting method (see Figs. 6.7 and 6.8). Point-counting data are biased toward coarser grain sizes as compared with data resulting from ribbon-counting measurements. Only the latter method also considers small and inconspicuous grain sizes. Measure the largest apparent diameter of the grains. To reach at least some statistical basis measure more than 100 but, better, up to 300 or more grains. Conversion of mm to φ values: Conversion is necessary to produce grain-size diagrams. Use conversion tables or nomograms. Grain-size distributions: Describe grain-size distributions by cumulative curves (plotted on probability paper) and try to separate these curves into linear segments (following the method of Visher 1969). The Visher method aims to separate subpopulations, which should reflect the proportions of grains transported by saltation, traction and suspension processes. Grain-size parameters: Calculate average size and sorting using the formulas presented in Table 6.1. Inclusive standard deviation and inclusive skewness are of particular importance. There are several computer programs available which calculate the relevant parameters. Diagrams: Draw Friedman-diagrams showing possible relations between sorting and average grain size. Draw Passega diagrams showing the relations between the coarsest grains and median diameter. An advantage of the Passega diagram is that both necessary parameters can be determined from thin sections of limestones. Cluster analysis: In order to obtain more objective results than by comparing only combinations of grain-size parameters, all data should be treated by statistical analyses showing similarities and allowing groups to be recognized which may reflect controlling depositional processes. Interpretation Be cautious: Consider the possibility that grain-size ranges of skeletal grains are predominantly dependent on individual growth rates and microstructures of organisms, or are caused by bioeroders which produce specific grain sizes. If you use grain-size parameters, take care, because the concepts for discriminating depositional environments or processes by grain sizes have been developed for siliciclastic sediments, and are based on sieve and settling methods. Some authors emphasize the necessity of transforming thin-section data (number percent) into sieve data (weight percent) prior to drawing and evaluating grain-size diagrams. Comparisons of thin-section data and transformed data of calcarenites show a shift of data (higher percentage of coarser grains in thin sections), but the general pattern shown in the diagram is not essentially changed (Fig. 6.6). Additional indications: If your mind is set on a very specific answer (‘look, that is a beach sediment’, or ‘seems to point to high-energy conditions although other criteria do not show this’), be critical and look for further indications, e.g. the hydrodynamic conditions revealed by the composition of the sediment, and by paleontological and paleoecological criteria (Sect. 12.1).
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Grain Size: Case Studies
often used to describe textural changes over time in the controls of microfacies types (Carozzi 1988). Up to now few but thought-provoking grain-size analyses of ancient carbonate rocks exist. The following case studies demonstrate the potential of this approach. Turbidites are discussed in Sect.15.7.3.4
Case studies: Grain-size analyses of Late Jurassic platform and slope carbonates from the Alps Late Jurassic carbonates in the central and western Northern Calcareous Alps (Austria) offer an excellent opportunity to discuss the value of grain sizes for differentiating of ancient depositional environments. Several isolated Bahamian-type carbonate platforms (Plassen Formation) in the central Northern Alps are connected with deep-marine basinal sediments (Oberalm Formation) by slopes consisting of allochthonous carbonates represented by mass-flow sediments (Barmstein Formation: Steiger 1981) and fine-grained breccia (Tressenstein Formation; Hötzl 1966). Another open-marine carbonate platform of approximately the same age is the Sulzfluh limestone (Ott 1969) at the boundary between Switzerland and Austria. This limestone has been interpreted as inter- and subtidal deposition but can sometimes be reinterpreted as beach sediment in the light of grain-size patterns. The microfacies of these limestones has been studied in detail. Grain-size analyses were applied to bioand lithoclastic calcarenites. The methods used are: interpretation of cumulative curves, parameter sorting and parameter diagrams as well as Passega diagrams. Sulzfluh limestone: The carbonate platform exhibits two major facies, predominantly mud-supported limestones and predominantly grain-supported limestones. The latter consist of grain- and rudstones rich in skeletal grains (calcareous algae, foraminifera, mollusks, corals) and synsedimentary clasts. The grain sizes of these biolithoclastic carbonates were studied using the mean, standard deviation, skewness and kurtosis. Parameter diagrams (Fig. 6.5) were compared with the Passega diagram (Fig. 6.6B). For modern calcareous sands, the sinusoidal pattern in the sorting versus mean diagram observed in the Sulzfluh samples (Fig. 6.5) is explained as a function of the average grain size and indicates mixing of variously sized grain populations (Folk 1962). This pattern is common in beach sediments, but also occurs in other shallow- and deep-marine settings (Hubert 1964). The same Sulzfluh samples have been used to construct the
Fig. 6.5. Inclusive standard deviation σi versus mean of twenty samples from the Late Jurassic Sulzfluh limestone. Most biolithoclastic calcarenites (Fig. 6.6A) are moderately well to moderately sorted and fall within field B. Field B characterizes the pattern produced by calcareous bioclastic sands from a modern carbonate beach in Yucatan (Folk and Robles 1964). Field A shows the pattern of samples from adjacent shallow submerged areas. Sorting classification after Friedman (1962). Note: Another often used sorting classification (Folk and Robles 1958) differs somewhat in boundary values for moderately sorted samples. Modified from Ott (1969).
Passega diagram shown in Fig. 6.6B. This pattern corresponds to beach sediments, too. Does this interpretation make any sense? The biotic composition (green algae, abundant porostromate cyanobacteria) and strong micritization indicate the existence of a very shallow sea. Wide grain-size ranges and moderate sorting can be explained by mixing and reworking of sedimentary particles. Many samples show evidence of early marine-vadose diagenesis (gravitative cements, internal silt). The co-occurrence of skeletal grains and ‘black pebbles’ (Sect. 4.2.8) points to the existence of a very near vegetated coast line. These criteria fit into the mental image of a sunny humid environment with sand cays at the margin of an open-marine subtropical platform. Tressenstein limestone and Barmstein limestone: Most Tressenstein limestones are characterized by finegrained lithoclastic calcareous breccias and coarse calcarenites consisting of poorly rounded, microfacially strongly variable clasts. Common fossils are corals, chaetetid sponges, calcareous algae and foraminifera. The composition of the Barmstein limestones, which intercalate with the Tressenstein limestones and occur as distinct beds within the basinal Oberalm limestones, is similar to the Tressenstein limestone. Significant differences between Tressenstein and Barmstein limestones relate to the size and sorting of the lithoclasts as well as sedimentary structures, which point to the depo-
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Passega Diagram
Fig. 6.6. Passega diagram of Late Jurassic limestones from the Alps. The one percentile (φ 1, C) measures the coarsest grain in the deposit; the median (φ 50, M) measures the total population and considers the fine grains in the tail of the distribution. A and B – Sulzfluh limestone. A: Moderately sorted, coarseto medium-grained calcarenite consisting of large and small lithoclasts. Note that some lithoclasts are encrusted and include microfossils. Scale is 2 mm. B: Passega diagram. The field included by the straight line is the pattern based on number frequencies. The only slightly divergent shape of the other field (sample points not shown) results from the conversion of thin-section measurements into weight percentages after Friedman (1958). The C-M pattern is considered to be characteristic for beach sediments. Modified after Ott (1969). C: C-M diagrams for Barmstein limestone (I) and the Tressenstein limestone (II). Pattern I is regarded as being characteristic for deposition from turbidity currents, pattern II has been interpreted as a possible coastal deposit affected by tractive currents. Tractive currents are caused by waves and currents touching the bottom. Modified from Hötzl (1966).
sition of most Barmstein beds as turbidites (Steiger 1981). This interpretation is supported by the C-M pattern shown in Fig. 6.6C). The C-M pattern of the Tressenstein limestone indicates the transport of coastal material and the deposition by tractive currents. This picture would fit into a depositional model considering the Tressenstein limestone as a sediment deposited on the gentle foreslope of a platform or on the near-coast slope of the inner part of a carbonate ramp.
6.1.2.3 Significance of Grain-Size Studies of Carbonate Rocks Grain-size studies of carbonate rocks are applied to: • Limestone classification: All textural classifications and some classifications used in differentiating porosity types and limestone resources need to know the percentages of grains of specific grain sizes (Sect. 7.3.2 and 8.1). • Recognition of depositional processes and hydrodynamic conditions (see this section and Sect. 12.1.1). • Differentiation of reef complexes: Grain-size ranges and average grain sizes vary in traverses crossing backreef, reef and forereef of some modern reef complexes. Comparable patterns have been observed, e.g. in Triassic and Jurassic platform-margin reefs. • Information on diagenetic changes: The size of calcite and dolomite crystals in recrystallized carbonate rocks and the apparent sorting coefficients indicate the degree of neomorphic alterations (Germann 1963).
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• Grain-size ranges in combination with geochemical data offer a possibility to characterize marble types and to recognize the provenance of antique marbles (e.g. Benea et al. 1993). Basics: Grain-size analysis Burger, H., Skala, W. (1976): Comparison of sieve and thinsection techniques by a Monte Carlo model. – Computers and Geosciences, 2, 123-139 Carozzi, A.V. (1988): Carbonate rock depositional models. A microfacies approach. – 604 pp., Englewood Cliffs (Prentice Hall) Folk, R.L. (1966): A review of grain size parameters. – Sedimentology, 6, 73-93 Folk, R.L., Robles, R. (1964): Carbonate sands of Isla Perez, Alacran reef complex, Yucatan. – J. Geol., 72, 255-292 Folk, L.R., Ward, W. (1957): Brazos river bar: A study in the significance of grain size parameters. – J. Sed. Petrol., 27, 3-26 Friedman, G.M. (1958): Determination of sieve-size distribution from thin-section data for sedimentary petrological studies. – J. Geol., 66, 394-416 Friedman, G.M. (1962): Comparison of moment measures for sieving and thin-section data in sedimentary petrological studies. – J. Sed. Petrol., 32, 15-35 Friedman, G.M. (1979): Address of the retiring president of the International Association of Sedimentologists: Differences in size distributions of populations of particles among sands of various origin. – Sedimentology, 26, 3-32 Hötzl, H. (1966): Zur Kenntnis der Tressenstein-Kalke (OberJura, Nördliche Kalkalpen). – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 123, 281-310 Longiaru, S. (1987): Visual comparators for estimating the degree of sorting from plane and thin sections. – J. Sed. Petrol., 57, 792-794 McManus, J. (1988): Grain size determination and interpretation. – In: Tucker, M. (ed.): Techniques in sedimentology. – 63-85, Oxford (Blackwell) Merta, T. (1991): A new universal method of thin-section- to sieve transformation of granulometric data. – Acta Geologica Polonica, 41, 117-146 Piller, W.E, Mansour, A.M. (1990): The northern Bay of Safaga (Red Sea, Egypt): an actuopaleontological approach. II. Sediment analysis and sedimentary facies. – Beiträge zur Paläontologie von Österreich, 16, 1-102 Passega, R. (1964): Grain size representation by CM patterns as a geological tool. – J. Sed. Petrol., 34, 830-847 Pye, K. (ed., 1994): Sediment transport and depositional processes. – 398 pp., Oxford (Blackwell) Rizzini, A. (1968): Sedimentological representation of grain sizes. – Memoria della Società Geologica Italiana, 7, 65-89 Syvitski, J.P.M. (ed., 1991): Principles, methods and application of particle size analysis. – 368 pp., Cambridge (Cambridge University Press) Visher, G.S. (1969): Grain size distribution and depositional processes. – J. Sed. Petrol., 39, 1074-1106 Zanke, U. (1982): Grundlagen der Sedimentbewegung. – 402 pp., Berlin (Springer) Further reading: K106
Frequency Analysis
6.2 Frequency Analysis of Microfacies Data Thin sections of carbonate rocks exhibit marked differences in the frequency of grains, matrix, cements, and pores. The value of microfacies for interpreting ancient depositional environments, discussing controls on distributional patterns of organisms, and understanding diagenesis and porosity relies strongly on quantitative data. Compositional studies of modern and ancient carbonates illustrate that the abundance of different grain types (called ‘constituent composition’ by Ginsburg 1956) is of particular significance for reconstructing depositional systems. Sect. 6.2.1 reviews the methods commonly used in determining the relative frequency of constituent particles and groundmass by counting, estimating image analysis and other methods. The advantages of multivariate and semiquantitative methods applied to frequency data are discussed in Sect. 6.3.
6.2.1 Methods of Frequency Analyses The raw data for frequency analyses are derived from the study of thin sections, peels or microphotographs. The use of digital images or flatbed color scanners allowing overall and detailed views and measures of thin sections is particularly recommended for limestones which differ strongly in the distribution and frequency ranges of constituents (e.g. reef or slope carbonates).
6.2.1.1 Counting Frequently used methods Four methods of determining frequency (modal analysis) are commonly used: point-, line-, ribbon- and area-counting. These procedures are illustrated schematically in Fig. 6.7. Point counting is an area measure. Ribbon and area counting are number frequency measures. Both point counting and area counting measure the relative proportions of grains of different size, but in point counting where an area proportion rather than a number frequency is measured, the relative proportion of the coarser grains is emphasized relative to the finer ones. (1) In point counting (Fig. 6.7A), the most widely used method, a regular two-dimensional grid is used, and every grain (or matrix and cement) which falls under a grid point is counted. The estimated volume of a particle type is proportional to the counted number of
Point Counting
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of the diameter of the largest grain. Narrow grid spacing tends to overrepresent the large grains. Limestones with cm-sized grains should be measured in polished slabs using wide grid spacing.. What about a spar-filled gastropod shell? Particles can be registered in point counting analyses by (a) all grid points falling on the particle as well as on voids within the particle, or (b) only those grid points which fall on the particle but not on the voids included in the particle (Fig. 6.8). Voids are separately counted, e.g. as micrite or cement. The first case corresponds to grainbulk measurements, the second to grain-solid measurements. If intraparticle porosity should be emphasized and the amount of cement within grains, grain-solid measurements are recommended. The same may be necessary for coated grains exhibiting siliciclastic nuclei (Wilson 1965). The grain-solid definition is related to the weight; the grain-bulk definition to the volume of the grain. The magic number or how many grid points must be considered? The number of points necessary for a statistically reliable result depends on the distribution and size of the particles. Accuracy is related to selected grid density and the number of points counted. The mean deviation decreases with an increasing number of points and increases with an increasing percentage of particles, that is up to 50%, whereupon the mean deviation again drops. In order to keep sampling errors small, several hundred (at least 300 to 500) points must be counted; for limestones with variously sized and poorly sorted grains as many as 1000 points and more. Fig. 6.7. Comparison of A - point counting, B - line method, C - area method, and D - ribbon method used in evaluating frequency values. Solid black grains are those not counted. Note that ribbon counting measures only those grains which are totally or almost totally included within the ribbon (hatched area). After Middleton et al. (1985).
particles of that type (Chayes 1956). The number of grid points intersected by specific grains is proportional to the relative area of the grain. For detailed measurements under the microscope special oculars are used together with a specially adapted stage micrometer. Computerized point-counting (‘Prior’ point counter) facilitates the time-consuming work (Kobluk and Kim 1991, McKinney 1993). Several computer programs for point counting are available, e.g. Point Count ‘99, developed for reef monitoring programs. In practice, pointcounting of microfacies samples is done on photographs of thin sections or on the screen of reading fiches covered with a point grid. Grid spacings used are (a) larger than the diameter of the largest grain, (b) approximately equal to the diameter of the largest grain, or (c) approximately half or one quarter
Fig. 6.8. A: Dual point counting. Point counting measurements of grains deal with the entire grain (including not only the shell but also voids filled with sediment or with cements (grain-bulk measurements) or B only the solid parts of the skeletal grain (grain-solid measurements). In the first case, all voids are counted regardless of their infilling (indicated by large dots). In the second case intraparticle voids (open circles) are measured separately as internal infillings or intraparticle cement. It is necessary to state clearly which of the definitions has been used in the study. After Jaanusson (1972).
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Frequency Index
There is a magic number of 300 counts which is believed to provide sufficient accuracy of frequency measurements. This often used figure is based on an equation for counting heavy mineral grains and on frequency studies of modern foraminifers. The reliability of point counting results can be judged using the chart developed by Van der Plas and Tobi (1965). This chart permits the relationships between the number of points counted and the accuracy of the result to be estimated (Fig. 6.9). For a discussion of the number of point counts needed for sufficient accuracy see Patterson and Fishbein (1989). One immediate and obvious shortcoming of point-counting data is often that taken on their own they do not display meaningful groupings without statistical treatment. A general impression of the statical behavior of individual point-count groups can be obtained by calculating summary statistics for each variable of the pointcounting groups (see Sprenger and Ten Kate 1990 for the software): Range (minimum and maximum values), average (mean, median and mode), measures of dispersion (variance standard deviation and coefficient of variations), and the confidence limits for the population mean and variance. These statistics are important if variations along a geological section are to be discussed. (2) In line counting (Fig. 6.7B) all grains are measured that are intersected by lines across the section spaced at intervals approximately equal to the diameter of the largest particle. Line-counting of microfacies samples is less useful than the other methods mentioned, but is a common and time-saving method in field-based quantitative facies studies aiming at an understanding of paleoecological patterns. Results based on line counting and point counting can differ by as much as 10% (Flügel and Hötzl 1976; Bernecker et al. 1999). Another successful approach to line counting are measurements of facies and microfacies variations in cores (Sect. 17.1.4.2). (3) In area counting (Fig. 6.7C) all grains lying within a defined thin-section area are evaluated. Area counting is a genuine number-frequency measurement. The method can sometimes be compared with ‘coarse grain analysis’ used to investigate recent carbonate sand-sizes sediments (Sarnthein 1971). A very promising variant of area counting is the foraminiferal abundance analysis (Sect. 10.2.2.1). (4) Ribbon counting (Fig. 6.7D) requires measurement of all grains occurring within ribbons oriented parallel to the sedimentary bedding. Individual ribbons
Fig. 6.9. Chart for judging the reliability of point counting results after Van der Plas and Tobi (1965). The reliability is expressed by the standard deviation s. Key: N = total number of points counted; P = estimated percentage of a grain in the thin section. Full curves = 2 s values, dashed curves = relative s values. Example 1 (asterisk): Consider an oolitic grainstone. Out of 500 points counted 140 points (= 28%) fall on ooids. Plotting these values on the chart leads to a standard deviation of 2 s = 4% (dark line, absolute error). This means that the ooid content of the samples lies between 28 ± 4% = 24 and 32% with a 95% confidence. If this error is acceptable, the counting can be brought to a halt. If you agree with a larger standard deviation of, e.g. 5%, you can stop after counting 300 points. Note that the use of the chart requires a point distance that is larger than the largest grain included in the measurement.
are vertically separated by unassessed ribbons of a width approximately equal to or twice the diameter of the largest grain or just spaced randomly. Grains extending from above into the ribbon are counted; those extending from below into the ribbon remain uncounted. Frequency index: A useful variant of area counting Carozzi and his students developed a technique of quantifying microfacies data by means of sets of variation curves drawn alongside the stratigraphic column. Quantification starts by determining two indices, the
Comparison Charts
clasticity index (Sect. 6.1.2.2) and the frequency index. These indices are determined subsequent to a rough classification of the samples into microfacies types by means of qualitative criteria. The frequency index is counted as the sum of the number of specific particles occurring in six adjacent fields in the thin section arranged in two rows of three fields. In grain-supported rocks, the area required to be counted should be at least 10 times larger than the average size of the grains. In practice, a field of 300-400 square millimeters is commonly used. Each grain category is separately counted. Grains considered include bioclasts and fossils, intraclasts, ooids, and oncoids as well as detrital minerals (e.g. quartz, glauconite). The frequency index of bioclasts, transported fossils and detrital minerals is regarded as an indication of the traction or suspension load of transporting agents in the depositional environment. For non-transported benthic or planktonic fossils the index expresses the frequency of a given population. The frequency and clasticity index and rock composition are used to distinguish microfacies types. For shallow-marine limestone these types are arranged according to their position in an ideal shallowing-upward sequence corresponding to the assumption that all or most shallow-marine carbonate successions consist of superposed shallowing-upward sequences (Carozzi 1958, 1986). The method has been intensively used and was applied to carbonate ramps and platforms, reefs, mass flow deposits as well as to lacustrine limestones (see case studies in Carozzi 1988). The technique is time consuming and the presentation of the results is somewhat arduous; but if the results are statistically tested and put in the context of other facies criteria, the frequency index and the other indices used by Carozzi provide valuable information about short-term changes in water energy and the inferred paleobathymetry.
6.2.1.2 Estimating Counting methods described in Sect. 6.2.1.1 need a considerable amount of time when applied to a large number of samples. The time factor is often a large drawback to collecting and handling quantitative microfacies data. What should be done instead? Just use estimates as many people do? Spots in front of your eyes: Using comparison charts Estimations of the frequency of grains and matrix in thin-sections are common methods in facies analysis, although not as accurate as point counting or image analysis systems.
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Visual comparison charts for sedimentary rocks have been published by various authors (Folk 1951; Terry and Chilingar 1958; Baccelle and Bosellini 1965; Schäfer 1969). Matthew et al. (1991), confronted with the problem of quantifying archaeological ceramics, developed charts whose use in microfacies studies is especially recommended. The comparison charts were prepared in different ways: (a) by cutting black paper and scattering the pieces over a defined area (Terry and Chilingar), (b) by image presentation of point-counting results (Baccelle and Bosellini), or (c) by a computer program considering covering of grains and different ranges of grain size (Matthew et al. 1991). The comparison charts of Baccelle and Bosellini, Matthews et al. and Schäfer assist positively in reliable estimations of frequencies. The conceptual background of the charts is shown in Figs. 6.10 and 6.11. The Bacelle and Bosellini charts are included in the CDROM accompanying this book. Advantages and disadvantages of the charts: The most commonly used Baccelle and Bosellini chart has the advantage that it shows images which can be attributed to common carbonate grain types. A pitfall of the charts is, that variations in grain size are only insufficiently taken in consideration. Checking the images by digital image analysis result in somewhat different percentage values (± 5%), particularly for wellrounded grains. The Matthew et al. charts have the advantage that they show percentages in the context of grain size ranges. The latter include ranges which are common in calcarenites and calcirudites. The ranges comprise different ranges between 0.5 mm and 10 mm. If you want to use these charts, (1) select the appropriate charts according to the shapes and association of the particles (the charts display various round and mixed round and polygonal figures), (2) determine the grain-size ranges in your sample, and (3) decide whether it is better to use black/white or white/black charts. The charts proposed by Schäfer are particularly designed for limestones with shells or echinoderms, e.g. mollusk coquinas and crinoid limestones. Particles are shown in black on a white background. Point counting, visual comparison charts or free-lance estimating? Series of experiments carried out with students and postgraduates in Erlangen produced good results with regard to the comparability of point counting measurements and estimates. Rough estimates are somewhat lower than those using the Baccelle and Bosellini charts.
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Comparison Charts
Fig. 6.10. Examples from the ‘Comparison charts for visual percentage estimation’ developed for limestones by Baccelle and Bosellini (1965). The entire series of 24 charts can be found on the CD. The charts compare (a) different grain categories, (b) different percentage values and (c) the ‘psychological distortion effects’ which result from looking at white objects on a black or gray background and black objects on a white or gray background. The major grain categories are shown in the images A to F: A - silt-sized and fine-sand sized particles, e.g. peloids, small intraclasts, or microfossils; B - coarse angular and subangular particles, e.g. intraclasts and angular skeletal grains; C - equal-sized ooids; D - differently sized ooids or oncoids associated with smaller particles, e.g. peloids; E - oriented shells, e.g. bivalves; F - associations of various grains, here predominantly bioclasts, a few coated grains and peloids. White/black effect: Contrasting particles appear larger or smaller depending on the background; G and H display the same percentage of skeletal grains with a black and with a white background. Low and high values: Most charts start with images at 10%; only the charts for the groups A and B show images below 10%. The largest percentage value shown is 50 or 60%. I and J: Few and abundant peloids contrasted; K and L: Few and abundant bioclasts.
These results derived for oolitic and bioclastic grainstone samples, however, should not be overrated, because the questions raised in the exercise were rather simple and concerned only the amount of the total of all grains without any differentiation of grain types. Less favorable results were produced by experiments dealing with texturally inhomogeneous limestones and asking for estimates of the frequency of particular grain types occurring in highly variable abundances. Rare and less common grains were considerably under- and overrated. In addition, the error increased with decreasing grain size. Very small grains, e.g. some foraminifera or sponge spicula, occurring in low percentages (2 to 10% according to point-counting measurements) were commonly excluded both in free-lance estimates
and in the estimates using comparison charts. The compositional variability of these samples was revealed only by areal counting and point counting. Another pitfall of estimating methods is that thin-section diversity is often blurred because of the missing data of low-percentage biota. In summary, estimating methods have their merits if two-state compositions are being analyzed (e.g. grain/ matrix ratio or amount of visible porosity), or if only the relative frequency of two or three different grain categories is being sought. Constituent analyses, however, commonly based on all available thin-section data need counting data that allow point-counter groups to be treated statistically (Sect. 6.3).
Comparison Charts
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Fig. 6.11. Examples from ‘Visual comparison charts’ developed by Matthew et al. (1991) for studying ceramic materials. In contrast to the Baccelle and Bosellini charts, these charts also compare differences in grain sizes and sorting. They consist of 24 pairs of charts showing black grains on a white background or vice versa. The charts show two major groups - images with rounded grains and images with mixed rounded/polygonal grains. The grains range in size from 0.5 to 10 mm. The images display mixtures of different-sized grains, shown at percentages of 10, 25 and 50% and within different grain-size ranges. For additional information and comment see Appendix (Comparison charts).
6.2.1.3 Image Analysis Image analysis is important for quantifying textural criteria and the composition of rocks at various scales. Digital image analysis aids in determining the frequency, shape, rounding, and size of grains and pores and in determining statistical parameters that can be plotted in graphs. The data are derived from microphotographs, field photographs, images of thin sections or peels studied with digital video camera, scanner or digital SEM. The back-scattered electron imaging mode of the SEM provides fully automatized compositional and morphological analyses of sandstones (Ehrlich et al. 1984; Krinsley et al. 1998) and carbonate rocks (Rogen et al. 2001). Several computer-assisted systems are commercially available (e.g. Kontron IBAS, GIPSY, PETROG or DIANA 2001). Most image analysis software packages allow grains (Van den Berg et al. 2002), pores, and cements to be differentiated by recognizing
differences in ‘gray levels’ in an image of thin sections under plane light or cathodoluminescence. Image acquisition of thin sections studied under the microscope usually takes place via a video scanner. Image digitization then converts the analog TV signal into a digital form. Further treatment includes image segmentation with the purpose of enhancing the image and discriminating between the phases that are to be measured in the sample. Limestones and dolomites need specific preparation (e.g. etching, staining) in order to enhance the contrasts between different grain categories and to count frequencies (Harvey and Steinmetz 1971). Grain types, matrix, and cement in carbonates can be made visible by using methods developed for aerial image interpretation (Conley and Davis 1973). Potential of image analysis for carbonate studies: The method allows a wealth of different characteristics of carbonate rocks including color and weathering
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Integrated Frequency Studies
of carbonate rocks to be quantitatively collected (Muge et al. 1997; Lebrun et al. 1999): Depositional textures, e.g. the degree of bioturbation (Magwood and Ekdale 1994; Francus 2001), the relationship between microfacies of heterogeneous limestones and their suitability as building stones (Fronteau and Paicheler 1999), or the intensity of the skeletonization of frame-building reef organisms (Bernecker and Weidlich 1994). Image analysis can be used to characterize the geometry and frequency of porosity in carbonate rocks from random thin sections (Anselmetti et al. 1998; White et al. 1998; Frykman 1992). Frequency counting based on digital image analysis is a common method. The results correspond rather nicely to frequency measurements based on pointcounting in thin sections (Dorobek et al. 1987).
embrace fossils which can be identified in thin sections only at gross taxonomic levels (e.g. echinoderms, brachiopods, rugose corals, dasyclad algae), at group levels (e.g. small very thin-shelled ostracods vs. large ostracods with homogeneous microstructure), or, rarely even at generic taxonomic levels. PFD differs from diversity based on taxonomic richness primarily in the broader level of taxonomic identification. Consideration of fossil diversity in thin sections yields a good first approximation to simple taxonomic diversity and is, therefore, often used as a rule of thumb measure of environmental conditions, although this method can be dangerous if no other supporting data exist. PFD can be determined just by counting the number of fossil types, but statistically based comparisons of samples require percentage data produced by area, point or ribbon counting.
6.2.1.4 Constituent Ranking, Diversity and Maturity
Compositional maturity The extent to which a grain-rich carbonate sediment approaches the constituent end-members (compositional maturity: Smosna 1987) reflects the complexity of the processes that operate on carbonate sediments. End-members are synsedimentary clasts, ooids, skeletal grains, peloids, micrite, and terrigenous minerals. The proportions of samples consisting of these endmembers give some indications of major depositional controls. A grainstone composed almost totally of ooids and no other grain types may reflect conditions which are most favorable for the accumulation of ooids. A mixed grainstone with various skeletal grains and some ooids points to poor conditions for autochthonous ooid formation, but good and varying conditions favoring the occurrence of diverse biota.
The relative frequency of grain types, the diversity of grain types and the dominance of specific grain categories carry environmental signals which should be considered in microfacies studies (see Box 11.1): These approaches are based on identificating the relative amount of grains rather than purely numeric data. Constituent ranking This technique requires ranking grains according to visual estimates of their volumetric importance. The most important constituent is coded as rank 1, the second most important as rank 2, and so on. Rank 4 includes all grains present but those of lesser importance than 3. The ranking is independent of the proportions of matrix or cement and thus expresses relationships between grains. Shifts in ranking indicate gradual environmental changes particularly within defined units (e.g. Waulsortian carbonate mud mounds: Lees et al. 1985) containing various and diverse grains. A statistical treatment of ranking data using presence/absence data and looking for ‘relay’ conditions assists in recognizing natural transitions between qualitatively defined microfacies types (Sect. 14.4). Petrographic Fossil Diversity In analyzing ancient environments fossil diversity is generally regarded to be most significant. Diversity measures based on microfacies data are not identical with those of the usual paleoecological studies. Petrographic fossil diversity (PFD) is defined as the total number of all dissimilar fossil types observed in thin sections (Smosna and Warshauer 1978). ‘Fossil types’
6.2.1.5 Integrated Frequency Studies of Reef Carbonates The complexity and variability of reef carbonates requires studying abundance and distribution of biota on different scales. Field methods used in modern reefs to describe ecological zonation or monitoring changes in the diversity include plotting techniques using quadrats (arranged randomly or in belts) as sampling and counting areas (corresponding to area or ribbon counting), as well as plotless techniques utilizing line transects with defined lengths or grids (corresponding to line- and point-counting). Quadrat-belt methods exhibit the most complete data sets with respect to taxonomic diversity, size and growth form of reefbuilders, and coverage (Sullivan and Chiappone 1992; Vogt 1995; Segal and Castro 2001).
Frequency Analysis: Practical Advises
Compositional studies of ancient reef carbonates are strongly biased towards microfacies-scale observations. Thin sections often characterize microenvironments rather than the entirety of communities occurring in an outcrop. Microfacies-scale data are invaluable mosaic stones in reconstructing development stages of ancient reefs. These data, however, should be combined with outcrop-based quantitative analyses using methods similar to those used in investigating modern reefs. Most authors use line- or point-counter techniques (e.g. Flügel and Hötzl 1976; Königshof et al. 1991; Perrin et al. 1995). One successful approach is mapping the reef fabrics on plastic sheets at a natural scale in the outcrop, followed by digital image analysis of these sheets (Bernecker et al. 1999). These quantitative data document the biotic composition and development of a reef structure (Weidlich et al. 1993; Fagerstrom and Weidlich 1999). Frequency analyses (area and point counting) of microfacies samples taken within the studied quadrats (Webb 1999) provide information on the composition of the reef sediment and the distribution of microbiota, and describe biotic interactions (e.g. role of encrusters). Similar to modern reefs (Piller 1994), microfacies allows subenvironments and the origin of the fine-grained matrix to be differtiated in greater detail.
6.2.2 Practical Advice Frequency analyses are very useful in evaluating microfacies data if the following points are taken into consideration: • Make a clear decision about which method is most appropriate to your problem. Be consistent in treating all your samples with the same methods. The choice of the method depends on the questions you are asking. • State precise questions at the beginning of your investigation. What is the final goal of your study? Do you need some quantitative data to be sure that you are right about the textural classification of your sample? Should frequency data be used to understand porosity evolution? Or are you more interested in reconstructing paleoenvironments and depositional sites? If you use frequency data just to classify and categorize samples, you will need only a few percentage values on the proportions of grains and groundmass and the frequency of the most abundant grain types (Sect. 8.1). The same holds for a first differentiation of porosity types (Sect. 7.3). Visual estimating methods will often be sufficient for satisfactory results. But if you are aiming at a mul-
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tivariate treatment of constituent analysis data, you will need many more values indicating the abundance of many constituent grains. Here, point counting or ribbon counting are superior to other methods. Image analysis is especially successful in porosity differentiation and combined outcrop- and microfacies studies. • Draw a well-defined line between particles and matrix. Very small particles are usually difficult to assign to specific grain categories. Most people include grains < 0.05 mm within the groundmass (see Jaanusson 1972 for a discussion of this problem). • Use ribbon or area counting for limestones consisting of coarse-sized calcirudites and breccias. • If you use point counting, think about which microfacies criteria are useful in answering your questions. A pre-selection of variables may be appropriate if you are intending to determine point count groups that reflect genetically related criteria (e.g. all biotic data which indicate that some of the grains within a calciturbidite were derived from the reef front and other grains from lagoonal sands). Point-counting is commonly carried out in thin sections cut perpendicular to the bedding planes. A minimum size of 5 x 5 cm is strongly recommended. • If you use the visual comparison charts of Baccelle and Bosellini, note that larger grains may be overestimated up to 10%, that the frequency of white grains within a dark matrix is frequently overestimated, and that underestimations may occur in samples yielding only a few grains within a ‘white’ matrix. • Be sure that you or your students are able to distingish grain categories adequately and that the operator errors are statistically tolerable. Try to assign as many grains as possible to specific categories, but do not hesitate to use pigeonholes for grains which can not be categorized in detail (e.g. ‘skeletal grains of unknown systematic affinity, ‘sparry grains, most probably recrystallized shells’, ‘micrite grains, may be peloids, intraclasts or totally micritized bioclasts’). • Do not be anxious about accurately determining skeletal grains. Constituent analyses aiming to recognize environmental controls and the depositional mosaic (e.g. on carbonate ramps) profit greatly from data reflecting the relative frequency and co-occurrence of specific biotic elements, such as species, morphotypes or growth types. • But consider that the abundance of bioclasts within a particular limestone type is strongly controlled by taphonomic factors, especially by skeletal architecture, environmental factors, and the time involved in taphonomic and diagenetic processes. A low preservation
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record of originally abundant organisms may be expected in environments affected by strong water energy (e.g. reef environments), lack of sedimentation (e.g. in deeper hardground areas) and in areas with strong bioturbation (e.g. deeper-water mud). The frequency of echinoderm fragments, therefore, may not reflect true abundance, but can be a measure of biodiversity (Nebelsick 1996). • Consider diagenetic alterations (e.g. partial or complete dissolution) of skeletal and other grains that can lead to misinterpretations of the primary composition of the sample. • Note that small silt- and clay-sized terrigenous grains (e.g. quartz) or small microfossils might not be visible in thin sections but may be abundant in acid residues. • If appropriate use adjective terms in describing the relative frequency of grain categories: Very rare (< 2% grains), rare (2–5%), sparse (5–10%), common (10– 30%), very common (30–50%), abundant (> 50%). For semiquantitave methods and estimations use: absent, rare, present, abundant. • Note the percentage of thin sections in which the constituent is present. Basics: Methods of frequency analysis Baccelle, L., Bosellini. A. (1965): Diagrammi per la stima visiva della composizione percentuale nelle rocce sedimentarie. – Annali della Università di Ferrara, Sezione IX, Science Geologiche e Paleontologiche, 1, 59-62 Bernecker, M., Weidlich, O., Flügel, E. (1999): Response of Triassic reef coral communities to sea-level fluctuations, storms and sedimentation. Evidence from a spectacular outcrop (Adnet, Austria). – Facies, 40, 229-280 Chayes, F. (1956): Petrographic modal analysis. – 113 pp., New York (Wiley) Jaanusson, V. (1972): Constituent analysis of an Ordovician limestone from Sweden. – Lethaia, 5, 217-237 Matthew, A.E., Woods, A.J., Oliver, C. (1991): Spots before the eyes: New comparison charts for visual percentage estimation in archaeological material. – In: Middleton A., Freestone, J. (eds.): Recent development in ceramic petrology. – British Museum Occasional Paper, 81, 211-264 Middleton, A.P., Freestone, I.C., Leese, M.N. (1985): Textural analysis of ceramic thin sections: evaluation of grain sampling procedures. – Archaeometry, 27, 64-74 Patterson, R.T., Fishbein, E. (1989): Re-examination of the statistical methods used to determine the number of point counts needed for micropaleontological quantitative research. – J. Paleont., 63, 245-248 Russ, J.C. (1999): The image processing handbook. 3rd edition. - 771 pp., Boca Raton, Florida (CRC Press) Schäfer, K. (1969): Vergleichs-Schaubilder zur Bestimmung des Allochemgehalts bioklastischer Karbonatgesteine. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1969, 173-184 Soille, P. (ed., 2000): Morphological image analysis. – 314 pp., Berlin (Springer) Van der Plas, L., Tobi, A.C. (1965): A chart for judging the reliability of point counting results. – American Journal of Science, 263, 87-90
Multivariate Analysis
Weidlich. O., Bernecker, M., Flügel, E. (1993): Combined quantitative analysis and microfacies studies of ancient reefs: an integrated approach to Upper Permian and Upper Triassic reef carbonates (Sultanate of Oman). – Facies, 28, 115-144 Further reading: K108
6.3 Multivariate Microfacies Studies Modal analyses and other frequency measurements supply the raw data for multivariate microfacies analyses. These analyses are not intended to replace the evaluation of data derived from qualitative data, but rather aim at simplification and efficient ordering of datasets. Statistical methods shed a new, ‘more objective’ light on problems related to classifying samples and interpreting compositional data, or may point out aspects which would not come out in a conventional approach to microfacies studies. Multivariate methods may be divided two broad groups: Classification refers to the recognition of discrete groups within a data set and therefore assumes discontinuity. The outcome of cluster analysis is a two-dimensional diagram, in which no relationships between the clusters and underlying environmental gradients can be discerned directly. Ordination aims to order or arrange samples or variables in a space in relation to environmental parameters or gradients. Ordination extracts major directions of variations from the original data, which are assumed to be in continuity. The outcome of ordination are two- or three-dimensional scatter plots which need intelligent and sound interpretations. 6.3.1 Methods: Variations between Constituents and between Samples Multivariate analyses start by creating a numerical element data matrix, based on frequency measurements. This matrix serves as basis for calculating correlation coefficients that describe similarity between variables and between samples. Various correlation coefficients (e.g. the product moment correlation) are used to establish a similarity matrix. This data matrix is used for cluster analyses in answering the questions: how are variables (grains) related statistically and how are samples related. Cluster and factor analyses applied to microfacies Variable-per-variable relations are shown by R-mode cluster analysis, sample-per-sample relations by Qmode cluster analysis derived from point-counting data. Variables are often standardized in order to negate the strong effects of magnitude. The degree of similarity
Multivariate Analysis
between variables or pairs of samples is expressed by various similarity coefficients. Similarity measures take two forms, quantitative or binary. Shi (1993) lists 39 measures with the goal of measuring resemblance between variable sets or all possible pairs of samples. Cluster analyses result in a grouping of variables (facies criteria) or samples with regard to their similarity and correspondence. The step-wise reduction of the most similar groups leads to a hierarchical classification best shown in a two-dimensional dendrogram (Fig. 6.12). Cluster analysis dendrograms are helpful in grouping samples with similar component distributions, but should be used for facies differentiation only in the context of evaluating qualitative microfacies criteria. The question of how groupings of facies data are related to independent factors can be approached by factor analysis. Factor analyses are powerful tools for (a) determining the minimum number of causal relationships needed to explain the majority of the observed microfacies variations, (b) to identify these causal relationships, and (c) to estimate the relative importance of microfacies variables. In facies studies, factor analysis is used to derive causal factors controlling the distribution of sedimentary constituents. This technique is used for resolving complex relationships within a set of variables or samples by constructing hypothetical factors which reduce the overall complexity. The factors are ranked according to the extent to which the factor accounts for the variance observed in the data set. An often used technique for extracting significant factors which can be interpreted in terms of facies controls, is principal component analysis. A visual inspection of factor plots indicates which objects are grouped together and may result from the comparison process. Different factors explain different percentages of variance observed in the data. Positive or negative factor loadings of the variables reflect the controls in the interrelationship of variables, e.g. biotic composition, hydrodynamic conditions and terrigenous input (Fig. 6.13). Choose between quantitative numerical and binary data! The use of quantitative methods has a few drawbacks: Important non-quantitative attributes, e.g. rock color, are excluded. Poorly preserved constituents, e.g. recrystallized bioclasts, often can not be allocated to a grain type category as would be desirable for pointcounting. Modal analysis of samples with large-sized particles require the use of large, oversized thin sections. Last but not least: point counting is rather time
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consuming, even with the appropriate computer equipment. These drawbacks can be overcome by a method using two-state or multi-state attributes (Klovan 1964; Bonham-Carter 1965). This method uses binary data. These data indicate the presence or absence of attributes, and are coded as 2 or 1; the existence of uncertain attributes is coded as 0. Multistage quantitative variables are subdivided into ranks (e.g. abundant, present, rare, absent, uncertain) and also are coded as simple presence/absence attributes. There are no general rules as to how the attributes should be divided and coded. Visual estimation is usually quite satisfactory. Binary data analysis includes (a) collection and establishment of a raw data matrix; (b) calculation of measures of similarity (e.g. Jaccard coefficient) either between all possible pairs of characters (R-mode technique) or between all possible pairs of samples (Qmode); (c) clustering of samples, and (d) plotting of dendrograms. This method and also the various numerical classification methods were tested using data sets from the Bahama Bank (e.g. Ramsayer and BonhamCarter 1974). Binary data have been successfully applied in facies discrimination particularly with the aid of association analyses (Gill 1993). Association analysis, developed to handle binary-coded multivariate data of plant communities, was used to discriminate modern carbonate sediments (Erez and Gill 1976), and to recognize ancient reefal paleoenvironments of Cretaceous and Miocene reef complexes (Jurassic: Neuweiler 1995; Cretaceous: Ekdale et al. 1976; Miocene: Buchbinder 1979). Classifications based on numerical frequency distributions lead to a breakdown of continua into discrete units, e.g. ‘standard microfacies types’. These units, however, do not reflect compositional relays characterized by overlapping compositions, caused by shifting environmental gradients. These relays become evident after the binary data have been treated by using optimized similarity matrices and correspondence techniques (Hennebert and Lees 1985; Sect. 14.4).
6.3.2 Significance of Multivariate Studies: Constituent Analysis as a Clue to Environmental Conditions and Depositional Settings Constituent grain analysis was developed as a tool for differentiating modern sedimentary environments. The technique relies on the assumption that in shallow-marine environments a high proportion of carbonate sediments is produced in situ as skeletal remains of local communities. The close correlation between spatial pat-
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Box 6.3. Selected examples of multivariate microfacies analyses of ancient and modern (*) carbonates. Cluster analysis: Behrens 1965; Boss and Lydell 1987*; Carannante et al. 2000; Carcuel et al. 1998; Cugny 1979; Cunningham and Collins 2002; Demirmen 1971; Everts et al. 1999; Ferretti 1989; Greb et al. 1996; Khaiwka et al. 1981; Nebelsick et al. 2000, 2001; Flood and Orme 1977; Parks 1967; Perry 2000*; Piller 1994*; Pandolfi et al. 1999; Purdy 1963*; Radke 1976; Rasser 2000; Reijmer et al. 1922,1994; Schott 1984; Smosna and Warshauer 1978, 1979; Utescher 1992; Zuschin and Piller 1994* Factor analysis: Cugny and Rey 1981; Fenninger 1970; Flügel and Hötzl 1976; Griffith et al. 1969; Harbaugh and Demirmen 1964; Imbrie and Purdy 1962; Jöreskog et al. 1976; Klovan 1964; Montaggioni et al. 1986*; Nebelsick 1989, 1992; Osborne 1967, 1969; Purdy 1963*; Rao and Amini 1995; Rao et al. 1973; Sarnthein and Walger 1973*; Toomey 1966 Correspondence analysis: Bonham-Carter et al. 1986; Hennebert and Lees 1991; Melguen 1973, 1974; Montaggioni et al. 1986*; Reijmer 1998; Reijmer and Everaars 1991; Reijmer et al. 1994.
terns of skeletal grain types, distribution of sedimentproducing organisms, and the quantitative composition of carbonate sediments is widely applied when studying large-scale, cross-shelf environments or documenting small-scale facies shifts. Many of these studies rely strongly on multivariate analyses of grain composition data; others concentrate on a semiquantitative survey of distribution patterns across depositional environments (e.g. reef belts or on-coast to off-coast shelf traverses: Ginsburg 1956; Boichard et al. 1985). Constituent analyses of ancient carbonates and distributional patterns of specific grain types address several major goals of microfacies analysis, which are discussed in detail elsewhere in this book: • differentiating small-scale depositional environments within ramp and platform carbonates (Sect. 14.2.2), • fingerprinting records of lost platform sediments recorded in allochthonous carbonates of slope and deepwater settings, recognizing sealevel fluctuations and correlating carbonate platform-to-basin settings by means of grain-composition logs (Sect. 16.3), • discriminating ecological zonations in reef complexes (Sect. 16.2), • recognizing growth phases of carbonate mounds (Lees and Miller 1985), • understanding climatic controls and differentiating tropical and non-tropical carbonates (Sect. 16.4) • evaluating lithologic and geophysical data of wells to understand and predict favorable stratigraphic reservoirs (Sect. 17.1.5).
Example for Cluster Analysis
>>> Fig. 6.12. Facies differentiation supported by cluster analysis. This case study concerns Late Eocene shallow-water carbonates deposited in the Alpine foreland basin (‘Molasse zone’) in Upper Austria. These carbonates are targets of hydrocarbon exploration. Samples from deep wells exhibit bioclastic limestones with highly diverse combinations of different skeletal grains (coralline red algae, foraminifera, bryozoans, corals) and various amounts of terrigenous quartz. Facies discrimination reflecting depositional and environmental constraints was only possible with the use of multivariate methods. Methods: The composition of 120 thin sections of the cores (5 x 5 and 12 x 6 cm) was analyzed by point counting (600 points, grain-solid method). Statistical analyses include calculation of Spearman’s Correlation Coefficient, hierarchical cluster analysis and factor analysis as well as Markhov chain analysis. The dendrograph shows the result of R-mode analysis reflecting the similarity in the modal composition of samples from four wells separated by a distance of between 7 and 15 km. Note the decreasing similarity of sample groups to the right. The analysis resulted in recognizing two main clusters: A-H, dominated by red algae, and I-N, characterized by higher amounts of terrigenous detritus and lower amounts of coralline algae. Clusters A-C are characterized by the joint occurrence of corallinacean and peyssonneliacean red algae, clusters D-G combine samples dominated by coralline branches (D) and those dominated by coralline detritus (EG). Facies types with a predominance of coralline algal detritus (indicated by black bar) are comparable with the presentday ‘maerl’ facies of temperate seas (Sect. 2.4.4.3). Cluster H unites all samples with more than 24% rhodoliths. Note that this cluster is indistinctly separated from the clusters AG, although it occurs within the same main cluster that combines other samples with red algae. Clusters I and J are characterized by large amounts of terrigenous quartz. Clusters KN are united by the occurrence of bryozoans and differ in the frequency of coralline algae, benthic foraminifera and bryozoans. The evaluation of the facies types, taking into consideration ecological data, the location of the wells and the stratigraphical succession of the samples in the cores resulted in a consistent facies model that described the depositional history of the basin during time. Modified from Rasser (2000). Photographs show some of the microfacies types: A – Terrigenous Coralline Detritus-Peyssonneliacean Facies. Components are fragmented and well-rounded coralline algal thalli (black; Mg-calcite), unrounded peyssonelliacean algae (gray; aragonite) and angular quartz grains (white). The facies is characterized by the absence of quantitatively clearly dominating grains. Geinberg. F – Coralline Detritus-Coral Facies. Coral-bearing algal rudstone with packstone matrix. The branched corals are encrusted by coralline algae. Maria Schmolln. J – Foraminiferal Quartz Sandstone Facies, dominated by nummulitid foraminifera embedded within a matrix consisting of silt-sized quartz grains, cemented by calcite. Larger foraminifera and coralline algae are negatively correlated, probably reflecting controls by substrate types. Helmberg. N – Bryozoa Facies with unilaminar erect growth forms. Fine-grained sandy matrix with glauconite. Helmberg. Basics: Multivariate microfacies studies Birks, H.J.B. (1985): Recent and possible future mathematical developments in quantitative paleoecology. – Palaeogeogr., Palaeoclimat., Palaeoecol., 50, 107-147
Example for Cluster Analysis
Cheetham, A.H., Hazel, J.E. (1969): Binary (presence-absence) similarity coerfficients. – J. Paleont., 43,1130-1136 Davis, J.C. (1986): Statistics and data analysis in geology. 2nd edition. – 646 pp., New York (Wiley) Flügel, E., Hötzl, H. (1976): Palökologische and statistische Untersuchungen in mitteldevonischen Schelf-Kalken (Schwelmer Kalk, Givet; Rheinisches Schiefergebirge). – Bayerische Akademie der Wissenschaften, Mathematischnaturwissenschaftliche Klasse, Abhandlungen, 156, 1-70 Gill, D. (1993): Discrimination of sedimentary facies by association analysis. – Mathematical Geology, 25, 471-482 Hennebert, M., Lees, A. (1985): Optimized similarity matrices applied to the study of carbonate rocks. – Geological Journal, 20, 123-131
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Hohenegger, J. (1974): Über einfache Gruppierungsverfahren von Fossil-Vergesellschaftungen am Beispiel obertriadischer Foraminiferen. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 146, 263-297 Imbrie, J., Purdy, E.G. (1962): Classification of modern Bahamian carbonate sediments. – American Assoc., Petrol. Geol., Mem., 1, 253-272 Lees, A., Hallet, V., Hibo, D. (1985): Facies variation on Waulsortian buildups, Part1: A model from Belgium. – Geological Journal, 20, 133-158 Nebelsick, J.H. (1992): Component analysis od sediment composition in Early Miocene temperate carbonates from the Austrian Paratethys. – Palaeogeogr., Palaeoclimat., Palaeoecol., 91, 59-69
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Example for Factor Analysis
Fig. 6.13. Factor analysis reduces the overall complexity of data by extracting factors which account for the variance observed in the data and facilitates the evaluation of the causal factors which control sediment composition. In this case study from the Miocene Zogelsdorf Formation northwest of Vienna, Austria, the R-mode principal component analysis of point counts in 118 thin sections is used to deduce environmental controls. The Zogelsdorf Formation represents carbonate deposits formed in a warm temperate environment (Pl. 93/7, Pl. 106/3). Factor loading of individual components are weighted with respect to the contribution of that component to the factor. Factor scores estimate the contribution of the factors to each original observation (the samples). In the case studies four factors were found to be primarily responsible for the distribution of skeletal grains, carbonate mud and terrigenous grains. The interpretation of these factors is based on the relationships of these components within the factors, the ecological constraints of the calcareous organisms, the distribution of the components within the geological section, and the comparison to modern environments. A – Factor loading of factor 1 vs. factor loading of factor 2. Factor 1 explains 26% of the variance and shows a highly positive loading of bryozoans, brachiopods and serpulids along with a positive loading of echinoderms. A highly negative loading is shown by terrigenous components. Factor 2 explains nearly 17% of the variance and is characterized by a very highly positive loading of barnacles and epifaunal bivalves, and a highly negative loading of foraminifers. B – Factor 3 vs. factor 4 plot. Factor 3 explains 14.3% of the variance; echinoderms and foraminifers exhibit a highly positive loading, the muddy matrix a highly negative loading. Factor 4 represents 10.4% of the variance; coralline algae are highly positively loaded. Negative loading is shown by terrigenous components. Interpretation: Primary controls responsible for the component distribution were terrigenous input, hydrodynamic energy, presence of seagrass and light intensity. Factor 1 reflects the detrimental effect of high terrigenous input to filter-feeding bryozoans, brachiopods and serpulids. Factor 2 is seen to reflect the exposure to high energy water conditions. The interpretation is supported by the dominance of barnacles and epifaunal bivalves and the occurrence of this facies in coastal settings. Factor 3 should explain the positive loadings for echinoderms and foraminifers. The presence of seagrass can be derived from the increased occurrence of regular echinoids and of epiphytic foraminifera. Factor 4 with a highly positive loading for coralline red algae could reflect light intensity related to water depth and suspension loads. C – Geological section and changes in factor scores during time. Modified from Nebelsick (1992).
Purdy, E.G. (1963): Recent calcium carbonate facies of the Great Bahama Bank. I. Petrography and reaction groups. – J. Geol., 73, 334-355 Rasser, M.W. (2000): Coralline red algal limestones of the late Eocene Alpine foreland basin in Upper Austria: component analysis, facies and palaeoecology. – Facies, 42, 59-92 Shi, G.R. (1993): Multivariate data analysis in palaeoecology and palaeobiogeography - a review. – Palaeogeogr., Palaeoclimat., Palaeoecol., 105, 199-234
Swan, A.R.H., Sandland, M. (1995): Introduction to geological data analysis. – 446 pp., Oxford (Oxford Press) Toomey, D.F. (1966): Application of factor analysis to a facies study of the Leavenworth limestone (PennsylvanianVirgilian) of Kansas and environs. – Special Distribution Publication, Kansas Geological Survey, 27, 1-28 Further reading: K147
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7 Diagenesis, Porosity, and Dolomitization
The preceding chapters provided information on the diagnostic criteria for ‘depositional microfacies’ . This chapter deals with ‘diagenetic microfacies’ that reflects changes during lithification and rock history (Pl. 28). Diagenesis refers to physical, chemical and biological processes. The understanding of these processes and their products has high economic importance, because diagenetic criteria account for many of the petrophysical properties of carbonate rocks and determine their value as reservoir rocks and use in industry. Diagenetic studies require a combination of various methods, including standard optical petrography, cathodoluminescence, SEM observations, stable isotope analyses and rare element composition. The following text concentrates on diagenetic features that can be studied in thin sections. Comprehensive reviews of carbonate diagenesis can be found in textbooks listed under Basics in the reference list at the end of this chapter. The chapter starts with a summary of carbonate mineralogy and diagenetic processes, followed by a survey of the major diagenetic environments. The treatment of porosity focuses on thin-section criteria and porosity types. We then take a look at how to recognize and interpret carbonate cements and diagenetic textures, discuss recrystallization and neomorphism, and examine the criteria of calcite spar formed by cementation or recrystallization. The analysis of dolomitization is limited to descriptive criteria, classifications, and an overview of some of the more common dolomitization models. The last part of this chapter deals with metamorphic carbonates and marbles.
7.1 Carbonate Mineralogy and Diagenetic Processes This section summarizes attributes of common carbonate minerals and discusses diagenetic processes affecting these minerals. The prevailing mineralogies of marine carbonate constituents have changed during the Phanerozoic; the last part of the section emphasizes the relevance of these secular variations for understanding cementation and porosity development.
7.1.1 Modern Carbonate Sediments and Ancient Carbonate Rocks Modern carbonates: Two carbonate minerals prevail in recent sediments, orthorhombic aragonite and trigonal calcite. Calcite shows marked differences in the magnesium content. Low-Mg calcite (LMC) and HighMg calcite (HMC) are usually separated by a value of 4 mol% MgCO3. The mineralogical composition of modern carbonate sediments depends on that of the skeletal and non-skeletal grains and the mineralogy of early cements. Common minerals are aragonite, magnesian calcite, followed by subordinate minerals like calcite and dolomite. Stable Low-Mg calcite dominates in many non-marine carbonates and is by far the most abundant carbonate mineral in deep-sea carbonates. The mineralogical composition of tropical and non-tropical shelf carbonates is strongly controlled by water temperatures. Shallow-marine tropical carbonates are mainly composed of metastable aragonite and calcite with high Mg concentrations. Non-tropical carbonates consist predominantly of High-Mg calcite (and minor aragonite) in shallow warm-temperate settings, and HMC and LMC in cool-temperate environments. Carbonate rocks: Since aragonite is metastable and HMC looses Mg with time, all carbonate sediments are eventually transformed to LMC. Carbonate rocks therefore consist of calcite and to a lesser extent, of dolomite. The mineralogical composition controls the diagenetic potential that characterizes the varying manner in which carbonates respond to diagenetic processes (Schlanger and Douglas 1974). Sediments with different mineralogical composition require different time intervals to reach equal stages in lithification. Sediments consisting of a mixture of aragonite and HMC (e.g. shallow-marine sands) as well as evaporite sediments have a high diagenetic potential because these minerals are easily and intensively dissolved or transformed. In contrast, sediments dominated by HMC and/or LMC (e.g. pelagic chalks) have a lower diagenetic potential. Because the diagenesis of carbonate and silica is interrelated at least in pelagic carbonates, the role of siliceous
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_7, © Springer-Verlag Berlin Heidelberg 2010
Stability of Carbonate Minerals
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constituents for the diagenetic potential should be thoroughly examined. Replacement of carbonate by silicification and the formation of cherts can result in an excess surplus of carbonate (Hobert and Wetzel 1989). Stability of carbonate minerals: The stability of carbonate minerals is different under marine or meteoric conditions. Textbooks favor a stability ranking for marine carbonates with calcite > Mg-calcite (<12 mol% MgCO3) > aragonite > very high Mg-calcite (>12 mol% MgCO3) and for meteorically influenced carbonates a ranking with calcite > aragonite > High-Mg calcite. However, laboratory experiments (e.g. Walter 1985) and preservation patterns of fossils show that these rankings describe only frequently observed stability successions
and do not allow for deviations. The commonly accepted sequence of diagenetic alteration (Mg-calcite > aragonite > calcite) is observed only in solutions supersaturated with respect to both calcite and aragonite. In solutions undersaturated with respect to calcite, aragonite skeletal grains with complex microstructures dissolve more rapidly than Mg-calcite. In solution supersaturated with respect to calcite but still undersaturated with respect to aragonite, mineralogic controls become more significant and Mg-calcites containing ≤ 12 mol% MgCO3 dissolve at rates similar to or greater than those for aragonites. Relations between carbonate minerals and organic matter: The precipitation of carbonate minerals is con-
Plate 28 Integration of Depositional and Diagenetic Microfacies Microfacies thin sections of carbonate rocks exhibit depositional criteria reflecting environmental constraints acting during sedimentation, and diagenetic criteria brought about by processes affecting carbonate sediments and rocks after initial deposition until after lithification. These processes take place in freshwater (meteoric), marine and burial environments and are recorded in thin sections by criteria related to pore-filling cementation, compaction and pressure solution, recrystallization, and dolomitization. The plate documents the integration of depositional and diagenetic features. 1 Reef limestone. The bulk of many reef limestones consists of reef-derived detritus rather than the autochthonous organic frameworks as documented by this sample. Depositional microfacies: The thin section shows skeletal grains, organic crusts, sediment and carbonate cements. Bioclasts are represented by a gastropod (transverse section; black arrow points to the inner layer, white arrows to the outer shell layer), corals (C) and a dasyclad green alga (D; transverse section of Diplopora). Biogenic crusts are microbial spongiostromate crusts (SPC) consisting of micritic and peloidal laminae of different thickness. The crusts exhibit biogenic overgrowth (BO) made of foraminifera, tiny tubiform fossils, sponges and microbes. Fine-grained sediment (S) within and between the skeletal grains consists of allomicrite with small shell fragments (lower part of the graded internal sediment in the gastropod) and microbial automicrite (evidenced by peloid microstructure and bacterial microfossils visible only at high magnifications). The sequence of depositional events comprises sedimentation of bioclasts, followed by biogenic overgrowth and microbial encrustation on skeletal grains resulting in a total occlusion of most of the interskeletal pores. Note the geopetal fabric formed by frozen spirit levels within the gastropod and in other small voids. Diagenetic microfacies criteria are cements, neomorphic alterations, dissolution features and fracturing. Calcite cements occur predominantly within the shell interior (marine radial fibrous cement, RFC, and late diagenetic recrystallized blocky cement, BC) and in some tiny voids within the matrix. Aragonitic fossils (corals and the inner layer of the gastropod shell, black arrow) were replaced by coarse sparry calcite. The original skeletal mineralogy is reflected by a different type of preservation: The originally calcitic outer shell layer (white arrows) differs in the finer crystal size and contains relicts of the microstructure. Note the different preservation of the primarily aragonitic gastropod shell and the originally High-Mg calcitic microbial automicrite crusts (SPC), both of which are now preserved as Low-Mg calcite. The reef limestone was affected by near-surface and marine diagenesis (inversion of aragonite; radiaxial fibrous cement) and burial diagenesis (coarse blocky calcspar cement). Mechanical compaction was hindered by rapid cementation of synsedimentary organic encrustation. Breakage and dissolution seams within the gastropod shell and between the gastropod and the coral point to incipient pressure solution. The complex microcrack network is of tectonic origin. Late Triassic (Dachstein reef limestone, Norian): Gosaukamm, Austria. Crossed nicols. 2 Lagoonal limestone. The sample represents a slightly dolomitized skeletal rudstone formed in a restricted lagoonal innershelf environment. Depositional microfacies: The sediment consists of dasyclad green alga (star-shaped cross section of Clypeina, CL), and various coated grains and pisoids characterized by closely spaced laminae (long arrow). Diagenetic microfacies: Rare isopachous fibrous cements (short arrows) indicate diagenesis in a marine phreatic environment. Intergranular voids are filled by clear granular cements (GC) representing former drusy cements that underwent crystal coarsening (aggrading neomorphism) during burial. Dolomitization postdates cementation as shown by the occurrence of idiomorphic dolomite rhombohedrons within micrite, grains and the cement of interparticle pores. Late Jurassic (Tithonian): Trnovo, southwestern Slovenia. ––> 2: Koch et al. 1989
Plate 28: Depositional and Diagenetic Microfacies
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trolled by physicochemical constraints and triggered by biological factors such as organic matter (OM) present in carbonate cements, skeletal and non-skeletal grains, and carbonate skeletons. OM occurs within grains and on grain surfaces and has been found in pores between submicroscopic subunits of cement crystals (Barnes et al. 1990a, 1990b for reviews). Current research is focussed on the location of OM within the carbonate host, the nature of OM associated with carbonates of different origins, and the influence of OM on the growth of inorganic and organic carbonate crystals (e.g. Ramseyer et al. 1997). Evidence for the leading role of OM in the formation of organic and inorganic carbonates has been revealed by experimental work showing that specific organic substances induce calcite and Mg-calcite precipitation, while others cause aragonite to precipitate. Experiments with ooids point to a relationship between the presence and composition of dissolved and adsorbed OM and the formation of specific structures (e.g. radial structures). Studies of mineral-organic matrix relationships of organisms indicate that specifically composed membranes tend to nucleate carbonate and control the mineralogy and ultrastructural arrangement of the biocrystals.
7.1.2 Common Carbonate Minerals A few criteria of common carbonate minerals were summarized in Sect. 3.2.4 and in Fig. 3.8. Fig. 4.9 outlines the mineralogical composition of skeletal grains. Carbonate mineralogy and chemistry are discussed by Reeder et al. (1983) and Morse and Mackenzie (1990). Aragonite: Orthorhombic aragonite differs from calcite in its higher specific gravity and hardness, lack of cleavage, and a high Sr/Ca ratio that varies with water temperatures. In skeletal aragonites, physicochemical processes play a major role in the incorporation of Sr into aragonite during biomineralization. Mg occurs only in low concentrations. Aragonite occurs in modern sediments in the form of skeletal material and cement. Marine aragonite often precipitates as needlelike crystals forming in isopachous and circumgranular rims around grains, as an intergranular mesh of crystals, and as botryoids. In recent carbonates and Cenozoic rocks aragonite is abundant, but is less common in Mesozoic rocks, and almost absent in Paleozoic rocks. The decrease in aragonite with geologic age is in part due to diagenetic processes, but probably also reflects secular mineralogical changes during time (see Sect. 7.1.5.1).
Carbonate Minerals
Aragonitic skeletal grains are susceptible to dissolution causing secondary porosity. The skeleton of totally aragonite fossils is replaced by calcite crystals that are commonly one or more orders of magnitude larger than the aragonite crystals they replace. Aragonitic grains can suffer two main fates: They are either dissolved to form complete or partial molds that are subsequently filled with cement, or they may be transformed to Low-Mg calcite by the simultaneous volume-per-volume dissolution of aragonite and precipitation of calcite along intervening solution films. Skeletal aragonite dissolution is widespread in shallowmarine carbonates and is commonly regarded largely as a near-surface freshwater process (James and Choquette 1983). Some evidence exists however, for skeletal aragonite dissolution on shallow sea floors (Palmer et al. 1988) and in hypersaline waters (Sun 1992). In deeper waters of the open oceans aragonite is dissolved at the aragonite lysocline and finally at the aragonite compensation depth (Box 2.6; Sect. 2.4.5.6). Mg-calcite (High-Mg calcite, HMC): Many biogenic calcites and non-biogenic calcites contain variable amounts of MgCO3. These calcites with more than a few mol% MgCO3 are generally referred to as magnesian (or magnesium) calcites. The boundary between Low-Mg calcite and High-Mg calcite is a matter of discussion (Richter 1979, 1984) as is the distinction between High-Mg calcite and ‘high High-Mg calcite’ (Sect. 3.2.4). Chave (1954, 1964) used 4 mol% as a boundary. Many Mg-calcites have an average composition of about 14 mol% MgCO3 (Bathurst 1975) but biogenic Mg-calcite varies strongly in the amount of MgCO3 contents (e.g. calcareous red algae 10–30%, echinoderms 10–15%). In modern environments magnesian calcite occurs in invertebrate and algal skeletons, cements and ooids; in non-marine environments in freshwater tufas and cements formed in terrestrial and aquatic settings. Modern marine inorganic Mg-calcite cements form predominantly in warm low-latitude waters that are supersaturated with respect to calcium carbonate. The Mg/Ca ratio of magnesian calcites varies directly with the Mg2+/Ca2+ ratio and the temperature of the ambient water. Temperature and latitudinal controls are evident for most shallow-marine organisms forming Mg-calcite skeletons (calcareous red algae, benthic foraminifera, bryozoans, echinoderms and barnacles; see Fig. 4.9). A marked compositional change occurs from the tropics to the poles. Higher latitude calcites formed in cold waters (as well as cements formed in deeper waters of the oceans) contain less magnesium than lower latitude calcites in tropical warm seawater,
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Carbonate Minerals
where MgCO3 mol% increases with increasing temperature (Videtich 1985; Morse and Mackenzie 1990). High-Mg calcite is rarely preserved in the geological record. Exposure of Mg-calcite to meteoric water leads to dissolution and eventual replacement by stable Low-Mg calcite, processes which may occur within geologically short time ranges of only a few 100 000s of years (Friedman 1964; Land et al. 1967; Richter 1979). Both incongruent dissolution in the stabilization of HMC skeletons without evidence of textural modification of skeletal microstructures and congruent dissolution of HMC and the precipitation of LMC with defined textural changes have been reported. The process of Mg loss during stabilization of Mg-rich calcites is not well understood (Oti and Müller 1985; Bischoff et al. 1993). Calcite (Low-Mg calcite, LMC): Low-Mg calcite is traditionally regarded as a stable carbonate phase with little tendency toward intracrystalline alteration of its chemistry, but dissolution-cementation processes may occur and produce intracrystalline porosity by partial dissolution (Pedone et al. 1994). Low Mg/Ca ratios of meteoric waters favor the precipitation of LMC in meteoric and burial diagenetic environments. In marine environments LMC carbonates form at platform margins, on slopes and on the deep-sea floor by primary LMC cements or rapid conversion of High-Mg calcite to calcite (Schlager and James 1978). LMC is a major constituent of pelagic chalks, limestones and marbles. In deeper waters of the open oceans calcite dissolution takes place at the calcite lysocline and the calcite compensation depth (Sect. 2.4.5.6 and Box 2.6 ). Calcites may contain trace amounts of Mn2+ and 2+ Fe . Calcite containing ferrous iron Fe2+ (up to several thousand ppm) is called ferroan calcite. Ferroan and non-ferroan calcites (as well as ferroan dolomite: ankerite) can be distinguished by staining with potassium ferricyanide (Dickson 1966; Adams and MacKenzie 1998) and their luminescence and intensity. Ferroan calcite occurs as cement and as fissure filling, replacement of calcitic matrix or replacement of aragonitic or Mg-calcitic skeletons. Transformation of aragonite to ferroan calcites requires the existence of an open system, e.g. late-stage void filling by sparry calcite in former aragonite molds. Ferroan calcite cements are often of late diagenetic origin and can indicate deep burial reducing conditions, because shallow settings of carbonate sequences favor oxidization, whereas ferroan calcite cements precipitates under reducing conditions in the pH range 7 to 8.
Dolomite: The mineral dolomite, CaMg(CO3)2, can be distinguished from calcite by its commonly euhedral rhombic crystal form in thin sections, X-ray diffraction, and staining techniques using titan yellow alizarine S, or potassium ferricyanide. Small amounts of iron produce brown or tan weathering colors of many dolomite and dedolomitized rocks.
7.1.3 Diagenetic Processes and Controls Major diagenetic processes affecting carbonate sediments and rocks are micritization (Sect. 4.2.3), dissolution and cementation, compaction, neomorphism, dolomitization, and the replacement of carbonate grains and matrix by non-carbonate mineralogies (e.g. silicification and chertification; see Sect. 13.1.2.1). Dissolution: Undersaturation of pore fluids with respect to carbonate leads to dissolution of metastable carbonate grains and cements. Dissolution is particularly effective in shallow near-surface meteoric environments, in deep burial and cold waters (Steinsund and Hald 1994) as well in the deep sea (Berelson et al. 1994), where seawater becomes undersaturated with respect to aragonite and Mg-calcite (Morse 2002). How to estimate dissolution effects? Use fossils as reference base. Many fossils (e.g. mollusks, corals, calcareous algae) have well-defined skeleton mineralogies (Box 4.9) and microstructures. Use the alterations in these criteria to reveal dissolution processes (Sect. 4.2.1; Fig. 4.10). Cementation comprises processes leading to the precipitation of minerals in primary or secondary pores and requires the supersaturation of pore fluids with respect to the mineral. Compaction and pressure solution (stylolitization) refer to mechanical and chemical processes, triggered by the increasing overburden of sediments during burial and increasing temperature and pressure conditions. Neomorphism (Folk 1965) is a term summarizing all transformations taking place in the presence of water through dissolution-reprecipitation processes between one mineral and itself or a polymorph. Recrystallization refers to changes in crystal size, crystal shape and crystal lattice orientation without changes in mineralogy. The term is often used in a very loose manner (Sect. 7.6). Examples for inversion are the gradual replacement of aragonite by calcite through solution and in-situ precipitation in an aqueous environment or the replacement of dolomite by calcite (see Sect. 7.8.3). Cementation as well as neomorphism produce calcite spar fabrics. For diagnostic criteria see Sect. 7.7.
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Dolomitization is a process whereby limestone or its precursor sediment is completely or partly converted to dolomite by the replacement of the original CaCO3 by magnesium carbonate, through the action of Mgbearing water (Sect. 7.8). Major controls on carbonate diagenesis are mineralogy and crystal chemistry, the chemistry of pore waters, water movement, dissolution and precipitation rates, grain size, and the interaction with organic substances. The extent and path of diagenetic reactions are determined by the thermodynamic stability of carbonate minerals being dissolved or precipitated (MacInnys and Brantley 1992), the saturation state of the diagenetic fluid and the available surface area for reaction. Dissolution experiments show that differences in the surface area (corresponding to the microstructure of sedimentary grains) appear to be more significant in controlling relative dissolution rates than mineralogic stability (Walter 1985). The effect of the parameters summarized above depends on the saturation stage and flow rates of the diagenetic fluids (Gonzalez et al. 1992; Lighty 1985).
7.1.4 From Soft Sediments to Hard Rocks To transform soft and friable carbonate muds and sands into micritic limestones and carbonate sands into grainy limestones pore spaces must be filled with carbonate cement. The occlusion of primary depositional porosities up to 60% and 80%, respectively, requires tremendous amounts of carbonate material. But where does all this material come from and where does lithification take place? For a certain time, extensive lithification was thought to be limited to near-surface and subaerial settings (Friedman 1964; Land et al. 1967; see James and Choquette 1990 for a review). The comparison of recent and Pleistocene carbonates seems to indicate that subaerial and near-surface lithification is of overwhelming importance and that most shallow-marine carbonates were lithified as a result of dissolution-precipitation effects. At the same time submarine lithification on or close to the sea bottom was demonstrated for shallow-marine environments (tidal flats, reefs) and deep-water settings (Fischer and Garrison 1967; Ginsburg 1971; see James and Choquette 1983 for overviews). The study of burial lithification started in the late 70s and in the 80s of the last century along with the development of new geochemical techniques and the investigation of pelagic chalks (Garrison 1981; Scholle and Halley 1985; see Choquette and James 1990 and Munnecke 1997 for reviews). The
Secular Variations
transformation of deep-marine calcareous ooze to chalk and subsequently to limestones is associated with a loss of porosity with increasing age and depth of burial (Schlanger and Douglas 1974). The source of the calcium carbonate necessary for burial lithification is seen in the dissolution and reprecipitation of CaCO3 by pressure solution, or dissolution promoted by increasing temperature with increasing depths and overburden, or dissolution/ precipitation of subsurface carbonate rocks by freshwater moving down an aquifer into the basin.
7.1.5 Oscillating Trends in Phanerozoic Carbonate Mineralogy Dominating mineralogies of marine non-skeletal and skeletal carbonate constituents have varied during earth history. Since the primary mineralogy governs early diagenetic processes, these long-term variations must be considered in interpreting carbonate cements, dissolution patterns and porosity development.
7.1.5.1 Secular Variations Based on ooid and cement composition, Sandberg (1975, 1983, 1985) presented strong evidence for secular variations in Phanerozoic non-skeletal carbonate mineralogy. The concept was discussed and underwent various modifications (Mackenzie and Pigot 1981; Wilkinson et al. 1985; Wilkinson and Given 1986; Burton and Walter 1987; Opdyke and Wilkinson 1990; Bates and Brand 1990), but most scientists agree on the inferred existence of calcitic seas typified by the dominance of calcite mineralogies during greenhouse times and aragonitic seas typified by aragonite and High-Mg calcite mineralogies during icehouse times. The Sandberg curve divides the past 600 million years into two greenhouse and three icehouse eras (Fig. 7.1A).
7.1.5.2 How to Recognize Former Aragonite and Mg-Calcite Mineralogy in Ancient LowCalcite Limestones? Estimating the starting mineralogy of matrix, carbonate cements, and grains of limestones assists in understanding early diagenetic processes and porosity evolution. Recognizing former aragonite in skeletal grains: The diagenetic behavior of aragonite shows a surprising uniformity among skeletons of diverse taxa. Re-
Secular Variations
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Fig. 7.1. Secular variations in the mineralogical composition of non-skeletal carbonates (A) and skeletal grains (B, C), and the frequency of reef cements over time. A: The Sandberg model (Sandberg 1983, 1985), based on the study of ancient ooids and carbonate cements, postulates changes between times dominated by aragonite and High-Mg calcite (Late Precambrian to Cambrian, Late Mississippian to Triassic, Mid-Cenozoic to Recent) and times dominated by calcite (Cambrian to Early Mississippian, Middle Jurassic to Early Cretaceous). Note that the prevailing mineralogical composition of non-skeletal carbonates of several time intervals is still a matter of discussion. B: The Stanley-Hardie model (Stanley and Hardie 1998) describes secular oscillations in the carbonate mineralogy of hypercalcifying reef-building and sediment-producing organisms. The model relates Phanerozoic skeletal mineralogy to a chain of causes that extended from mid-ocean ridge processes, via seawater chemistry to the mineralogical and biological composition of reef communities and bioclastic carbonate deposits. Note slight differences in the temporal distribution of aragonite- and calcite-dominated times as compared with the Sandberg model. C: The Hardie model (Hardie 1996) explains the oscillations in the prevailing mineralogical composition of non-skeletal carbonates and synchronous oscillations in the mineralogy of marine potash evaporites by variations of the Mg/Ca ratio and the Ca concentration in seawater. The boundary between the nucleation fields of Low-Mg calcite and aragonite + High-Mg calcite is shown as a horizontal line at Mg/Ca = 2. D: The amount of reef cement and its importance in reefbuilding vary through time (Kiessling 2002). The plot is based on the PaleoReef database (Kiessling and Flügel 2002) which considers low, medium, or high amounts of reef cements. In the plot ordinal values have been transformed into approximate scale values. Vertical bars indicate 95% confidence intervals. Times characterized by cement-rich reefs correspond roughly to phases of biocementstone reefs (Webb 1996). The abundance of reef cements is apparently not related to specific cement mineralogies. Note the conspicuous difference in the medium to high amounts of marine cement in Paleozoic and Triassic reefs as compared to the usually low mean cement content in Jurassic to Cenozoic reefs. Chronostratigraphic time scale for all plots after Golonka and Kiessling (2002).
274
placement of aragonite fossils and microbial fabrics by calcite produces identical neomorphic structures (Chafetz et al. 1993). Aragonitic skeletons (e.g. mollusks, corals, green algae) have a consistent alteration behavior (Sandberg 1984). The replacement crystals are one or more orders of latitude larger than the aragonite crystals they replace. The crystals generally show pitted surfaces, usually fluid inclusions, and occasionally aragonite relicts on etched sections. Elevated Sr contents may be present. Recognizing former aragonite in cements: Aragonite relicts can be verified by X-ray diffraction, elevated Sr (EDX), electron diffraction pattern (TEM) and similarities with known aragonite relicts based on morphology, etching behavior and brightness in SEM. Diagnostic criteria, arranged in a hierarchy of decreasing reliability (Sandberg 1983), are (1) a relative coarse calcite mosaic crosscutting original structures with inclusions of original aragonite; (2) calcite mosaics as for 1, no aragonite relicts, but elevated amounts of Sr; (3) calcite mosaics as for 1, and (4) molds or subsequently filled molds. Recognizing former Mg-calcite in fossils (Lohmann and Meyers 1977; Leutloff and Meyers 1984; Yoo and Lee 1993; Dickson 1995): Indications of former HighMg calcite in skeletons are the replacement of skeletons by ferroan calcite with preservation of the original structures, inclusions of microdolomite, low Sr/Mg ratios, and fossils that have modern analogs known to consist of Mg-calcite.
Diagenetic Environments
Clues to the lime-mud precursor mineralogy of finegrained limestones: Paleozoic and Mesozoic microcrystalline limestones show differences with regard to common crystal sizes (Honjo 1969; Boss and Wilkinson 1991). The matrix of many Paleozoic fine-grained limestones consists of microspar in contrast to Mesozoic limestones with a micrite-sized matrix (Sect. 4.1). This difference has been explained by differences in the original mineralogy of carbonate muds. The dominant mineralogy of mud precursors of micrites has a strong influence on the resulting microfabric (Lasemi and Sandberg 1993). A distinction between ADP (AragoniteDominated Precursors) and CDP (Calcite-DominatedPrecursors) micrites is therefore necessary. ADP micrites are characterized by aragonite relicts, pitted crystal surfaces, common microspar fabrics and high amounts of strontium. CDP micrites are inferred by the absence of aragonite relicts, non-pitted crystal surfaces, the dominance of very small crystals (< 4 µm) and low amounts of strontium.
7.2 Major Diagenetic Environments Diagenetic environments are surface or subsurface zones affected by specific diagenetic processes and typified by assemblages of characteristic criteria that can be studied in thin sections and by geochemical methods. The concept was developed by Matthews (1967, 1971), Purdy (1968), and Land (1970). Diagnostic criteria are cement types and cement fabrics (Sect. 7.4.3).
Box 7.1. Selected case studies on processes occurring in different diagenetic environments. Meteoric environment: Buchbinder and Friedman 1980; Budd and Land 1990; Budd and Vacher 1991; Budd et al. 1993; Buddemeyer and Holladay 1977; Burns and Swart 1992; Dravis 1996; Dunham 1969a, 1969b, 1971, 1976, 1979; Evans and Ginsburg 1987; Friedman and Kolesar 1971; Humphrey et al. 1986; James and Choquette 1990; Keupp 1978; Koch and Rothe 1985; Land 1970; Longman 1980; Matthews 1968; Quinn 1991; Reeckmann and Gill 1981; Richter 1979; Schroeder 1973; Steinen 1979; Steinen and Matthews 1973; Thorstenson et al. 1972; Vollbrecht 1990 Beachrock and marine vadose zone: Alexandersson 1973; Amieux et al. 1989; Beier 1985; Binkley et al. 1980; Cooper 1991; David and Kinsey 1973; Gischler and Lomando 1997; Gray and Adams 1995; Hollail and Rashed 1992; Kneale and Viles 2000; Meyers 1987; Neumeier 1999; Pigott and Trumbley 1985; Richter 1979; Schillings 1994; Scoffin and Stoddart 1983 Mixing zone and freshwater lenses: Back et al. 1986; Chafetz et al. 1988; Kimbell and Humphrey 1994; Richter et al. 1990; Steinen et al. 1978; Stoessel et al. 1989; Stoessel and Ward 1990 Shallow marine environment: Burns and Swart 1992; Evans and Ginsburg 1987; Ginsburg 1957; Ginsburg and James 1976; Ginsburg and Schroeder 1973; Longman 1980; Mullins et al. 1985; Vollbrecht 1990; Vollbrecht and Meischner 1996 Deep marine environment: Adelseck 1978; Archer 1991; Backerand Mill 1986; Berger 1975; 1978; Broecker and Takahashi 1978; Cook and Egbert 1983; Lingen 1975; Milliman 1971; Schlager and James 1978; Schlanger and Douglas 1974; Scholle 1977 Burial environment: Halley 1984; Hemleben and Auras 1984; Hobert and Wetzel 1989; Hutcheon 1989; James and Choquette 1990; Wanless 1983
275
Diagenetic Environments
A Diagenetic environment
Location
Pore Filling
Processes
~ Time needed
Meteoric vadose environment
Above water table, Pores filled with between land surface freshwater and/ and meteoric or air phreatic zone
Solution zone (soil): Extensive solution; removal of aragonite; formation of vugs. Precipitation zone (near surface): Minor cementation
103 - 105 years
Meteoric phreatic environment
Below water table, Pores filled may tend downwards with freshwater 100s of meters
Solution zone (e.g. sinkholes, caves): Solution; formation of molds and/or vugs. Active zone (upper part of meteoric phreatic environment): Dissolution of aragonite and Mg-calcite; rapid and diverse cementation; precipitation of calcite; creation of molds and vugs. Stagnant zone (deeper part and in arid climates): Little cementation; stabilization of aragonite and Mg-calcite
103 -105 up to 106 - 107 years
Marine phreatic environment
On the shallow or deep sea floor or just below
Pores filled with marine water
Shallow-marine environment: Waters oversaturated with 101 - 104 respect to CaCO3; rapid cementation by aragonite and years Mg-calcite; diverse cement types. Deep-marine and cold-water environments: Waters undersaturated with respect to CaCO3; strong dissolution of aragonite and calcite at two dissolution levels
Burial environment
Subsurface beneath reach of surfacerelated processes, down to realm of lowgrade metamorphism. May tend downwards 1000s of meters
Pores filled with brines of varying salinity, from brackish to highly saline
Shallow burial (first few meters fo tens of meters) and deeper burial (sediment overburden of hundreds to thousands of meters): Physical compaction; chemical compaction (pressure solution); cementation; porosity reduction
B
106 -108 years
Fig. 7.2. Major diagenetic environments. A – Simplified scheme. Many of the studies dealing with the diagenesis of carbonate rocks are environment-specific and concentrate on processes affecting particular hydrogeochemically defined diagenetic environments. Carbonate diagenesis operates in the meteoric environment, the marine environment and the burial environment. In the meteoric environment pore space is occupied by freshwater and air (meteoric vadose zone above water table; hatched) or by freshwater (meteoric phreatic zone). The marine-vadose zone at the land-sea boundary and the mixing zone in coastal areas and shallow near-coastal subsurface exhibits meteoric and marine criteria. In the marine phreatic environment water is supersaturated with respect to CaCO3 in shallow seas and undersaturated in cold and deep seas. The subsurface burial environment comprises the subsurface beneath the reach of surface-related processes down to the realm of low-grade metamorphism. Conventionally shallow burial and deep burial are differentiated. The term near-surface diagenesis refers to processes at or close to the sea floor and in the meteoric environment within the reach of surface-related processes related to depositional or weathering interfaces. Here, cementation is highly facies-specific. The terms eogenic, mesogenic and telegenic, introduced by Choquette and Pray (1970), refer to early near-surface, burial and uplift/unconformity-related processes (Sect. 7.3.2). B – Major processes occurring in different diagenetic environments. The time involved in diagenetic processes varies significantly in different diagenetic zones. Early diagenetic solution/precipitation processes in meteoric vadose and shallow marine phreatic environments need far less time than late diagenetic deeper burial diagenesis, which can last millions of years. Similarly, unconformity-related meteoric phreatic processes may continue over very long time intervals. Early cementation in intertidal and shallow subtidal environments occurs within a range of almost recent to several tens to a few thousand years. Synsedimentary botryoidal cements on marginal slopes of platforms may grow over several tens of years, resulting in synsedimentary stabilization of steep carbonate slope deposits at or above angles of repose (Grammer et al. 1993).
276
Lucid synopses of the most common diagenetic environments were published by Matthews (1974), Longman (1980) and Harris et al. (1985). Box 7.1 lists some key papers. 7.2.1 Meteoric, Marine and Burial Diagenesis The location of the major diagenetic environments is shown in Fig. 7.2, the processes that control these environments in Fig. 7.3. Diagenetic zones pass vertically and laterally into others. Many carbonate rocks have changed from one diagenetic zone to another during their history because of sea-level fluctuations, tectonic movements, and increasing or decreasing burial.
7.2.1.1 Meteoric (Freshwater) Diagenesis Meteoric is a Greek word meaning freshwater derived from the Earth’s atmosphere. Freshwater diagenesis takes place in continental areas, along shelf margins, upon platforms with islands, and on atolls or isolated platforms where sediments rise above sea level (Fig. 7.3). Much of our knowledge of meteoric carbonate diagenesis is based on studies of exposed Pleistocene limestones (e.g. Bermudas, Barbados, Jamaica, Red Sea; Land 1986; Vollbrecht 1990). Meteoric diagenesis starts with the loss of magnesium from High-Mg calcite followed by the gradual disappearance of aragonite and the replacement of aragonite by calcite. Freshwater diagenesis occurs in the meteoric vadose zone and the meteoric phreatic zone as well as the shal-
Fig. 7.3. Relationships of major diagenetic environments in the shallow subsurface of an ideal permeable carbonate sand island. Differences in pore space fluids and in the possibility for circulation of pore waters cause different dissolution and precipitation processes, which are recorded by specific types and assemblages of carbonate cements. The subaerial sediments of the meteoric vadose zone are affected by rain water. Sediments in the meteoric phreatic zone are subject to active water circulation (e.g. inner ramps, outer platforms, windward parts of reefs) or stagnant conditions (e.g. lagoons, deeper ramps). Modified from Longman (1980).
Mixing Zone
low subsurface, characterized by the mixing of freshwater and marine water. In the vadose zone water is concentrated by capillarity at grain contacts; in the phreatic meteoric zones water fills all pores. Both zones comprise subenvironments distinguished by differences in water movement (active circulation or stagnant conditions), solution and precipitation processes and resulting cements and porosity types (Longman 1980). Meteoric freshwater lenses: In coastal and island settings meteoric freshwater lenses extend down hundreds of meters (Fig. 7.3). Freshwater phreatic lenses are caused by underground streams entering the nearcoast subsurface area and floating on saltwater of marine origin beneath or bounded at the base by impermeable layers (Matthews 1968). Recent island freshwater lenses show different zones of solution, stagnation, and active flow (Longman 1980; Budd 1988; Vacher 1988). Examples are known from Florida, the Bahamas, Barbados and Bermuda (Land 1970; Matthews 1968; Halley and Harris 1979). The diagenetic pattern of freshwater lenses (Saller 1984; Budd 1991) is explained by the hydrological Ghyben-Herzberg model, which refers to freshwater lenses floating on seawater on a homogeneous porous island or cay aquifer. The model suggests that the freshwater lens extends below mean sea level approximately 40 times the height of the water table above mean sea-level. Small drops in sea-level, therefore, may cause relatively deeply extending lenses recorded by specific cement associations.
7.2.1.2 Mixing Zone and Marine Vadose Environment The mixing of waters with dissimilar chemistries has the potential for geochemical activity. Specific diagenetic conditions characterized by the mixing of meteoric and marine waters may exist in shallow subsurface near-coastal settings (mixing zone) and at the coastal interface of land and sea (marine beaches, shoreface), characterized by rapidly changing conditions (marine vadose environment). The mixing of waters in the mixing zone may be partly responsible for the precipitation of aragonite and dolomite cementation in vugs (Kimbell and Humphrey 1994), dolomitization (Badiozamani 1973; Cander 1994; see Sect. 7.8.2.2) and chertification. Mixtures of salt water and fresh water contribute significantly to the dissolution of carbonate rocks on coasts (Sanford and Konikow 1989). The diagenetic criteria of the marine vadose zone are similar to those of the meteoric vadose zone (pendant and meniscus cements), but the cementation is
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Burial Diagenesis
Box 7.2. Factors and conditions controlling the burial diagenesis of carbonates. Factors depending on the composition of the sediment 1 Mineralogy: The diagenetic potential of shallow-marine carbonates consisting of aragonite and Mg-Calcite is lower than that of deeper water carbonates consisting of Low-Mg calcite. Pelagic mud of chalks consisting of Low-calcitic coccoliths generally exhibits few pressure solution criteria and little interparticle porosity, and therefore good preservation of primary porosity. Pressure solution is accelerated by Mg-poor fluids. Carbonates with relatively low amounts of clay minerals are more susceptible for compaction and pressure solution than carbonates with higher amounts of clays (10% and higher). 2 Grain size and texture: Fine-grained poorly sorted sediments are more rapidly dissolved and compacted because of their larger grain surfaces and higher amounts of pore water. 3 Porosity and permeability: Control diagenetic reactions, dependent on pore-water amount and flux rates. 4 Presence of early cement or early dolomite: May increase rigidity and reduce or hinder stylolitization. Conditions 1 Pressure: Lithostatic, hydrostatic and tectonic pressure. Overpressed (geopressed) sediments may be formed in areas of rapid sedimentation, in sequences where porous carbonates are bordered by permeability barriers (e.g. hardgrounds, lateral stylolites, argillaceous sediments, evaporites). Pore fluid pressure may weaken hydrostatic pressure. 2 Temperature: The solubility of calcite appears to decrease with increasing temperature. The increase in the overburden favors temperature- and pressure-controlled mineral reactions (gypsum –> anhydrite, opal C –> opal A –> quartz, smectite –> mixed-layer minerals –> illite). Increasing temperature controls the transformation of stable organic matter into unstable organic matter and hydrocarbons. Free CO2 resulting from these processes and organic acids contribute to the dissolution of carbonate minerals and the creation of porosity. 3 Pore water chemistry: Influences on transport and the speed of pressure solution, which is more rapid in Mg-rich pore waters of marine environments than in Mg-poor burial pore waters.
more rapid. Typical rocks originating in this setting are beachrocks (see Sect. 2.4.2.1).
7.2.1.3 Marine Diagenesis Marine phreatic diagenesis takes place at the shallow or deep sea floor or just below, and on tidal flats and beaches. High hydrostatic pressure, low water temperature, and high partial pressure of CO2 in the deep sea lead to dissolution and to significant differences in the calcium-carbonate preservation of deep-sea sediments. These differences are used to distinguish two dissolution levels, the lysocline and the calcite compensation depth (see Sect. 2.4.5.6).
7.2.1.4 Burial Diagenesis Carbonates that have undergone burial diagenesis are ‘out of sight’ and somehow ‘out of mind’ (Scholle and Halley 1985). Microfacies criteria indicating burial diagenesis are specific cements (Sect. 7.4.2.1) and diagenetic textures pointing to compaction and pressure solution (Sect. 7.5). The study of burial diagenesis is highly significant in the reconstruction of rock history and the evaluation of rock properties and porosity. Reviews on burial diagenesis have been published by Wanless (1983) and Hutcheon (1989).
Burial environments are conventionally subdivided into shallow burial and deep burial, but the boundary is by no means well defined. The shallow burial zone includes the first few meters to tens of meters of burial. Shallow burial near-surface diagenesis are influenced by changing pore water chemistry in the mixing zone, temperature and pressure processes. Diagenetic processes in the deep burial zone are controlled by conditions summarized below and in Box 7.2. Pore water composition in deep burial environments differs substantially from that of shallow diagenetic environments, where cementation usually takes place in dilute meteoric waters that are oxic or only slightly reducing. Deep-burial subsurface pore waters have elevated salinities and act in a reducing manner. As a consequence redox-sensitive elements are mobilized and Mn and Fe tend to be incorporated in the calcite or dolomite cements. The cements are characterized by distinctive morphologies (Sect. 7.4.2). Changing Mn and Fe compositions of subsurface fluids result in mineral zoning recorded by CL patterns, specific fluid inclusion salinities (Dorobek 1989), and stable isotope signals (Scoffin 1987). Processes in deeper burial environments include: • Physical compaction due to sediment overburden: Reduces thickness of sediments, porosity and permeability, leads to breakage and distortion of grains
278
Porosity of Carbonate Rocks
(readily seen in grainstones), and produces compressed fabrics (Sect. 7.5.1). • Chemical compaction: Starts at various depths of overburden (commonly several hundreds and thousands of meters, but also at an overburden of 100 or 200 m). Reduces thickness of sediments, porosity, and permeability. Produces stylolites and other pressure-solution structures. Pressure solution provides carbonate for burial cementation (Sect. 7.5.2). • Cementation: Produces coarse calcite spar cements. The cements are enriched in iron and manganese, poor in strontium. Fluid inclusions are common. Coarse or poikilotopic calcite, drusy and other calcite mosaics, saddle dolomite. • Minor solution porosity: Caused by dissolution of carbonate and calcium sulphate minerals. • Burial dolomitization: Anhedral crystalline fabric; generally coarse crystals.
7.2.2 Early and Late Diagenesis Diagenesis continues throughout the life of a sediment or sedimentary rock. No clear boundary can be placed within the continuum of diagenetic alterations that change carbonate sediments to carbonate rocks, but the terms early and late diagenesis are generally used in order to indicate the relative timing of diagenetic processes. The term early diagenesis refers to diagenesis occurring immediately after deposition or immediately after burial (Berner 1980). Reviews of early diagenesis were written by Bathurst (1989) and Dickson (1984). Late diagenesis occurs a long time after deposition, when the sediment is already more or less compacted into a rock and corresponds to burial processes acting in the subsurface over a long geological time. The relative importance of early diagenesis versus late burial diagenesis is controversial (Longman 1980, Scholle and Halley 1985), but most carbonate rocks appear to be more intensively affected by burial than by early diagenetic processes.
7.3 Porosity of Carbonate Rocks Porosity is a prerequisite for the course of diagenesis. Porosity studies of carbonate rocks are crucial in understanding diagenetic processes and highly significant in evaluating reservoir rocks (Wardlaw 1979; Ehrlich et al. 1991; Chilingarian et al. 1992; Moore 2001; see Sect. 17.1.3.1). Thin sections of limestones and dolomites allow diagenetic pore types to be qualitatively and quantita-
Fig. 7.4. Microporosity in a micritic reservoir rock. Porosity is caused by equant to equant-elongate or tubular or platy pores between calcite crystals (intercrystalline porosity). In this sample pores are very small (0.5 to 5 µm, mostly < 3 µm), but effective porosity is high because of well-developed pore throats and good permeability. The rhombohedral crystals are of different size and have well-developed crystal faces, forming an inequigranular idiotopic crystal fabric sensu Friedman (1965). Microporosity occurs in the matrix of lime mudstones, wackestones and packstones as well as in carbonate grains. Processes involved in the creation of microporosity are dissolution, dolomitization and incipient weathering. SEM. Subsurface. Late Cretaceous: Ras Al Kaimah, United Arabian Emirates. Scale is 10 µm.
tively differentiated, and the relationships of thin-section porosity and porosimetry data to be discussed. The following text concentrates on the basic terminology and the practical pore type classification in thin sections.
7.3.1 Porosity Categories, Pore Geometry and Permeability 7.3.1.1 Basic Definitions Porosity is the percentage of the bulk volume of a rock that is occupied by interstices, whether isolated or connected. This definition describes the total porosity which must be separated from the effective porosity. The latter is the percentage of the total rock volume that consists of interconnected pores. Many carbonate rocks no longer exhibit open pores, but former interstices which have been filled with cement. An estimation of this minus-cement porosity, representing the sum
Porosity
of pre-cementation porosity and expressed by the total of pore-filling cement percentage of the bulk volume, is necessary for understanding and describing porosity evolution during time (see Pl. 30/1, 3). Thorough reviews of the creation and destruction of carbonate porosity were published by Jodry (1972), Bebout et al. (1979), Wardlaw (1979), Roehl and Choquette (1985), Moore (1979, 1989, 2001), Purser et al. (1994), Budd et al. (1995), and Lucia (1999). Primary and secondary porosity: The porosity of sedimentary rocks falls into two major groups: • Primary porosity forms during the predepositional stage ( e.g. intragranular pores in foraminifers, corals, or ooids) and during the depositional stage (depositional porosity), e.g. intergranular porosity, framework growth porosity. • Secondary porosity is formed during diagenesis at any time after deposition. The time involved in the formation of secondary porosity may be tremendously long, and can be subdivided into three stages called eogenetic, mesogenetic, and telogenetic by Choquette and Pray (1970) and summarized in Box 7.3. Important processes generating secondary porosity are dissolution, dolomitization/dedolomitization, fracturing
279
and brecciation. Secondary porosity formation by dissolution occurs at any point in the burial history and can substantially enhance reservoir properties. Prime prerequisites for the formation of secondary porosity are the existence of a solution undersaturated with respect to carbonate, and fluids that transport the solutions. Deep subsurface processes resulting in undersaturation are mixing of waters of dissimilar composition, dissolution by acidic waters through decarboxylation, and thermal maturation of organic matter related to igneous intrusions and thermal metamorphism (AlShaleb and Shelton 1981; Surdam et al. 1984). Thermally driven diagenesis of shales associated with dewatering of clays and transformation of smectite to illite may form dilute solutions that lower the saturation state and thus cause fluid flow. Enhanced porosity originates from the enlargement of primary interparticle porosity, and the creation of secondary porosity, predominantly by selective dissolution of carbonate grains (e.g. aragonitic shells) during subaerial exposure, or dissolution of evaporite minerals. Solution porosity (Pl. 15/4, Pl. 34/6, Pl. 36/2, Pl. 123/2) refers to moldic pores, vugs, channels and caverns. Vugs and channels are crosscutting non-fabric
Box 7.3. Porosity and the timing of diagenetic processes (after Choquette and Pray 1970). During their geological history carbonate rocks can be affected several times by mesogenetic and telogenetic processes causing modification, destruction and renewed construction of secondary porosity. Eogenetic (near-surface) Near-surface diagenetic processes of relatively short duration occur in the time between the deposition of sediments and their burial within the zone of active surface-related processes and surface-promoted fluid migration. The upper limit is the subaerial or subaqueous interface, the lower limit lies at the point where surface-related meteoric or marine waters cease to circulate by gravitation or convection. Sediments are mineralogically unstable; their porosity is modified by dissolution, cementation and dolomitization. Diagenetic environments active within the eogenetic zone comprise the meteoric vadose, the meteoric phreatic zone, and the mixing zone. Porosity in the eogenetic zone is medium to high. Mesogenetic (burial) Diagenetic processes taking place during burial, away from the zone of major influence of surface-related processes are characterized by rather slow porosity modifications, but often radical porosity destruction due to compaction and compaction-related processes. The mesogenetic zone corresponds roughly to the deep burial diagenetic environment. Telogenetic The term refers to the time during which long-buried mineralogically stable limestones and dolomites of the mesogenetic zone are exhumed in connection with unconformities (unconformity-related diagenesis), following tectonic uplift and controlled by surface-related fluid migration. The rocks are affected by surface-related meteoric processes. Long-buried and exhumed carbonate rocks are significantly influenced by solution and precipitation processes, often in association with the formation of unconformities as seen in paleokarst systems. Karstification modifies porosity and produces solution pores ranging in size from tiny voids to extremely large caverns (Sect. 2.4.1.3). Because sea-level changes can affect deposits from the surface down to a burial of several hundreds of meters, eogenetic criteria can overprint mesogenetic and telogenetic signatures, as seen, for example, in the karst development in Florida and the Bahamas.
280
Porosity Classification
Box 7.4. Terms used to describe carbonate porosity in thin sections. The arrangement within the list follows the porosity classification of Choquette and Pray (1970), see Fig. 7.5. Fabric-selective pores Interparticle (intergranular) porosity: Porosity between individual particles or grains of a sedimentary rock. Corresponds generally to depositional primary porosity, but also includes secondary porosity (e.g. resulting from partial dissolution of aragonitic ooid cortices). Interparticle porosity is high in modern mud-free carbonate sands (up to about 50%) as well as in modern mud-bearing sediments (about 40 to about 75%). The preservation of open interparticle pores in ancient carbonates is promoted by the absence of water in the pores in dry climates, by a protective seal of clay or evaporites, or by early oil emplacement. Common pore sizes are 0.05 to 1 mm. Pl. 13/6; Pl. 29/5, 6, 7; Pl. 34/1, 2, 4; Pl. 40/5; Pl. 41/1; Pl. 61/1, 2; Pl. 80/2; Pl. 93/5; Pl. 120/2; Pl. 134/7. Intraparticle (intragranular) porosity: Primary pore space corresponding to defined parts of skeletons (intraskeletal porosity, e.g. chambers of foraminifera: Bachman 1984) or to open spaces created by the removal of less calcified internal elements (e.g. central part of Halimeda –> Pl. 7). Pl. 29/2, 7; Pl. 67/1. Depends on the morphology and microstructure of tests and skeletons of organisms as well as on the ultrastructure of the grains (e.g. ooids or aggregate grains). Common pore sizes are <0.01 to 1 mm. Pl. 29/2, 7; Pl. 58/ 2, 5. Growth framework porosity: Primary porosity associated with the growth of reef-building organisms. Framework porosity may be high (in modern coral reefs) or low (in reefs dominated by encrusting organisms). It tends to become quickly reduced by infilling of sediment and by carbonate cements. Pl. 42/1; Pl. 48/7; Pl. 65/3. Fenestral porosity: Primary porosity bound to synsedimentary open-space structures (Sect. 5.1.5), and commonly associated with supratidal and intertidal, algal- and microbial-related, mud-dominated sediments. Pl. 20/1, 2, 7; Pl. 29/9, Pl. 30/5; Pl. 41/2; Pl. 46/2; Pl. 50/5; Pl. 123/2. Shelter porosity: A type of primary porosity created by the shelter and umbrella effect of relatively large grains which prevent the sediment infilling of pore space underneath lying. Shelter porosity is favored by the existence of large platelike fossils (e.g. larger foraminifera, platy algae). Pl. 17/2; Pl. 29/1, 4; Pl. 105/2. Intercrystalline porosity: Porosity between more or less equal-sized crystals, often related to early and late diagenetic recrystallization and dolomitization processes. Common pore sizes are <1 to 10 µm. Pl. 39/1-3. Moldic porosity: Results from the selective removal, commonly by solution, of grains, e.g. fossils or ooids. Requires a distinctive mineralogical or microstructural difference between the solubility of grains and matrix or cements. Molds form preferentially in rocks of mixed mineralogies in meteoric-phreatic, but also in burial settings. Common pore sizes are 0.1 to 10 mm. Pl. 29/3, 8, Pl. 30/6. Biomoldic porosity refers to porosity caused by removal of fossils. Pore sizes <0.5 to several millimeters. Pl. 30/1; Pl. 149/5. Dissolution of (mainly aragonitic) ooids results in the formation of oomoldic porosity, particularly in meteoric vadose and meteoric phreatic environments. Common pore sizes <0.20 to >0.5 mm. Pl. 29/3, 4. Non-fabric selective pores Fracture porosity: Results from the presence of openings produced by the syndepositional, depositional or post-depositional burial breaking of rocks (Sect. 5.3.2). Often caused by brittle fracture of shells as a result of increasing overburden before cementation, folding, faulting, salt solution, or fluid overpressing. Common in brittle homogeneous carbonates, e.g. chalks. Fractures may be healed by late diagenetic calcite. Fracture porosity is the main porosity type in many reservoir rocks. Pl. 25/2, 6. Channel porosity: A system of secondary pores in which the openings are markedly elongate and have developed independently of texture or fabric. Vuggy porosity: Caused by irregularly distributed early and late diagenetic dissolution cutting across grains and/or cement boundaries and creating millimeter- to meter-sized holes that must be studied on different scales (McNamara et al. 1992, Dehghani et al. 1999). Dissolution may start from molds or interparticle pores. Decimeter-sized vugs of different size (globular; vertically elongated; irregularly elongated) and corroded walls that are lined by marcasite may be caused by synsedimentary to early diagenetic biogenic methane exhalation originating from the decay of organic matter (Liu et al. 1988). Pl. 15/4; Pl. 29/1; Pl. 30/1, 3. Cavern porosity: Non-fabric selective porosity characterized by large caverns. The term cavern applies to man-sized or larger openings of channel or vug shapes formed predominantly by karstic solution processes. Fabric-selective or not Breccia porosity: Depositional, solution and karst breccias may yield high porosities. Pl. 26/3; Pl. 114/3. Boring porosity: Micro- and macroborers contribute to the formation of very small to centimeter-sized borings. Microborers are effective in producing microporosity. Pl. 52/7, 8. Burrow porosity: Various organisms create organic burrows in relatively unconsolidated sediment. Pl. 19/2. Shrinkage porosity: Shrinkage of carbonate mud in tidal flats can result in the formation of characteristic pore systems.
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Pore Types
MODIFYING TERMS Size Modifiers (for regularly-shaped pores < cavern size)
Genetic Modifiers Process
Direction of stage
Solution s Cementation c Internal sedimentation i
Enlarged Reduced Filled
Classes
x r f
Megapore
mg
mm 256 large
lmg
small
smg
large
lms
small
sms
32 16
Time of formation Primary P pre-depositional Pp depositional Pd
Mesopore Secondary eogenetic mesogenetic telogenetic
S Se Sm St
Genetic modifiers are combined as follows Process + direction + time Examples: solution enlarged cement-reduced primary internal sediment-filled eogenetic
ms
1/2 1/16 Micropore
mc
Use size prefixes with basic porosity types: mesovug –> msVUG; small mesomold –> smsMO; microinterparticle –> mcBP sx crP ifSe
Abundance Modifiers: e.g. percent porosity: (15) or ratio of porosity types: (1 : 2) or ratio and percent: (1 : 2) (15%)
Fig. 7.5. Pore types and porosity classification (Choquette and Pray, 1970). After Adams and McKenzie (1998). Fabric-selective porosity is controlled by primary or secondary fabrics. Primary depositional fabrics include interparticle, intraparticle, growth framework, fenestral, and shelter pores. Fabric-selective secondary porosity is formed by intercrystalline and moldic pores. Nonselective porosity is independent of depositional fabrics and includes fracture, channel, vug and cavern pores. Breccia, boring, burrow, and shrinkage pores may be fabric-selective or not. These qualitative criteria can be substantiated by the additional use of genetic, size and abundance modifiers. Note that the definition of micro- and mega- (or macro-) pores does not always coincide with the limits used by Choquette and Pray. More generalized definitions of microporosity prefer pore sizes of 1 mm or less (e.g. Pittman 1971). Examples of the use of the Choquette and Pray classification are shown on Pl. 29/1 and Pl. 30/1. See Box 7.4 for definitions of pore types.
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selective pore systems. Solution pores are responsible for the porosity of several of the world’s largest oil fields. Timing and location of porosity evolution Porosity originates in different time intervals and reflects the history of carbonate rocks (Box 7.3).
7.3.1.2 Pore Geometry and Permeability Pore geometry influences the ability of the rock to contain water and produce oil and gas. Reservoir properties are controlled by pore geometry, particularly the mean size of pore throats (smaller spaces or constrictions of minimal cross-section area connecting larger pores), and pore interconnections (Jodry 1972; Bliefnick and Kaldi 1996). Reservoir characteristics depend how pore spaces are arranged and interconnected, which is a function of the size and shape of grains and crystals. These criteria also control reservoir-rock productivity and recovery efficiency and must be included in the differentiation of porosity types (Wardlaw and Cassan 1978; Kopaska-Merkel 1994). Note: Pores account for pore volumes, pore throats for permeability! Permeability determines the ability of a fluid to pass through a porous medium. It is measured as the rate at which fluids pass through a sediment or rock. Permeability is calculated according to Darcy’s Law and commonly expressed in millidarcy. It is largely related to the size and shape of conducting pore spaces, the shape and size of pore throats, and the specific surface area within a pore space (Etris et al. 1988). Pore-throat size is inversely proportional to capillary pressure (Wardlaw and McKellar 1981; Kopaska-Merkel and Friedman 1988). Flow rates decrease with the increase of fine-grained matrix. Pure micrites exhibit the lowest permeability. The micrite content of carbonate rocks, therefore, may reflect original permeability conditions. In turn, early cementation is strongly dependent on the primary carbonate facies types. These relationships are the basis of the porosity classification proposed by Choquette and Pray (1970). Flow rates depend strongly on pore distribution: Interparticle and intercrystalline pores should provide higher permeability than molds or intraparticle pores. Pore geometry, especially the size and the shape of the interconnections between adjacent pores, is a major controlling factor. Permeability decreases during diagenesis because of compaction and cementation. Intergranular cements
Porosity Measurement
cause a reduction in permeability. Pore-bridging cements seem to be more effective in reducing singlephase permeability than pore-lining or pore-filling cements (Panda and Lake 1995).
7.3.1.3 Porosity Measurements and Pore Types in Thin Sections Porosity can be measured by physical methods or in thin sections. The former method determines the volume of gas displaced from the pore spaces of a known bulk volume of rock by capillary pressure. Information on physical porosity measurements can be found in Jennungs (1987). Porosity and pore geometry are determined by mercury capillary-pressure curves (Wardlaw and Taylor 1976, Kopaska-Merkel 1994), studies of three-dimensional resin casts, and thin sections showing the size, shape and distribution of the pores (Rieke et al. 1972; Wardlaw 1976). Resin pore casts reveal a wide variety of carbonate pore geometries (Walker 1978; Beckett and Sellwood 1991). SEM techniques are ideal for examining both permeability and porosity. The combined use of SEM and thin section information is useful in illustrating porosity modifications due to solution and dolomitization effects (Armstrong et al. 1980). Pore throats are generally too small to be examined visually in thin sections. Quantitative determination of porosity from thin sections is commonly obtained by point counting or visual estimation. A useful method for a rapid and accurate computer-assisted determination of porosity in a two-dimensional, digitized image of thin sections is described by White et al. (1998). The relationships of thin-section porosity to mercury-injection porosity and between mercury porosimetry and pore types were discussed by Wardlaw and Li (1988), Etris et al. (1988), McCreesh et al. (1991), and other authors combining Petrographic Image Analyis (Ferm et al. 1993; Ehrlich 1984) with statistical analysis. Differentiating porosity distribution and cement volume in thin sections and polished sections (Halley 1978) is considerably facilitated by staining techniques (Ruzyla and Jezek 1987) and digital imaging (Ruzyla 1986; Frykman 1992; Bowers et al. 1994; Carr et al. 1996; White et al. 1998; Anselmetti et al. 1998). These methods allow open pore space (Pl. 30) as well as pores closed by carbonate cements to be recognized and differentiated.
Porosity Classification
283
7.3.2 Porosity Terminology and Classification More than 50% of the world hydrocarbon reserves occur within pores of limestones and dolomites. The assessment of the type and amount of porosity, therefore, is essential for microfacies studies. Through hydrocarbon exploration several classification schemes of pore space were developed with the aim of the describing porosity and permeability (e.g. Archie 1952; Thomas 1962; Choquette and Pray 1970; Mazzullo et al. 1992; Lucia 1983, 1999). The Archie classification was modified by Roehl (1985). The classification is based on the relationships between pore size and rock-matrix texture observed in well cores and cuttings or on sawed surfaces at low magnifications of 10 to 15 x. The Choquette and Pray classification illustrated and described in Fig. 7.5 has received wide acceptance in the oil and gas industry as well as in academia. The classification distinguishes fabric-selective and nonfabric-selective porosity. The first type is controlled by the depositional or post-depositional fabric, the latter by processes which cut across the fabric of the rock. Fabric-selective porosity includes interparticle (Pl. 29/ 1, 5, 6, 7; Pl. 33/1), intraparticle (Pl. 29/2; Pl. 30/2), intercrystalline (Pl. 39/1, 3), fenestral (Pl. 30/4), growth framework, shelter and moldic porosity (Pl. 29/3, 4, 8; Pl. 30/1, 2, 6). Moldic porosity is formed through the selective solution of aragonite grains, e.g. mollusks (biomolds), or ooids (oomolds). The majority of these porosity types results from early shallow diagenesis, but burial diagenesis can also have an effect (Moore and Druckman 1981). The most common non-fabric selective pore types are vugs (Pl. 29/ 1, 9; Pl. 30/1, 3; Pl. 33/6) and channels. Neither of the two types are conformable in position, shape, or boundaries to particular fabric elements. Vugs are more equidimensional pores; the term channel is used for pores that are markedly elongate or continuous in one or more dimensions. Both pore types are large enough to be visible by the naked eye. Vuggy porosity can be subdivided into connected and disconnected types. Other types of non-fabric-selective porosity are fracture, channel, and cavern porosity. All these types are often associated with the uplift and karstification of carbonate successions. Pores in carbonate rocks are either open or occluded by cements or sediment. The evaluation of cement-filled pores (minus-cement porosity) assists in understanding the course of porosity development.
Fig. 7.6. Ternary porosity plot classification (Kopaska-Merkel and Mann 1993). Carbonate ramp. Late Jurassic (Smackover Formation): Alabama, USA.
Ternary diagrams of carbonate pore types provide information on the shapes and origin of pore systems (Kopaska-Merkel and Mann 1993). The plots are based on quantitative data derived from point counting of thin sections. Using the most common kinds of pores (Fig. 7.5), the position of point clusters on the plots record variations in the diagenetic destruction of the sedimentary rock fabric as well as differences in cementation and dissolution between wells in different areas. Interparticle pores dominate many only slightly altered pore systems of grain-supported carbonates. Moldic porosity dominates pore systems primarily affected by leaching of unstable grains (e.g. aragonitic ooids: Pl. 29/4) and by cementation. Intercrystalline porosity often represents the ultimate result of pervasive and fabric-independent dolomitization. Any three pore types can be used to define the apexes of the plot. Rock/petrophysical classes: Lucia (1999) relates porosity, geological data, and petrophysical properties of carbonates. The author underlines the importance of pore space related to the rock fabrics, which controls porosity, permeability and saturation. He distinguishes two major pore space groups, the interparticle pore space and the vuggy pore space, the latter group subdivided into separate-vug pores and touching-vug pores. The interparticle porosity group includes three rock/ petrophysical classes defined by permeability and water saturation. Quantitative porosity groups used in the Shell Standard Legend for interparticle, intercrystalline, moldic and vuggy porosity are up to 5% - low; up to 10% middle; up to 15% - good; >15% excellent. The quantitative characterization of pore systems is facilitated by digital image analysis (Anselmetti et al. 1998). Text continued on p. 288
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Plate 29
Pore Types
Distinguishing Carbonate Pore Types
The plate displays different pore types, using the Choquette and Pray classification. See Fig. 7.5 and Box 7.4. 1 Primary (fabric-selective) and secondary (non-fabric-selective) porosity. Poorly cemented bryozoan rudstone with different types of bryozoans. Primary pores are pores between particles (interparticle or intergranular pores, IP), and within particles (intraparticle pores, black arrows). Secondary, still open pores appear as vugs (VP) originating from solution enlargement of interparticle pores forming vugs that grade into elongate channel pores. Note that outlines of vuggy pores intersect and destroy grain outlines (white arrows). Primary pores have been partly filled with isopachous equant calcite and calcite blocky cement, and partly reopened by solution. Large idiomorphic calcite crystals (top right) were formed subsequent to equigranular cement. The preservation of open interparticle porosity was favored by the stable mineralogy of the skeletal grains (Low-Mg calcite) and burial solution-enhancement of interparticle pores. The dominance of bryozoans, grain association, and weak primary cementation are common criteria in bryozoan-dominated temperate and coolwater carbonates (see Sect. 16.4.1). Tertiary (Middle Miocene): Albacete near Tobara, Alicante, SE Spain. Use this picture as an exercise! You should find the following answers to the questions summarized on Pl. 30: (A) Primary interparticle pores occluded by cement. Open secondary vuggy and intraskeletal pores. (B) Vuggy pores. (C) Primary reduced porosity. Secondary vuggy solution-enlarged porosity. (D) Primary pores enlarged and enhanced by solution, grading into irregular vugs. (E) Vuggy porosity postdates cementation, no compaction. (F) Open porosity about 20%. 2 Open intraparticle (intraskeletal) porosity (P) within a rugose coral. The oblique section shows septa (S) and horizontal tabulae (T). These skeletal elements consist of recrystallized calcite crystals growing perpendicular to the median line of the original skeletal elements. The coral microstructure is still preserved because of the stability of Low-Mg calcite. Some intraparticle spaces were filled with cements, but most intraskeletal pores are still open. Middle Devonian: Belgium. 3 Fabric-selective oomoldic porosity. Ooid grainstone. The primary aragonitic ooid cortices have been leached (arrows) by vadose meteoric solution. Leaching of cortices can create secondary intraparticle porosity, which is either occluded by cement or leads to oomolds. No evidence for strong compaction. Interparticle pores were filled with thin isopachous cement rims and fine-grained drusy mosaic cement. Early Cretaceous (Purbeck facies, Berriasian): Subsurface, Bavaria, Germany. 4 Fabric-selective oomoldic porosity (OM). Ooid grainstones/packstones are important hydrocarbon reservoirs. Principal effective porosities of carbonate reservoirs in the Four Corners Area of the Colorado Plateau include interparticle, interand intraskeletal (algal blade), oomoldic and intercrystalline dolomite types (Weber et al. 1995). Oomoldic porosity on paleodepositional highs is characterized by high porosity values (average 11.3%, McComas 1963) associated with low permeability (average 3.4 millidarcys). High porosity is caused by meteoric leaching of originally aragonitic ooids. Low permeability is due to the insufficient pore interconnection system because of the scarceness of leached mold walls. Note that some ooids have not been leached. This may indicate differences in the primary mineralogy of the ooids. Middle Pennsylvanian (Paradox Formation): Goosenecks of San Juan, Four Corners Area, Utah, U.S.A. 5 Intercrystalline and interparticle nannofossil porosity. Soft Chalk is a very fine-grained, light-colored and friable carbonate sediment, composed largely of minute Low-Mg calcite skeletons of planktonic coccolithophorids (Cribrosphaerella ehrenbergi (Arkhangelsky); see Pl. 7/3-5). Porosity occurs between fragments of nannofossils and between calcite crystals. Small cement crystals form overgrowth on coccolith elements. Intraparticle pores are partly filled with coarse blocky cement. High-porosity chalks are characterized by well-preserved nannofossils indicating minimum fragmentation and dissolution, in contrast to lowest-porosity chalks containing poorly preserved nannofossils. The hydrocarbon production of many chalk reservoirs (e.g. in the North Sea) is due to the preservation of high primary porosity (caused by a unique combination of pore fluid overpressuring and exclusion of water from pores by migrating oil) and permeability enhancement by post-lithification fracturing (Feazel 1985). SEM. Late Cretaceous (Maastrichtian) chalk: Logster, Denmark. 6 Interparticle porosity (IP). Grains are superficial ooids. Note open pores between the grains. Thin white rims around grains are isopachous meteoric cements. Subrecent grainstone: San Salvador, Bahamas. 7 Open interparticle (IP) porosity between skeletal grains (predominantly corallinacean red algae, CA) and intraskeletal porosity within foraminifera (arrows). Corallinacean algal grains are covered by thin isopachous Mg-calcite cement rims. Pleistocene calcarenite: El Haouira, Cap Bon, northern Tunisia. Crossed nicols. 8 Moldic porosity formed by selective dissolution of ooids and bioclasts. Reservoir rock characteristics of dolomitized oolitic grainstones in the Fore-Sudetic area of western Poland are controlled by vadose dissolution of aragonite and Mgcalcite grains. Late Permian (Zechstein Main Dolomite): Subsurface, Laba, Poland. 9 Keystone vug porosity. Keystone vugs (KV) are mm-sized fenestral pores occurring in tidally influenced carbonate beach sands. They form when tides rise and air or gas bubbles enclosed within the fine-grained sediment try to escape from the sediment. The air bubbles push grains aside, thus creating voids. The grains at the top of the vug form a keystone arch that prevents collapse. Preservation of keystone vugs points to rapid cementation. Keystone vugs are generally larger than birdseyes (see Pl. 20/7, 8). Example: Particles are predominantly peloids and micritized ooids as well as a few aggregate grains. Tidal sand bars. Early Cretaceous (Purbeck facies, Berriasian): Subsurface, Kinsau, Bavaria, Germany.
Plate 29: Pore Types
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Plate 30
Porosity in Thin Sections
Porosity Evaluation in Thin Sections
The impregnation of carbonate rocks with blue-dye stained resin before sectioning makes the evaluation of porosity straightforward and facilitates the recognition of pore distribution. Questions which should be answered in evaluating thin sections with regard to porosity are: A Are the pores completely or partially occluded by cement or sediment or are they still open? B Which of the pore types shown in Fig. 7.5 occur and which types dominate? C Classify the sample according to the Choquette and Pray scheme (Fig. 7.5)! D Do primary depositional pores still exhibit their original shape or are they enlarged and enhanced by solution, thus grading into irregular vugs? E Does porosity postdate cementation and compaction (e.g. fracture porosity)? F Estimate the percentage of open (and/or closed) rock porosity! 1 Bioclastic limestone with terrestrial gastropods (Hydrobia, H). A blue color reveals open pores. The snails were deposited near the coast and covered by sediment. The dark particles (bottom left) are micritized algal stems. The primary aragonitic shells were dissolved and replaced by calcite under meteoric conditions. The interior of the shells was filled with peloids prior to the final deposition of the shells. The peloids are fecal pellets exhibiting different preservation; light peloids are calcitized aragonite pellets. The fine-grained sediment exhibits a distinct geopetal fabric consisting of dark micrite at the bottom and lighter, more coarser sediment at the top. The remaining interparticle voids between the shells are lined by thin isopachous cement rims of bladed calcite crystals and blocky spar. Porosity: Moldic porosity is indicated by white arrows. Some pellets show open intraparticle porosity (top right). Black arrows point to vuggy porosity. Brackish-lagoonal environment. Late Tertiary (Early Miocene): Mainz Basin, Germany. Answers: (A) Most pores filled with sediment and cement, some pores open. (B) Moldic, interparticle, vuggy, rare intercrystalline. Moldic porosity now dominates. (C) Primary depositional interparticle porosity mainly filled and reduced, open secondary moldic porosity. (D) Primary pores with original shape. (E) No compaction. (F) Original porosity about 50%, open porosity < 5%. 2 Detail of the sample shown in –> 1. Gastropod shell (Hydrobia) filled with peloids and calcite cement. Note the difference in crystal size and shape of cements within the shell and in the void at the top of the geopetal structure. Many of the peloids and the complete gastropod shell are dissolved. The shape of the gastropod is preserved because of thin protecting organic coatings. 3 Predominantly vuggy porosity (blue) caused partly by solution-enlargement of interparticle pores. Solution enhancement has changed the original fabric-selective porosity to non-fabric-selective porosity. Some intraskeletal porosity is preserved in the whorls of foraminiferal tests. In terms of the Choquette and Pray classification the open porosity is classified as solution enlarged secondary vuggy porosity characterized by mesopores (see Fig. 7.5). This bioclastic grainstone consists of skeletal grains and some micritic intraclasts. Bioclasts are predominantly miliolid foraminifera and a few bryozoans (B). The chambers of the miliolid foraminifera were partly occluded by cement. Non-miliolid foraminifera have been almost completely replaced by calcite (arrows). Tertiary (Miliolid limestone): Carrière Saint-Pierre-Aigle, Paris Basin, France. 4 Fenestral porosity in a travertine formed by spring waters in shallow ponds (see Pl. 2/1-3). Small open fenestral pores occur within a laminated fabric. Note the different cementation within large and small pores: Larger pores are bordered by large scalenohedral calcite crystals, small pores by small bladed crystals. Some scalenohedral crystals are zoned (black arrow). The micritic layers border larger voids that are incompletely filled with blocky calcite crystals. Fenestral porosity occurs in caliche deposits, freshwater tufa, and travertines, but is most common in carbonate mud- and bindstones of tidal flats. Travertine. Terme di Rapolano, north of Rome, Italy. 5 Fenestral porosity in travertine. Isopachous calcite cements lining open pores in places. The crystals are more or less equidimensional, indicating precipitation in a meteoric phreatic environment. Remains of organic-rich laminae are shown as a brown color. Same sample as –> 4. 6 Moldic porosity (blue). Because of the metastability of aragonite the gastropod shell was dissolved, creating moldic porosity. Same locality as –> 4.
Plate 30: Porosity Evaluation
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7.3.3 Porosity in Limestones and Dolomites Primary and secondary porosity of carbonate rocks are strongly facies-controlled and depend on the depositional patterns and the diagenetic history (e.g. Sun et al. 1992). Porosity relies upon the composition and fabric of rock constituents, the distribution and quantity of cement as well as the abundance and distribution of fractures and stylolites. Porosity in carbonate sediments: The volume of pore spaces in Holocene unconsolidated carbonate sediments may exceed that of the sedimentary particles. The average porosity of calcareous sands and muds lies in the range between 40% and 80% (see Box 7.5). Porosity is largely interparticle, with some intraparticle contribution. Porosity and permeability data show a marked inverse relationship. The correlation of high porosities with the lowest permeabilities reverses the relationships in ancient carbonate rocks, which despite great variations in porosity-permeability interrelationships show generally positive correlation of high porosity with high permeability within the same pore type. The main contrast between Holocene carbonate sediment and ancient limestones is the very high porosity of modern muddy sediment, indicating an enormous pore-volume reduction in the transition from carbonate sediment to carbonate rock. Grain-supported sediments undergo relatively slightly early diagenetic changes in porosity and permeability. Porosity in limestones: Porosities change with increasing age and/or burial depth of the sediment (Scholle and Halley 1985). The decrease in total porosity versus depth is described by the ‘normal curve’ exhibiting a decrease from > 60% at the water/sediment interface of deep-water calcareous muds to about 35% at a depth of 1000 m, about 15% at 2000 m and < 5% at 3000 m. This curve has many deviations, depending on pore water pressure, as exhibited e.g. by chalks (Scholle 1977; Feazel and Schatzinger 1985). Mechanical compaction can be responsible for about one third of porosity in micritic limestones. The total porosity of many limestones is often less than 10% whereby grain-supported limestones often exhibit higher porosities than mud-supported limestones. Common porosities of wackestones formed in inner shelf and lagoonal environments are < 3%. The preservation of primary pores requires that postdepositional diagenesis be limited in its pore-destroying effects, that compaction be kept at a minimum, and that fluctuations between exposure and submergence create a balance between the formation of solution pores
Porosity in Limestones and Dolomites
Box 7.5. Porosity and permeability of modern carbonate sediments based on samples from Florida and the Bahamas (Enos and Sawatsky 1981). Interrelationships between porosity and permeability in recent carbonates is largely controlled by depositional texture, particularly the amount of fines <62 µm. Grainstones: Porosity: range 40–53%, mean 44.5%. Permeability: range 15 800–56 600 md, mean 30 800 md. Mud-free skeletal sands near the shelf break consisting of Halimeda, foraminifera and mollusk grains. Packstones: Porosity: range 45–67%, mean 54.7%. Permeability: range 31.5–9,300 md, mean 1,840 md. Grain-supported sediments containing some mud. Wackestones: Porosity: range 64–78%, mean 68%. Permeability: range 37.6–6,570 md, mean 228 md. Mixture of calcareous mud and coarse-shell grains, low energy areas. Very fine wackestones: Porosity: range 67–73%, mean 70.5%. Permeability: range 0.63–1.37 md, mean 0.87 md. Loosely pelleted muddy sediment of sea grass-free low-energy shelf areas. Supratidal wackestones: Porosity: range 61–66%, mean 63.5%. Permeability: range 617–24,100 md, mean 5,500 md. Large-scale pore network due to desiccation and lamination.
and the destruction of porosity by shallow burial cementation. A preservation of porosity in shallow burial environments is a consequence of minimal burial, reduced burial stress, increased framework rigidity, exclusion of pore water, Low-calcite mineralogy, permeability barriers, and pore resurrection. Porosity in mudsupported limestones affected by a prolonged burial diagenesis is commonly strongly reduced. Porosity in dolomites: Many dolomites have higher average porosities and permeabilities than limestones, because of differences in the size, shape and arrangement of crystals. Porosity tends to increase slightly in the initial stages of dolomitization of limestones, but increases abruptly with higher amounts of dolomite. At this stage, the dolomite is characterized by a sucrosic texture composed of equally-sized rhombohedra with intercrystalline porosity originating by dissolution of associated calcite. Dolomitization is a selective process and fine-grained constituents tend to be preferentially replaced. Common pore types in carbonate rocks: Interparticle, vuggy, intercrystalline, and framework pores are common in limestones. Porosity types in reef limestones are represented by growth framework pores, meteoric and vadose solution vugs, and fracture pores. Based on an evaluation of 38 large oil and gas reservoirs located
289
Cementation
in reef limestones, most of the effective porosity is connected to fracture and solution pores, followed by still preserved framework pores (Flügel 1989). Common pore types in dolomites are intercrystalline pores in fine-grained early diagenetic and replacement dolomites; vuggy pores caused by leaching of evaporitic sulfate minerals in supratidal areas; moldic pores originating from leaching of aragonite and calcite particles subsequent to the matrix dolomitization of intertidal and shallow subtidal dolomites; and interparticle pores due to selective solution of grains.
7.4 Pore-Filling Processes: Cementation Obvious signs of the effect of diagenesis are pore-filling cements as seen in thin sections of carbonate rocks (Fig. 7.7). Cement is a chemical precipitate from solution; it grows in pores, and requires supersaturation of the pore fluid with respect to the cement mineral. Cement is distinct from neomorphic spar, which is an insitu replacement of calcium carbonate in the solid state.
Fig. 7.7. Carbonate cements in Devonian fore-reef limestones. Carbonate cements are essential constituents of ancient reefs. In some reefs cementation is the major factor in the formation of the reef itself and in maintaining steep reef slopes. Reef cements appear to be particularly abundant along the seaward margins of ancient reefs as demonstrated by the cement-rich rudstone shown here deposited on fore-reef flanks. This reef was formed on the top of a submarine volcanic swell (Schwarz 1927, Krebs 1969). Larger clasts are reworked stromatoporoids and tabulate corals. The originally high interparticle porosity (about 40 %) is occluded by several isopachous cement fringes consisting of large radiaxial calcite crystals characterized by undulose extinction. The conspicuous black partings are explained as organic or microbial coatings formed during interruptions of cement growth. Petrographical and geochemical data suggest a precipitation of near-surface shallow burial Mg-calcite cements in semi-closed systems. Major controls were the original mineralogy and fluctuating pore-water salinities and temperature (Mirsal 1978). The red sediment is related to synsedimentary submarine volcanism (Schneider 1977). The depositional and diagenetic sequence embraces: (1) Deposition of reworked reef builders, (2) micritization of the clasts, (3) influx of gray and red internal sediment within large interparticle pores, (4) growth of thin fringes of submarine isopachous fibrous cement, (5) shallow burial and periodic growth of radiaxial cements and of (6) rhombohedral calcite cement (Krebs 1969). These limestones are famous decoration stones. Polished sample. Late Devonian (Frasnian): Wirbelau near Weilburg/Lahn, Germany.
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Features characterizing cements include: • often clear and clean appearance, with well-defined, often (but not always) straight crystal boundaries, • sharp contact between spar and particles, • spar between grains does not penetrate into or cut across grains, • presence of two or more generations of spar, • straight crystal edges and frequent triple junctions with 180° angles (enfacial junctions), • long axes of crystals often normal to grain surfaces, • increasing crystal size away from grain surface. Cements represent an important record of the diagenetic history of carbonate rocks. Studies of thin sections are of paramount importance in recognizing diagenetic environments and determining paragenetic sequences in carbonate rocks. Success in reconstructing the diagenetic history depends heavily on the ability to recognize the primary crystal shapes of the cements and their time succession as reflected by cement stratig-
Marine and Meteoric Cementation
raphy. Because crystal shapes observed under the light microscope are often a product of neomorphism (Sect. 7.6), a combination of conventional light microscopy and cathodoluminescence microscopy is strongly recommended. The two methods permit distinctly different levels of observations. Light microscopy exhibits the optical properties as well as shapes and fabrics of crystals under transmitted or polarized light. Cathodoluminescence microscopy offers a possibility to differentiate between primary growth forms of the cement crystals, cements which have been affected by recrystallization, and other neomorphism processes (Pl. 5). Cathodoluminescence facies, defined by crystal shape, fabric and luminescence, is increasingly being used to interpret diagenetic environments (Lavoie and Bourque 1993). The data derived from petrographical analysis must be also combined with geochemical data (stable isotopes, and trace and minor elements). Oxygen and carbon isotopes have proven extremely valu-
Plate 31 Modern Marine and Meteoric (Freshwater) Cementation: SEM and Thin-Section Criteria Marine pore-filling carbonate cements are precipitated from environments ranging from marine to meteoric to burial settings. Shallow-marine cements consist of aragonite (–> 1, 2) and High-Mg calcite cements (–> 3, 4). Calcite cements ( –> 5, 6) are common in the subaerial and near-surface meteoric as well as in burial environments. Examples are from Fuerteventura, Canary Islands (Spain). Modern beach sands along the arid coasts of the Atlantic island consist of volcanic and carbonate grains. The sands are cemented by High-Mg calcite or aragonite into beachrocks (–> 3-4) forming sheets up to 30 cm in thickness. Cementation takes place in the intertidal zone and along the shores of wind-protected bays. Most skeletal grains are fragments of red algae, mollusks, foraminifers and echinoderms. Precipitation of the cements is caused by the evaporation of sea water in interparticle pores at low tide. Subaerially exposed Pleistocene dune sands exhibit minor meteoric cementation (–> 5-6). 1 Aragonite fibrous cement. The crystals form a meshwork of needles. This type of crystal is not commonly recognized in ancient carbonate rocks. Marine-phreatic zone. SEM. 2 Thin section of –> 1. Fibrous cement rim covers rounded litho- and bioclasts. Note the longer crystals along grain contacts and within narrow intergranular voids (appearing black). 3 Microcrystalline High-Mg calcite cement. Closely spaced rhombohedral crystals, sometimes forming rosette-like aggregates. Note the conspicuous open intercrystalline porosity. Beachrock, formed at the interface of land and sea within the marine-vadose environment. SEM. 4 Thin section of –> 3. Equant crusts of microcrystalline High-Mg calcite cement (black arrows) around carbonate and noncarbonate grains. A part of the open interparticle voids (stippled white) is occluded by automicrite connecting micrograined sedimentary particles (black). Beachrock. Microbial automicrite is common in intertidal solution cavities of recent beachrocks (Neumeier 1999). 5 Dogtooth calcite cement. Scalenohedral and elongated rhombohedral crystals with blunted terminations. Length/width ratios are approximately 3:1. Note that dogtooth cement can no longer be taken as unequivocal evidence of meteoric environments, because it also occurs under marine-phreatic, shallow-burial and hydrothermal conditions (Reinhold 1999). Meteoric horizons in Pleistocene eolian calcarenites. SEM. 6 Thin section of –> 5. Dogtooth cement developed on coralline algal bioclasts (CA) and non-carbonate grains. In the meteoric phreatic zone below the water table, pores are filled with water and large cement crystals can grow on all surfaces. Note different spacing and different length of the crystals. Arrows point to broken crystals indicating dissolution or preparation effects. Open interparticle pores appear black.
Plate 31: Cements in SEM and Thin-Section
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able in discriminating carbonate cements formed in different diagenetic environments. Carbonate cements are also common in clastic sequences. They profoundly influence the quality of hydrocarbon reservoirs. For reviews and case studies on carbonate cementation in sandstones see Morad (1998).
7.4.1 Controls on Carbonate Cementation Common cements are aragonite, High-Mg calcite, LowMg calcite and dolomite. Less common are ankerite, siderite, kaolinite, quartz, anhydrite, gypsum and halite. The growth of carbonate cements is favored by high pH and higher temperatures, quartz or chert cements by low pH and low temperatures. The cementation of carbonates at or near the sediment-water interface requires an enormous input of CaCO3 and an efficient fluid flow mechanism. The source of CaCO3 in the marine realm is seawater, in meteoric and burial environments it is usually a solution of the sediment itself.
Cement Fabrics and Types
Precipitation of cements and dissolution of carbonate sediments is primarily controlled by • the composition of pore-fluids, • the flow rate of water through pores, triggered by energy levels. Water energy (supplying seawater from which carbonate cements can be derived) and the rate of sedimentation are considered as important factors in shallow-marine environments, particularly in reefs, • the primary porosity (size and geometry of pores) and permeability, • the number of ions (oversaturation resulting in precipitation, undersaturation resulting in dissolution) and carbonate supply rates, • the salinity of solutions and the Mg/Ca ratio. If Mg/ Ca ratios are low, calcite can crystallize freely, forming large equant crystals. If Mg/Ca ratios are high, calcite and aragonite crystals are prevented from growing sideways, so that fibrous, acicular, micritic or peloidal cements are common. The mineralogy and Mg content of the cements are predominantly controlled by the Mg2+/Ca2+ ratio of the solution Text continued on p. 295
Plate 32 Modern Marine and Meteoric Cementation: Cement Fabrics and Cement Types Cement fabrics indicate whether the pores were totally or only partially filled by fluids. Cement rims consisting of symmetrical cements around grains (–> 1-4) originate in pores completely filled with marine or meteoric water. In pores containing both water and air, water collects at grain contacts, precipitating asymmetrical meniscus-style cements (–> 5, 6). 1 Marine aragonite fibrous cement and microcrystalline Mg-calcite cement. A uniformly broad rim consisting of fibrous aragonite crystals (AC) growing normal to pyroclastic grains (PG) is covered by a thin rim of microcrystalline Mg-calcite (MC). Aragonite cement was formed in the upper subtidal or lower intertidal zone, High-Mg calcite cement within the intertidal zone. Partially dissolved aragonite crystals indicate the influence of fresh water. SEM. 2 Thin section of –> 1. Isopachous (= crust of equal thickness) fringe of cements. Note the difference in the thickness of aragonite (AC) and Mg-calcite cement (MC) rims. The cement grew within voids of pyroclastic sands. Open parts of the void filled with resin appear white. 3 Marine microcrystalline Mg-calcite cement (MC) and acicular aragonite cement (AC). Due to the evaporation of vadose water in tidal zones, broad Mg-calcite rims were formed around coralline algal grains. A change from vadose to marinephreatic conditions favored the precipitation of aragonite cement. Note the square-ended habit of the aragonite crystals. If this shape is preserved in limestones after neomorphism to calcite, it provides evidence of a former aragonite mineralogy of the cement (Assereto and Folk 1976; Mazzullo and Cys 1977). SEM. 4 Thin section of –> 3. Dark Mg-calcite cement rims (white arrows) and isopachous aragonite cement (black arrows) around coralline red algal clasts characterized by their fine network structure (CA; Pl. 54/6. 9). Coralline red algae consist of High-Mg calcite. The overgrowth by Mg-calcite cement points to substrate control of cement growth. Open interparticle pores appear black. 5 Meteoric-vadose meniscus cement characterized by concave calcite cement bridges (arrows) connecting adjacent grains. The cement was precipitated from water films at grain contacts distributed between the grains in a meniscus-like fashion. The overall effect of meniscus cements is to round off the pore spaces. Meniscus cements originate mostly in vadose environments but can also be formed microbially in other environments (Hillgaertner et al. 2001). Open interparticle pores black. SEM. 6 Thin section of –> 5. Meniscus cement (arrows) between coralline algal bioclasts (CA) and volcanic lithoclasts (VL). Open interparticle pores black. All samples from modern beach sands in Fuerteventura, Canary Islands, Spain.
Plate 32: Marine and Meteoric Cementation
293
294
Cement Types
Acicular: Needle-like crystals, growing normal to the substrate. Crystals elongated parallel to the c-axis, exhibiting straight extinction. Terminations are pointed or chisel-shaped, twinning is common. Width < 10 µm, length about 100 µm and more. Often forming isopachous crusts. Predominantly aragonite, but also Mg-calcite. Marine phreatic. Pl. 31/2, Pl. 34/1. Fibrous: Fibrous crystals, growing normal to the substrate. Crystals show a significant length elongation, usually parallel to the c-axis. Crystal shape is needle-like or columnar (length to width ratio > 6:1, width > 10 µm). Size commonly fine to medium crystalline. Often forming isopachous crusts; common in inter- and intraparticle pores. Aragonite or High-Mg calcite. Mostly marinephreatic, but also meteoric-vadose and marine-vadose (columnar crystal shape). Syn.: Radial fibrous. Pl. 2/4, Pl. 31/1-2, Pl. 32/1-4, Pl. 50/6. Botryoidal: Pore-filling cement made of individual and coalescent mamelons exhibiting discontinuous horizons, e.g. dust lines ranging in size from tens of microns to several centimeters. The cement consists of individual and compound fans, which in turn are composed of elongated euhedral fibers with a characteristic sweeping extinction in cross-polarized light. Aragonite. Usually marine (common in cavities of reefs and steep seaward slopes), but also known from burial environments. Syn.: Spherulitic. Pl. 145/1-3. Radiaxial fibrous: Large, often cloudy and turbid, inclusion-rich calcite crystals with undulose extinction. Size medium to coarse crystalline. Sometimes extending several millimeters in length, usually about 30 to 300 µm. Crystal length/width ratio 1:3 to 1:10. Crystals show a pattern of subcrystal units. Within each subcrystal that diverges away from the substrate an opposing pattern of distally-convergent optic axes occurs, caused by a curvature of cleavage and twin lamellae. Undulose extinction of subcrystals or subcrystal units are used in distinguishing three radiaxial subtypes (see text). Often forming isopachous crusts. Phreatic-marine and burial. Pl. 27/2, Pl. 34/2; Fig. 7.9. Dog tooth: Sharply pointed acute calcite crystals of elongated scalenohedral or rhombohedral form, growing normal and subnormal to the substrate (grain surfaces, atop earlier cements). Crystals are a few tens to a few hundred micrometers long and have acute and sometimes blunted terminations. Often meteoric and shallow-burial but also marine-phreatic and hydrothermal. Syn.: Bladed scalenohedral cement, bladed prismatic calcite cement, dentate cement, scalenohedral calcite cement. Pl. 2/3, Pl. 31/5-6, Pl. 34/8. Bladed: Crystals that are not equidimensional and not fibrous. They correspond to elongate crystals somewhat wider than fibrous crystals (length/width ratio between 1.5:1 to 6:1) and exhibiting broad flattened and pyramid-like terminations. Crystal size up to 10 µm in width and between less than 20 and more than 100 µm in length. Crystals increase in width along their length. Commonly forming thin isopachous fringes on grains. Usually High-Mg calcite but also aragonite. Marine-phreatic (abundant in shallow-marine settings) and marine-vadose. Pl. 33/1, 3, 5, 8; Pl, 34/7. Dripstone: Pendant cement characterized by distinct thickening of cement crusts beneath grains or under the roofs of intergranular and solution voids. The cement forms on droplets beneath grains after the bulk of the mobile water has drained out of the pores, leaving a thicker water film at the lower surface of the grains. Forms typically gravitational, beard-like patterns. Predominantly calcite. Formed below the zone of capillarity and above the water table within the meteoric-vadose zone (often associated with meniscus cement), but also in the meteoric-phreatic and sporadically in marine-vadose diagenetic environments (e.g. inter- and supratidal, and beachrocks: aragonitic dripstone cement). Syn.: Gravitational cement, microstalactitic cement, microstalactitic druse cement, stalactitic cement. Pl. 34/6, Pl. 126/1. Meniscus: Calcite cement precipitated in meniscus style at or near grain-to-grain contacts in pores containing both air and water. Exhibits a curved surface below grains. Resulting intergranular pores have a rounded appearance due to the meniscus effect. Characteristically formed in the meteoric-vadose zone but may also occur in the phreatic-meteoric and the vadose-marine environment (beachrock). Pl. 14/1, Pl. 32/5-6, Pl. 33/4, Pl. 126/1. Drusy: Void-filling and pore-lining cement in intergranular and intraskeletal pores, molds and fractures, characterized by equant to elongated, anhedral to subhedral non-ferroan calcite crystals. Size usually >10 µm. Size increases toward the center of the void. Displays a characteristic fabric (see Fig. 7.12). Near-surface meteoric as well as burial environments. Syn.: Drusy calcite spar mosaic, drusy equant calcite mosaic. Pl. 10/2.
Fig. 7.8. Cement types. Part 1.
295
Cement Types
Granular: Calcite cement consisting of relatively equidimensional pore-filling small crystals. Common in interparticle pores, generally without distinct substrate control. Formed in meteoricvadose, meteoric-phreatic and burial environments. Can also originate from recrystallization of pre-existing cements. Pl. 10/2. Blocky: Calcite cement consisting of medium to coarse-grained crystals without a preferred orientation. Characterized by variously sized crystals (tens of microns to several millimeters), often showing distinct crystal boundaries. Xenotopic and hypidiotopic crystal fabrics common. High-Mg calcite or Low-Mg calcite. Typically in meteoric (meteoric phreatic and vadose) and burial environments; rare in marine hardgrounds and reefs. Precipitated after the dissolution of aragonite cements or grains or as late diagenetic cement filling remaining pore space. Blocky textures can also originate from recrystallization of pre-existing cements. Pl. 20/1, Pl. 28/2, Pl. 34/1. Syntaxial calcite overgrowth cement: Substrate-controlled overgrowth around a host grain made by a single crystal (usually High-Mg calcitic echinoderm fragments). Overgrowth often in crystallographic lattice continuity with the host grain. Echinoderm overgrowth is often zoned. Color differences between the skeletal grain and the overgrowth cement can be conspicuous. Overgrowth cements from near-surface marine, vadose-marine and meteoric-phreatic environments are inclusion-rich and cloudy, in contrast to clear overgrowth from deep burial environments Syn.: Grain overgrowth cement, syntaxial echinoderm cement, syntaxial cement rim, syntaxial overgrowth rim cement. Pl. 31/3-4, Pl. 34/3-4, Pl.144/5; Fig. 7.10. Peloidal microcrystalline cement: Characterized by a peloidal (or pelleted) fabric composed of tiny peloids (size <100 µm) within a microcrystalline calcite matrix. The peloids consist of micrite-sized crystals bearing a radiating halo. Shallow-marine. Common in modern and ancient reefs. Possible interpretations: Chemical and/or microbially induced precipitation (Sect. 4.2.2). Pl. 8/5. Microcrystalline or micrite cement: Micron-sized curved rhombic crystals. Forms thin coatings around grains, lines intraskeletal pores, fills pores completely or constructs bridges between grains (contributing to meniscus cement). Mg-calcite. Micritic cement fringes should be distinguished from micrite envelopes (Sect. 4.2.3). Often associated with peloidal cements. Pl. 31/ 3-4, Pl. 32/1-4, Pl. 33/2.
Fig. 7.8. Cement types. Part 2.
from which precipitation occurs (Folk 1974; Mucci and Morse, 1984), but is also determined by the supply of carbonate ions as shown by the relationships between the degree of supersaturation, crystal mineralogy and morphology (Given and Wilkinson 1985), the mineralogy of the substrates (e.g. aragonite cement overgrowth on aragonite grains and Mg-calcite on Mg-calcite grains). Substrate control of cement substrate is an important factor in reefs and subtidal carbonate sands, where the mineralogy of the carbonate cement tends to reflect the mineralogical composition of the reef skeletons or the bioclastic grains (Friedman et al. 1974), the rate of precipitation, the composition of organic matter within grains and on grains. Organic-rich pore waters hinder carbonate precipitation even from supersaturated waters, the activity of microorganisms mediating carbonate precipitation (Sect. 9.1.1).
• nucleation, controlled by the mineralogy of the substrates (aragonite or calcite). Nucleation and growth rates of cement crystals in turn control cement crystal morphology, • crystal growth dependent on the presence of sufficient Ca ions, and therefore, on porosity/permeability conditions.
Precipitation of carbonate cements may be described as a succession of three stages: • oversaturation of a pore fluid within a pore space,
Criteria used for describing cement types are • shape (needle-like, columnar, microcrystalline, tiny crystals around a micritic center),
•
• •
•
7.4.2 Morphology and Fabrics of Cement Types Description and differentiation of cement types are based on the morphology and arrangement of cement crystals. Common cement types formed in meteoric, marine or burial environments are shown in Fig. 7.8, some of their fabrics in Fig. 7.12.
7.4.2.1 Cement Types
296
• termination (pointed, acute, chisel-shaped, pyramidlike, flattened), • appearance (clear, cloudy, with inclusions), • twinning, • crystal form (scalenohedral, rhombohedral, anhedral, subhedral; syntaxial overgrowth), • growth relation of crystals or subcrystals with respect to the substrate (normal, subnormal, oblique, arranged in fans, without preferred direction), • extinction type (straight, undulose, sweeping), • size (equidimensional, non-equidimensional, changes in size toward the pore center) and width/length ratio of crystals. The cements shown in Fig. 7.8 comprise five groups: • Cements consisting of crystals growing from a free substrate into the pore space (acicular, fibrous, botryoidal, radiaxial fibrous, dogtooth, bladed), • pendant cements (dripstone) underneath grains and grain-connecting cements (meniscus), • pore-filling cements forming calcite mosaics and characterized by crystals increasing in size toward the pore center (drusy), equidimensional crystals without substrate control (granular), or crystals without a preferred orientation (blocky), • overgrowth cements in optical continuum on a single crystal (syntaxial calcite overgrowth), • micritic cements exhibiting tiny peloids within a microcrystalline matrix (peloidal microcrystalline cement) or consisting of micron-sized calcite crystals forming fine-grained mosaics, rims or meniscus structures. Discussion of specific cement types: Acicular and fibrous cements: Some authors consider acicular and fibrous cements to be identical. The cements are similar in regard to the shape, appearance and growth relation of cement crystals, but differ in the termination and width of the crystals. Isopachous fibrous and radiaxial cements are common in ancient reef rocks (Fig. 7.7). They are dark colored in hand specimens and honey-colored in thin sections and form multiple bands consisting of densely packed fibrous crystals radiating from the substrate into growth cavities or solution cavities. In modern shallow-marine settings isopachous fibrous fringing cements formed by Mg-calcite (Schroeder 1973; Schroeder and Purser 1986; Aissaoui et al. 1986) or aragonite are common in marine-phreatic environments. In ancient reefs fibrous cements are composed of Low-Mg calcite (Walls and Burrows 1985). The difference in the mineralogical composition of modern and ancient isopachous fibrous cements is gen-
Carbonate Cements
erally explained by stabilization due to the action and influence of meteoric waters. Botryoidal cements: Pore-filling cements consisting of individual and coalescing millimeter-sized mamelons of compact, fibrous aragonite were first described by Ginsburg and James (1976) from Holocene reef cavities in Belize. Many botryoids contain thin discontinuity horizons marked by ‘dust lines’ of dark, irregularly shaped inclusions and/or slight irregularities in crystal growth, or small patches or thin layers of Mgcalcite micrite, or fragments of bioclasts (e.g. planktonic foraminifera) concentrated on discontinuity surfaces. Modern botryoidal aragonite cements contribute significantly to a very rapid cementation of steep marginal slopes with declivities between 25° and 45° (Grammer et al. 1993) and to the formation of seismic reflector horizons. Ancient botryoidal cements are known from Precambrian, Devonian, Carboniferous, Permian, Triassic and Tertiary reefs and slope deposits. Modern red algae belonging to the family Peyssoneliaceae have a hypobasal layer below the thallus, which may be strongly calcified although the thallus is non-calcified. Calcification consists of aragonite botryoids. This indicates that some ancient reef cements (especially those
Box 7.6. Common carbonate cements associations formed in different diagenetic environments. For the definitions see Fig. 7.8 and Fig. 7.12. Meteoric vadose: Types: Dripstone, meniscus, microcrystalline, bladed, fibrous, drusy, blocky. Cement crystals small and equant. Cements irregularly distributed, concentrated at grain contacts or beneath grains. Fabrics: Gravitational, meniscus, drusy and granular mosaics, syntaxial overgrowth. Marine vadose: Types: Dripstone, meniscus, fibrous. Fabrics: Isopachous, gravitational, meniscus, granular mosaics, syntaxial overgrowth. Meteoric phreatic: Types: Dogtooth, blocky, dripstone, meniscus, fibrous, microcrystalline, peloidal microcrystalline. Cement crystals are a bit larger than vadose crystals. Cements more homogeneously distributed than vadose cements. Fabrics: Circumgranular; isopachous, gravitational, equant, granular and blocky mosaics; syntaxial overgrowth. Marine phreatic: Types: Fibrous, acicular, bladed, botryoidal, radiaxial, dogtooth, microcrystalline. Fabrics: Circumgranular, isopachous, crusts, splays, botryoidal. Burial: Types: Blocky, poikilotopic, dogtooth, radiaxial. Fabrics: Drusy, equant and granular mosaics; syntaxial overgrowth.
Radiaxial Cement
297
Fig. 7.9. Radiaxial cement formed on the slope of a distally steepened ramp. Radiaxial calcite cements are common in depositional or diagenetic cavities of ancient reefs and slope settings. They consist of cloudy or turbid, fibrous to bladed crystals several millimeters long, growing perpendicularly to the substrate and exhibiting undulose extinction. The fabric is defined by a combination of curved twins, convergent optic axes and divergent subcrystals within single crystals. A: The pore space was filled with a first generation of marine inclusion-rich crystals, then covered by thinner (arrows) and thicker (double arrows) dark zones of reddish silt. B: A second radiaxial generation characterized by shorter and expanded crystals (arrows) can be seen under crossed nicols. It was formed subsequent to the cover of reddish silt on top of the crystals. The deposition of silt indicates influx from adjacent, subaerially exposed, and karstified platforms (Flügel and Koch 1995). The remaining pore space is occluded with coarse blocky cement. Polarized light. Late Triassic: Steinplatte, Tyrol, Austria. Scale is 1 mm.
developed in Late Paleozoic phylloid algal reefs) may be related to now-extinct algae (James et al. 1988). A specific type of botryoidal cement known from Late Paleozoic platform carbonates is array calcite (Davies and Nassichuk 1990) characterized by coarse spherulitic and botryoidal arrays of calcite replacing radial-concentric (botryoidal to spherulitic) pore-filling aragonite cements with a convex outer margin and internal concentric growth lines. Radiaxial fibrous cement (RFC): These calcite cements (Fig. 7.9) are most pervasive and volumetrically the most important cements in many Paleozoic and Mesozoic reef and fore-reef limestones (Pl. 28/2), but radiaxial fibrous cement is also known from ancient beachrocks and sheet voids (Gray and Adams 1995). They have been thought to be cements of syndepositional marine, shallow-burial marine (Halley and Scholle 1985), or subaerial meteoric origin. Current interpretations have focused on a neomorphic replacement by calcite of radiating bundles of acicular Mgcalcite that precipitated as early submarine cement (Davies 1977; Davies and Nassichuk 1990). Evidence of a marine-phreatic origin of RFC is the occurrence of
RFC as a first generation of cement, uncompacted shelter-pore fabrics, growth of RFC in marine and skeletal substrates, growth as isopachous rims around pores, interlayering of RFC and marine internal sediment, and analogues to modern submarine Mg-calcite cements. The present consensus favors a precursor Mg-calcite mineralogy. Evidence for a High-Mg calcite precursor is the textural similarity to modern Mg-calcite cements (Aissaoui 1988), microdolomite crystals within RFC, and relatively high amounts of Mg. Radiaxial fibrous cements were described in detail by Kendall and Tucker (1973) and revised by Kendall (1985), who differentiated three categories distinguished by the extinction patterns of subcrystals within each larger crystal (Pl. 35/2): • Radiaxial fibrous calcite exhibiting a pattern of subcrystals within each subcrystal that diverge away from the substrate in an opposing pattern of distallyconvergent axes, yielding a corresponding curvature to twin lamellae and cleavage (Bathurst 1959, 1977, 1982; Kendall and Tucker 1973; Kendall 1985). • Fascicular-optic calcite (Kendall 1977; Chafetz 1979; Marshall 1981) characterized by the presence of
298
Overgrowth Cement and Megacement
• Syntaxial overgrowth as a filling of secondary pore space created by the selective dissolution of micrite matrix in the vicinity of echinoderm grains (solution corona model: Walkden and Berry 1984). The model involves mechanisms of early near-surface meteoric dissolution of the micritic matrix accompanied by preferentially downward enlarged overgrowth on the surface of the echinoderm grains. • Neomorphic replacement of micrite operating across a thin solution film (Bathurst 1958; Orme and Brown 1963). This model has been used in many studies. • Syntaxial overgrowth into primary sheltered voids beneath echinoderm fragments (Görur 1979). The selective displacement of micrite beneath the skeletal grains may be facilitated by the high permeability of the pores of the echinoderm stereome acting as pathways for undersaturated vadose water that precipitates gravitative cements. Fig. 7.10. Echinoderm overgrowth rim cement. Syntaxial overgrowth on an echinoid fragment exhibiting relicts of the original net-like microstructure (1). The overgrowth cement (2) occurs in optical continuity with the monocrystalline host grain (3), as indicated by the corresponding orientation of the faint cleavage planes. The skeletal grain acted as a substrate for radiaxial cements whose crystals are cut in various directions (4). Crossed nicols. Upper part of a carbonate ramp. Late Triassic: Steinplatte, Tyrol, Austria. Scale is 1 mm.
a diverging pattern of optical axes coinciding with that of subcrystals. • Radial-fibrous calcite consisting of turbid crystals without undulose extinction (Davies 1977; Mazzullo 1980). Syntaxial calcite overgrowth cement: Overgrowth of clear rim cements in optical continuity with large monocrystalline host grains (e.g. crinoid ossicles, echinoid plates; Fig. 7.10; Pl. 33/7) is a common texture in wackestones, packstones and grainstones. These cements are easy to recognize, because echinoderm grains are inclusion-rich and differ in color from clear inclusion-poor cement (Pl. 34/3, Pl. 95/6). The origin of cement overgrowth on echinoderm fragments is explained by • syntaxial overgrowth before the deposition of marine muds (Evamy and Shearman 1965, 1969). The authors showed by staining that each individual bar of the echinoderm stereome generated its own microcrystal. Contact of the microcrystals during subsequent growth causes identical orientation and integration, producing the characteristic ‘single crystal’ cements of many echinoderm grainstones. Departures from this model were described by Braithwaite and Heath (1987).
Most authors interpret the syntaxial overgrowth cements to be of meteoric diagenetic origin (e.g. Longman 1980; Kaufman et al. 1988). However, overgrowth cements on echinoderms are also common in ancient limestones formed in marine environments (Walker et al. 1990). These thin cloudy overgrowth cements often represent the first generation of overgrowth on echinoderm grains followed by a continued growth under meteoric conditions. Megacements: Cements consisting of conspicuously large, millimeter- to centimeter-sized calcite crystals have been reported from different environments: • Large seafloor calcite cements were recorded from the Proterozoic and interpreted by strong oversaturation of the seawater (Grotzinger 1994; Sami and James 1994) and from Lower Triassic shelf sediments believed to have been deposited in oxygen-restricted settings subsequent to the Permian-Triassic catastrophic event (Woods et al. 1999). • Another example of megacements are crusts of palisade cements consisting of large, up to 10 cm long, black calcite crystals occurring in sheet cracks, cavities, and neptunian dikes of basinal limestones, and as reworked constituents of debris flows (Wendt and Belka 1991). These cements contain ghost structures of former marine fibrous aragonite cements that were neomorphically replaced by calcite megacements, and later corroded, broken and deposited in debris flows. • Similar megacements are represented by raggioni cements (Assereto and Folk 1980; Davies and Nassichuk 1990; Iannace 1991, 1993) characterized by giant rays growing in crusts and lining the walls of centimeter- to decimeter-sized cavities within tepee
299
Cement Fabrics
formed in the upper part of the vadose zone can be transported in suspension and deposited in voids as internal silt-sized sediment that may show grading and smallscale cross-lamination. Crystal silt may filter down 12 m below submerged paleosurfaces (Aissaoui and Purser 1983) and is often used as an indication of vadose diagenetic conditions (Fig. 7.11). This interpretation should be applied with some caution, because similar or identical silt-sized internal fillings are also known from burial settings.
7.4.2.2 Cement Fabrics
Fig. 7.11. Crystal silt within a solution void. The void is lined with scalenohedral dogtooth cement (arrows point to the variously shaped terminations of crystals) and was subsequently filled with very small calcite crystals. Note the conspicuous textural difference between the silt and the dark microcrystalline matrix of the limestone. The void was formed in carbonates representing the slope between a shallow-marine platform and a basin. Dogtooth cement and crystal silt record the influx of meteoric water from exposed parts of the nearby platform (Flügel and Koch 1995). Late Triassic: Steinplatte, Tyrol, Austria. Scale is 1 mm.
structures. Raggioni are typically 20–30 mm long, may reach 250 mm. The width of the largest ray crystals is about 4 mm. Terminations are square-ended or feathery. Raggioni are interpreted as pseudomorphs of calcite after aragonite cements formed in meteoric and marine-vadose, hypersaline as well as marine-phreatic waters. Crystal silt – internal sediment or cement ? The term crystal silt (Fig. 7.11; Pl. 34/8 Pl. 39/8) describes a poorly sorted gray internal matrix consisting of tiny, often rhombic, sometimes angular and very angular calcite crystals of 5–40 µm size. It occurs in molds, solution-enlarged molds, fenestral voids, interparticle pores, and small reef cavities. Crystal silts have various origins. They are known from continental (e.g. caliche), transitional (e.g. tidal) and marine environments, and have been interpreted as vadose crystal silt of supratidal origin (Dunham 1969), a product of the destruction of limestones and dolomites (Upper Jurassic; Lang 1964), marine internal sediment (because of the association of silt and marine microfossils), or reworked microcrystalline cements. Calcite crystals,
By examining the spatial relations between cements, grains, substrate and pores, cement fabrics can be subdivided into five groups (Fig. 7.12): 1 Symmetrical cements around grains represented by isopachous and circumgranular cement rims, distinguished by the cement type and thickness of the cement rim. 2 Asymmetrical cements concentrated on the underside of grains or developed at grain contacts. 3 Cement crusts, fans and hemispheres forming millimeter- to centimeter thick structures. 4 Calcite mosaics, differentiated by size, orientation and the type of crystal boundaries. 5 Syntaxial overgrowth cement, usually associated with echinoderm grains.
7.4.3 Cement Types and Diagenetic Environments Cement types can be diagnostic of particular diagenetic environments, but caution is required, because identical cement types may be formed in different diagenetic environments, as documented, e.g. by equant drusy spar cement precipitated in near-surface meteoric environments and under deep burial conditions. Careful attention must be paid to isotopes, trace elements, and cathodoluminescence fabrics to correctly deduce the precipitation environment. Diagnostic cement criteria: Meteoric cements exhibit • commonly Low-Mg calcite, because meteoric water has a very low Mg/Ca ratio, Mg is leached and HighMg calcite is converted to Low-Mg calcite, • aragonite dissolution and precipitation of drusy calcite spars, • blocky, granular or mosaic fabrics that fill the pores completely,
300
Cement rims around grains (symmetrical cements) Isopachous: Characterized by single or multiple cement rims growing with equal thickness around grains. The cement rim may consist of fibrous, bladed, or microcrystalline crystals. Thickness of the rim within the range of tens of microns to several millimeters. Common in marine-phreatic and marinevadose environments. Circumgranular: Characterized by a cement rim around grains, consisting of equidimensional crystals forming the first generation of pore-lining cements. The rim is commonly thinner than isopachous cement rims. Meteoric phreatic environment.
Cement rims restricted to the underside of grains and void roofs (asymmetrical cements) Gravitational: Pendant beard-like cements (A) beneath grains. Often associated with bridging cements (B, meniscus cement) which connect adjacent grains. Gravitational and bridging cements (crossing pores and connecting grains), e.g. meniscus cement and microcrystalline cement, are irregularly distributed and absent in many pores. Meteoric-vadose, meteoric-phreatic and marine-vadose environments.
Cement Fabrics Splays: Fan-like structure consisting of fibrous outward spreading calcite crystals. The structures may occur isolated or within marine cement crusts. Botryoids: Complex structures consisting of various dome-shaped hemispheres built by radiating fibrous calcite (originally aragonite) crystals and crystal fans. Formed on free surfaces as well as in marine cavities.
Pore-filling cement mosaics Drusy mosaic: Characterized by pore-filling calcite crystals increasing in size toward the center of interparticle pores or voids. Crystals with compromise boundaries (plane intercrystalline boundaries generated by two crystals growing alongside each other). Burial and near-surface meteoric environments. Equant mosaic: Characterized by small pore-filling calcite crystals of approximately equal size. Subhedral and anhedral crystals with welldeveloped boundary faces. Granular mosaic: Characterized by small pore-filling calcite crystals without a preferred orientation and no substrate control. Meteoric-vadose, meteoric-phreatic and burial environments.
Large cement structures exhibiting geometrical patterns Crusts: Millimeter- to centimeter thick crusts consisting of calcite cements (fibrous, radiaxial fibrous, microcrystalline) growing on extended substrates (e.g. pore walls, hardgrounds, shells). Cement crusts may consist of one or more growth zones and may display internal differentiations (e.g. alternating crusts formed by fibrous or radiaxial cements and festooned cellular crusts). Marine and meteoric environments. Chevron crusts, characterized by V- and inverted V-shape patterns, may sometimes be caused by neomorphic processes.
Overgrowth Syntaxial echinoderm overgrowth: Fabric characterized by the dominance of syntaxial calcite cement, formed as overgrowth usually on echinoderm skeletal grains within the sediment. In many places in optical continuity with a substrate of the same mineralogy.Vadose, meteoric-phreatic, and burial environments.
Fig. 7.12. Cement fabrics.
• isopachous and syntaxial cements in the phreatic zone and gravitational and microstalactitic cements in the vadose zone, • vadose cements formed in freshwater environments (meteoric-vadose cements) and in near-coast marine environments (marine-vadose cements), both above the water table, • sparry calcite cements commonly found either in the meteoric-phreatic zone at near-surface conditions, or buried relatively deep (10s–100s m), where meteoric water has become more saline. Marine-vadose and mixing zone cements are mostly gravitational cements and include cements with bladed relict textures, cements with the incipient growth of
dogtooth cements, radiaxial fibrous cements, and gravitational cements with algal relict textures. Marine cements have characteristic crystal habits and geochemical signatures. Common cement fabrics assisting in the identification of ancient marine cements are bladed calcite, fibrous calcite, microcrystalline cement, and some syntaxial overgrowth cements. The best indicator of marine diagenesis are hardgrounds (see Sect. 5.2.4.1). Common burial cements exhibit • clear coarse calcite spar (calcspar) and coarse equant spar (ESC) characterized by dull cathodoluminescence,
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Effect of Burial Diagenesis
• equant spar forming drusy or intergranular mosaics, • bladed prismatic calcite spar, dark cathodoluminescence. Larger and clearer as compared to first generation bladed marine cement, • poikilotopic fabric, characterized by large crystals, • syntaxial overgrowth cement, most commonly with echinoderm grains, • coarse dolomite cement (saddle dolomite and xenotopic, characterized by a mosaic of 0.1 to 1 mm-sized crystals), • coarse anhydrite cement. Equant spar calcite cements (ESC) are characterized by equigranular limpid crystals exhibiting a distinct extinction. Crystal boundaries are irregular, in places serrated. Crystals are large, up to several millimeters. The cements are common in voids, often occluding the last remaining pore spaces. Interpreting equant spar calcite cements is difficult, because they are formed under deep burial conditions as well as in near-surface environments. A high ferrous iron content is common in burial ESC, but can also exist in early cements formed in near-surface anoxic conditions (Scholle and Halley 1985). A deep-marine origin of equant spar cement has been reported by FreemanLynde et al. (1986) in Bahama escarpment limestones.
The effect of burial diagenesis on platform and basinal carbonates: The burial diagenesis of platform and basinal carbonates exhibits similar effects (increase in temperature and pressure with depth, porosity reduction, squeezing of pore waters, dehydration of minerals, increase in the influence of saline pore waters, and chemical compaction following mechanical compaction), but is different with regard to other criteria: Platform carbonates are characterized by (1) rapid sedimentation; (2) variable pore waters ranging in salinity from freshwater and brackish to marine and hypersaline; (3) fine- to coarse grained sediments often associated with evaporites; (4) metastable and stable carbonate minerals; (5) variable diagenetic and compaction potential; (6) possibility of the influx of meteoric water; (7) carbonate source for cement provided initially provided by meteoric dissolution, later by pressure solution; (8) decreasing porosity with depth and overburden, but strongly varying with pore water flux rates and different compaction intensities depending from different primary sedimentary fabrics. Basinal carbonates are characterized by (1) slow sedimentation; (2) initially marine pore waters; (3) finegrained sediments often associated with clay and or-
ganic matter; (4) stable carbonate minerals; (5) low diagenetic and compaction potential; (6) no possibility of influx of meteoric water; (7) carbonate source for cementation provided by pressure solution; (8) exponential decrease with depth and overburden. Very deep burial diagenesis Studies of very deeply buried carbonates have been conducted along two parallel lines: (1) Experimental compaction of natural carbonate sediments (Bhattacharyya and Friedman 1984) and (2) investigation of carbonate rocks from boreholes using fluid inclusions and thin-section petrography (Friedman 1987). Different burial criteria have been found in Paleozoic carbonates studied in deep wells (6 000 to 9 000 m; temperature around 210°; hydrostatic pressure around 2.5 kilobar: Friedman et al. 1984): • Grainstones exhibit twinning and cleavage of calcite spar crystals. Mechanical displacement along cleavage planes occurs within single echinoderm crystals causing grain elongation. Twin lamellae appear more narrowly set in deep burial than in shallow burial settings. Multiple displacement of differently oriented twin lamellae is common in deep burial samples. Further criteria are bent twinning lamellae and diminution within echinoderm single crystal. • Wacke- and packstones often display no thin-section criteria, pointing to alteration of skeletal grains. Bryozoan and ostracod fragments may have been protected by the surrounding micrite. The size of micrite matrix crystals is enlarged up to 30 µm. The texture resembles that of fine-grained marbles and is characterized by patchy extinction.
7.4.4 Facies-Controlled Diagenesis Cementation is strongly facies-related. Marine cementation can be traced from beaches to deeper water and depends on the porosity and fabric of the sediment. Beaches and tidal settings are characterized by alternating meteoric-vadose, marine-vadose, and marine conditions that contribute to the formation of beachrocks, hard carbonate crusts and diverse vadose and marine cement types. Common cements are microcrystalline, bladed or fibrous Mg-calcite, and aragonite cements.
7.4.4.1 Carbonate Platforms and Ramps The spatial distribution of interparticle cement in carbonate sediments is primarily a function of the en-
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Platform Diagenesis
Fig. 7.13. Diagenesis across the Bahama Banks (modified after Demicco and Hardie 1994) demonstrated by the distributional patterns of diagenetic processes, products, and early diagenetic cement types over the shallow platform and in adjacent deep-marine slope and basinal settings. Inter- and intragranular cements occur in tidal flat and shelf lagoon peloid muds, in bank-margin sands, in bank-margin reefs, in foreslope and periplatform debris. These cementation sites have several criteria in common: sedimentation and erosion rates are low, host sediments have the same mineralogy as cements (aragonite and High-Mg calcite) and can thus act as seeds, and large volumes of supersaturated waters are pumped through the sediment. However, the specific conditions that promote the precipitation of CaCO3 are different in each of these environments. Partial lithification of tidal flat, lagoonal and sand shoal sediments produces hard surface crusts, hardgrounds, and hardened grains. Intragranular cementation in fecal pellets and sand-sized mud aggregates turn soft muds of the outer lagoon deposited under low-energy conditions into hard sand grains with significantly different hydrodynamic properties. In contrast to the tidal and lagoonal sediments, the framework and debris of bank-margin reefs are extensively cemented due to the immense volumes of seawater flushed by tides and waves through the reef, thus providing the prerequisite for the growth of pore-lining cements.
vironment as exemplified by the diagenesis across the Bahama Banks (Demicco and Hardie 1994; Fig. 7.13). Cementation by calcite or aragonite precipitation is coeval with sedimentation. Early marine cementation occurring in shallow-marine plaforms and ramps at or near the sediment-seawater interface depends on several prerequisites: • Very low sedimentation and erosion rates. Low sedimentation rates and temporal non-deposition connected with increased biogenic encrustations (e.g. in deepermarine environments) will enhance the possibility of submarine cementation (formation of hardgrounds). • Large volumes of sea water are pumped through the sediments. Early cementation depends on sufficient carbonate. This is supplied by transport through the sediment by waves, tides, currents, or evaporation processes. Strong cementation, therefore, is common at seaward/windward sides of tidal and shallow subtidal areas (e.g. carbonate sand shoals), but less frequent on the protected leeward part of shallow-marine environments. • Permeability, determining the flow of the sea water
through the sediment and a sufficient supply of cementing carbonate. • Appropriate water temperatures and salinities. • Biogenic activities producing pores (e.g. borings and burrows) and stabilizing substrates (e.g. biofilms, algal mats, sea grass areas) and preventing reworking of carbonate grains. Sponge borings, e.g. may play a critical role in the aragonite-to-calcite inversion process of shells (Rehman et al. 1994) providing conduits for meteoric alteration. • Metabolic activity (photosynthesis and respiration) as well as disintegration of organic material causing dissolution of carbonate and controlling pCO2 and the degree of supersaturation (Jahnke et al. 1994).
7.4.4.2 Reefs The most significant amount of cementation and therefore porosity reduction, occurs in reef environments with higher rates of agitation and/or lower sedimentation rates (Lighty 1985). Common patterns are (a) high
Reef Cements
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Fig. 7.14. Reef cements. Many Middle and Late Permian reefs are characterized by frameworks consisting of low-growing encrusting organisms and up to 60% of coeval marine cement. A: Sequence consisting of submarine radiaxial fibrous cement (1) growing on Archaeolithoporella and blocky calcite cement (2). The marine cement forms yellowish isopachous crusts (0.5 to 2.5 mm) surrounding small patches of sessile organisms. The crust consists of ray crystals containing ghosts of thin straight-edged, fibrous crystals with a width of up to 25 µm. The crust shows a faint zonation. Note the distinct and sharp upper boundary of the cement crust. The fabric of this cement indicates an aragonite precursor cement. In contrast, the overlying coarse blocky cement (2) consists of euhedral, nonferroan calcite crystals that have suffered aggrading neomorphism also partly cannibalizing its underlying marine cement. The outer sometimes rounded and irregular boundary of these blocky cements may be explained by solution effects. The remaining pore space was filled by fine-grained, silt-sized carbonate with abundant, sometimes euhedral dolomite crystals (3). B: Radiaxial fibrous cement growing on biogenic encrustations mainly made by Archaeolithoporella (see Pl. 98/8). The outer part of the laminated structure has been eroded and filled with peloids (arrow) prior to the formation of the marine radiaxial fibrous cement exhibiting a distinct aragonite precursor fabric. The limestone was affected by fracturing and shearing. Microcracks are closed by several calcite generations. Interpretation: The sample comes from a reef formed on the upper platform slope (Flügel et al. 1984). The framework was stabilized by pervasive marine cementation. The white, bladed calcite cement originated in a shallow burial environment influenced by meteoric waters. The infilling of a dolomitized carbonate silt points to a change towards vadose conditions. Middle Permian (Guadalupian): Straza near Bled, Slovenia. Scale is 2 mm.
Box 7.7. Selected case studies of reef diagenesis. An overview is given by Schroeder and Purser (1986). Marine cementation, concurrent internal sedimentation, and bioerosion are the most important factors controlling early marine lithification within many reef environments. Holocene: Buddemeier and Oberdorfer 1986; Dullo 1986; Friedman 1974; Ginsburg et al. 1971; Jindrich 1983; James et al. 1976; Land and Goreau 1970; Lighty 1985; Macintyre et al. 1968, 1977; Marshall 1986; Marshall and Davies 1981; Schroeder 1972 Tertiary: Aissaoui et al. 1986; Armstrong et al. 1980; Asquith and Drake 1985; Pomar et al. 1994; Schroeder 1986; Sun and Esteban 1994 Cretaceous: Alsharhan 1985; Enos 1986; M’Rabet et al. 1986; Reitner 1986 Jurassic: Benito et al. 2001; Joachimski and Scheller 1987; Koch and Liedmann 1996; Koch and Schorr 1986; Matyskiewicz 1989; Schroeder et al. 1996; Sun and Wright 1998 Triassic: German 1968; Henrich and Zankl 1986; Lakew 1994; Mazzullo et al. 1990; Russo et al. 2000; Satterley et al. 1994; Sotak and Lintnerova 1994; Zankl 1971; Zeeh et al. 1995; Zorn 1971 Permian: Chaiuachi and M’Rabet 1989; Cross and Klosterman 1981; Dunham 1972; Given and Lohmann 1986; Liu and Rigby 1992; Mazzullo and Cys 1977; 1979; Mazzullo et al. 1990; Scherer 1986; Schmidt 1977; Stemmerik and Larsen 1993; Tucker and Hollingworth 1986; Wendt 1977 Pennsylvanian: Davies 1977; Davies and Nassichuk 1990; Soreghan and Giles 1999; Wahlman 2002 Mississippian: Ahr and Ross 1982; Bathurst 1959; Miller 1986 Devonian: Hurley and Lohmann 1989; Kaufmann et al. 1999; Kerans et al. 1986; Krebs 1969; Machel 1986; Mountjoy and Walls 1977; Wong and Oldershaw 1981; Schneider 1977; Walls and Burrowes 1985; Webb 2002 Silurian: Cercone and Lohmann 1986; Frykman 1986; Petta 1980; Savard and Bourque 1989
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water agitation and low sedimentation rates = strong cementation (e.g. reef front) and (b) low agitation, high sedimentation rates and substantial deposition of finegrained sediment = low cementation (e.g. back reef). Active cementation in reefs seems to be confined to the outer few feet of the sediment. Water energy and circulation control the intensity of cementation, local permeability controls the degree and type of cement on a small scale. Wave pumping and biological activity have a direct bearing on the formation of reef cements. Another important factor influencing cementation and
Reef Cements
mineralogy of cements is the size of framework pores. Small pores can trap sediments, ultimately producing cemented cavity fills and cement crusts. In large pores fine-grained sediment will be quickly flushed through the voids and cement will be precipitated (James et al. 1976; Marshall and Davies 1981; Marshall 1983). Reef cementation takes place just centimeters beneath the sea floor; the framework reefs are commonly lithified within one-half meter or less. Substrate control is a common feature as demonstrated by the precipitation of aragonite cements on aragonitic corals, and Mg-Calcite
Plate 33 Marine and Meteoric Carbonate Cements The plate exhibits typical carbonate cements formed in marine environments (–> 1, 3, 7, 8), in the marinevadose zone (–> 2, 3) and in meteoric environments (–> 4 ). Note open, usually interparticle porosity in all samples. 1 Circumgranular bladed High-Mg calcite cement (arrow) surrounding conspicuously well-rounded skeletal grains (foraminifera – F; coralline algae – CA). Note open interparticle pores (P). Roundness/sphericity values correspond to the field D, N and E of the Krumbein visual chart (Fig. 4.30). Bladed cements often form isopachous rinds and constitute the first generation of marine cements. Shape and orientation of substrate crystals may strongly control crystal habits (Wilson and Palmer 1992). Marine-phreatic environment. Bermudas. 2 Thin isopachous microcrystalline calcite cement (black arrows) is developed around various bioclasts (coralline algae – CA, miliolid foraminifera – MF, micritized bivalve shell – BS, open pores – P). Interparticle pores are filled with skeletal debris and peloids. Note the low sphericity and medium roundness of the red algal fragments (fields L, P, R of the Krumbein chart, Fig. 4.30). White grains are airborne quartz. Interparticle pores appear black. Marine-vadose environment. Pleistocene eolian dune: Carthage, northern Tunisia. Crossed nicols. 3 Thick bladed equidimensional cement (white arrow) on skeletal grains (foraminifera – F), formed in a marine-vadose environment and thin, poorly developed Mg-calcite cements (black arrow) on black lithoclasts (‘black pebbles’; see Sect. 4.2.8). No cements are present on the bivalve shell (BS). Open interparticle pores (P) filled with resin. Marinevadose environment. Beachrock: Holetown, Barbados. 4 Meteoric-vadose cements. Grains are connected by meniscus cement (black arrow) causing pore rounding, and pendant cements (white arrows), formed on droplets beneath grains within the meteoric-vadose environment. Pendant (or gravitational) cements are absent on the upper surface of the grains but thicken down the sides to the undersurface. Note the good preservation of calcitic bivalve fragments (BF) without micrite envelopes, in contrast to the poor preservation of originally aragonitic shell fragments (SF) exhibiting micrite envelopes around casts and filling with coarse calcite (CC). Interparticle pores (P) are still open. Holocene: Belmont, Bermudas. 5 Partly dolomitized foraminiferal grainstone with Amphistegina. Matrix and Mg-calcite skeletal grains are replaced by anhedral, finely crystalline dolomite mosaics (appear black), whereas the radial hyaline tests of the foraminifera are preserved as calcite. Arrows point to isopachous, bladed to fibrous dolomite cements developed in interparticle voids. Geochemical data suggest that seawater of slightly variable temperature and/or slight degrees of evaporation, and enriched in bicarbonate originating from the oxidation of methane, was the principal agent for the replacement dolomitization (Machel and Burton 1994). Late Pleistocene (fore-reef facies): Golden Grove, Barbados. 6 Partly dolomitized Amphistegina limestone exhibiting large secondary vuggy porosity. The large dissolution void is filled with thin isopachous bladed dolomite cement (white arrows) and occluded by recrystallized coarse, sparry calcite cement (black arrows). Skeletal grains are coralline algae (CA) and foraminifera (F). Same locality as –> 5. 7 Marine-phreatic isopachous carbonate cement (arrow) with irregular surface, surrounding bioclasts. Syntaxial overgrowth on echinoderm fragment (right). The dredge sample comes from sediment deposited on a seamount exposed to seawater. Diagenesis of seamount sediments is highly complex and includes cementation, transformation of carbonate phases, phosphatization, silicification and mobilization of Fe-Mn oxides (Grötsch and Flügel 1992). Early diagenetic cements of the drowned platforms are rarely preserved but exhibit a still open interparticle porosity (black). Cretaceous (Albian reef facies): Darwin Guyot, northwestern Pacific. 8 Isopachous rim cement composed of bladed crystals (arrows), growing within the mold of a dissolved aragonitic shell (AS). Open moldic porosity is preserved due to the micrite envelope (ME) coating. Primary calcitic mollusk shells (CM) show remnants of microstructures. Cretaceous (Late Albian): Charly Johnson Guyot, northwestern Pacific.
Plate 33: Marine and Meteoric Cements
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cements of Mg-calcitic coralline algae. Reef cements are bladed or microcrystalline Mg-calcite or aragonite. Microcrystalline cement is commonly pelleted. The fabric of the reef cement is controlled by its access to circulating water. Voids with greater permeability may be filled by isopachous or equant cements, more confined voids with larger and irregular crystals. Thick botryoidal crusts of aragonite occur in larger reef cavities. Common reef cements are radiaxial cements (Sect. 7.4.2.1), botryoidal cements(Fig. 7.14), isopachous fibrous aragonite cements, and isopachous fibrous High-Mg calcite cements forming blades. Syngenetic aragonite cements forming fans and composite fans of acicular crystals are common constituents of Permian and Triassic
Cold-Water, Warm-Water Diagenesis
reef limestones (Fig. 7.14; Mazzullo and Cys 1977; Flügel et al. 1984; Davies and Nassichuk 1990; Liu and Rigby 1992). Mg-calcite is the dominant cement in most modern and many ancient reefs, and occurs mostly as rim cements, microcrystalline crusts, or as peloidal cements that geopetally fills cavities in the reef framework. The syngenetic origin of reef cements is indicated by (a) multiple cement generations alternating with cement exhibiting biogenic encrustations, microborings, and peloidal sediment; (b) reworking of cement crusts and subsequent cementation of the clasts. ‘Dust lines’ around and between the cement fans are interpreted as being of microbial origin (initial growth of aragonite crystals within biofilms).
Plate 34 Marine, Meteoric, and Burial Carbonate Cements This plate displays marine (–> 1, 7), meteoric (–> 5,6) and burial cements (–> 1, 3, 4). Thin sections allow a first interpretation of crystal morphology and fabric with regard to the diagenetic environment. A critical assessment of the environment requires cathodoluminescence fluid inclusions analyses as well as and stable isotope studies. 1 Bioclastic grainstone. Interparticle pores between fossils enclosed within oncoids (T: trilobite) are filled with submarine calcite cement (thin rim consisting of acicular crystals, black arrow) and blocky burial cement (white arrow). The trilobite fragments show the typical ‘shepherd’s crook’ shape (see Pl. 94/3). Late Silurian (Ludlow): Burgwik, Hobbugen, Gotland Island, Sweden. 2 Radiaxial calcite cement growing on trilobite fragments. The large radiaxial cement crystals have curved boundaries and show a cloudy appearance due to inclusions. Note the distinct pores typical of some trilobite carapaces. The shell consists of prismatic calcite crystals growing with their c-axes perpendicular to the skeleton surface. Early Devonian (Emsian): Erfoud, Anti-Atlas, southern Morocco. 3 Syntaxial overgrowth (arrow) on echinoderm fragments. Because echinoderm fragments act as single crystals, they are commonly nucleation sites of calcite cements. Clear calcite is added in optical continuity to the echinoderm fragment. Sediment adjacent to the monocrystal crinoid may have been dissolved as the syntaxial overgrowth developed. The crinoid fragment at the left side of the photograph shows distinct cleavage lamellae. Syntaxial overgrowth by clear crystals points to burial diagenesis. In contrast, syntaxial overgrowth in near-surface marine, meteoric or mixing-zone environments is generally inclusion-rich and cloudy. Cambrian: Handler Range, Antarctica. 4 Syntaxial overgrowth (arrows) on echinoderms. When viewed under crossed nicols, the cement crystals show the same interference colors as the enclosed single crystal echinoderms and are part of the same crystals typical of syntaxial overgrowth. Prolongation of twin lamellae into the syntaxial cement as well as the inclusion-free clear crystals indicate the continuing growth of the echinoderm calcite under burial conditions. Jurassic: Switzerland. 5 Leached corals, filled with internal sediment and calcite. The walls of the corals are almost completely replaced by calcite. Some septal structures are preserved as relicts at the periphery (arrow). Dissolution occurred during the subaerial exposure of a reef and karstification (Bernecker et al. 1999). Infilling of (red) micrite sediment took place subsequent to a renewed change to submarine conditions. Late Triassic: Tropfbruch Adnet, Salzburg, Austria. 6 Gravitational cement represented by pendant cement (arrows) beneath variously shaped vadoids formed in a marinevadose shelf interior environment. Pendant cements reflect growth outward to capillary water-air interfaces of partly filled interparticle pores. Pisoid grainstone. SMF 26. Late Permian: Entrance Carlsbad Cavern, New Mexico, U.S.A. 7 Marine phreatic isopachous bladed calcite cement (arrows) around dasyclad algae and aggregate grains. Remaining interparticle pores were filled with crystal silt. Crystal silt composed of tiny calcite crystals, sometimes associated with microfossils. This internal sediment is often used as an indication of vadose diagenetic conditions (as in this figure), but is also known from burial environments. Middle Permian: Velebit, Croatia. 8 Solution void, occluded by scalenohedral calcite cement (black arrow) and silt-sized sediment (white arrow) consisting of tiny mechanically deposited calcite crystals and broken-off tips of cement crystals (crystal silt). See –> 7. Lowermost Cretaceous: Subsurface, Kinsau, Bavaria, Germany.
Plate 34: Marine, Meteoric, Burial Cements
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7.4.4.3 Cold-Water vs. Warm-Water Diagenesis Carbonate minerals are predominantly High-Mg calcite (and minor aragonite) in shallow warm temperate carbonates and High-Mg to Low-Mg calcites in cool temperate carbonates. Mineralogical composition is controlled by water temperatures and water depths. Carbonate saturation of seawater in non-tropical settings is less pronounced or even undersaturated (Opdyke and Wilkinson 1990). Diagenesis, therefore, is predominantly destructive (Alexandersson 1978). In comparison with tropical shallow-marine carbonates the diagenetic potential is low and sediments may remain essentially unlithified for a long period of time. Nevertheless, intraskeletal and intergranular carbonate cemen-
Diagenetic Pathways
tation may be common in some temperate and coldwater carbonates (Rao 1981, 1993, 1994; Rao and Adabi 1992; James et al. 1992, 1994; Freiwald 1998). In contrast to warm-water reefs, Mg-calcite radiaxial cements appear to dominate in cooler waters (Beauchamp 1994; Nelson and James 2000). Dolomite is rare, known from the cool temperate water of the Coorong, South Australia, and locally present in New Zealand.
7.4.5 Diagenetic Pathways and Patterns Pore waters and their products, including cement and porosity evolve over time resulting in short- and longterm diagenetic sequences and stages. Changes over
Plate 35 Diagenetic Pathways and Spatial Variations A detailed study of diagenetic spatial variations is necessary for understanding porosity development over time and the highly complex diagenetic history of carbonate rocks. Instructive examples of these processes were presented by Schroeder (1988). Spatial variations depend on the individual characteristics of grains and cements (mineral and chemical composition, ultra- and microstructure, surface properties), collective characteristics of sediments and sedimentary rocks (component mix due to ecological and transport selection, fabrics as determined by primary sedimentary and skeletal structures), synsedimentary to early diagenetic processes (internal sedimentation, bioturbation, bioerosion), early to late diagenetic processes (cementation, dissolution, fracturing, decomposition of organic matter), hydrology (wave and tide ranges, sea- and ground-water levels, mixing zones), as well as substrate mobility and solubility. Many of the causal factors are controlled by biologic, meteorologic and oceanographic factors. The samples shown in this plate illustrate a few variations in cements and neomorphic alterations as well as in grain preservation, both between and within particles. Diagenetic pathways are highly complicated. Therefore attributing the depicted diagenetic features to specific diagenetic environments has been restricted to the samples described here and should not be too generalized. 1 Cementation in a Holocene beachrock from Uyombo, Kenia. The photograph shows inherited aragonite spherulitic cement (IAC) within an intraskeletal pore of a coralline red alga (CA) and fibrous marine Mg-calcite palisade cement (MC) surrounding carbonate particles but not the quartz grains (Q). Meteoric sparry calcite cement (SC) fills the remaining pore spaces. Crossed nicols. 2 Particle alteration in a Holocene beachrock from Uyombo, Kenia. A fragment of a mollusk shell exhibits a variety of modes of preservation/alteration: The original fabric is partly preserved and rimmed by a micrite envelope. A secondary solution pore is partly empty and partly filled by blocky sparry calcite cement. Interparticle pores are filled with fibrous Mg-calcite cement. CA: Coralline alga; Q: quartz. Crossed nicols. 3 Modern Bermuda algal cup reef rock. Adjacent skeletal pores within barnacles are filled completely by aragonite needle cement (1), by a basal layer of micrite and subsequent Mg-calcite cements with calcified algal filaments (2, 3), or by Mgcalcite cement on calcified algal filaments and subsequent aragonite needle cement (4). Crossed nicols. 4 Branching reef coral Procladocora tenuis. The corallite is filled by micrite (1), fine-grained sediment (2), sparry calcite (3) and neomorphic calcite (4). The Paleocene coral knobs studied passed through a sequence of diagenetic environments ranging from marine to meteoric and back to marine, on to shallow burial conditions at various depths, and then back to the surface. In the course of this development various portions of the corals were affected at different times by neomorphism, replacement, dissolution, cementation, and/or fracturing. As a result of this history, various incomplete diagenetic sequences occurred in spatial proximity, many of which can be recognized in a single thin section. Infilling of intraskeleton pores with fine-grained sediment (2) and micrite (1) occurred in the marine environment. This micrite may be a sedimentary or diagenetic product. The sparry calcite mosaic (3) represents neomorphically altered Low-Mg calcite cements (indicating meteoric processes). The large crystals (4) with irregular boundaries and relicts of septa are caused by neomorphic alterations, probably in the burial environment. Crossed nicols. Paleocene: Bir Abu El Husein area, southern Egypt. –> 1–4 Schroeder (1988)
Plate 35: Diagenetic Pathways
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Diagenetic Textures
Fig. 7.15. Diagenetic changes of coral skeletons over time demonstrated by diagenetic patterns of dendroid scleractinian corals. The sketch summarizes different preservation types and the course of diagenetic processes. The generalized drawings are based on thin sections (see Pl. 36/4) of Cretaceous coral limestones used as building stones in the giant wall of the Hittite capital Hattusa near Bogazköy, central Turkey. A: The septal structure of the originally aragonitic dendroid corals showing biogenic encrustations is rarely completely preserved, whereby a geopetal infill indicates toppling of the corallites. B: Preservation of coral structures can be enhanced by the precipitation of fine-crystalline drusy calcite mosaics between the septa. More often, however, the interior of the corallites is completely dissolved, producing open molds. The molds are stabilized by variously thick interior cement tapestries. Coral dissolution may occur under meteoric conditions both in subaerial and burial environments. The tapestries at the bottom of some molds are covered by gray crystal silt pointing to dissolution in a near-surface burial setting influenced by meteoric-vadose waters. C: Some of the molds have been completely occluded by blocky burial cements. D: Still open molds are now filled with red internal sediment, in places overlying the gray crystal silt. The red sediment is unfossiliferous weakly laminated micrite that might represent a reworked karstic residual sediment. E: This stage is characterized by the formation of calcite-filled microcracks transecting both corals and packstone matrix during increasing burial conditions. F: Deep burial compaction and shearing produces angular fragments showing plastic bending and grain breakage. G: The last stage is characterized by tectonic fissures filled with clasts consisting of reworked corals and matrix. After Flügel und Flügel-Kahler (1997).
time are manifested by cement successions visible on small- and medium scales as well as by large-scale trends in mineralogy, cementation, and porosity due to shifts of diagenetic environments during rock history. Examples are shown in Fig. 7.15 and in Pl. 35.
ing diagenetic features to changing diagenetic environments (Schroeder 1986; Koch and Schorr 1986) and provides a more profound understanding of porosity evolution.
“Gone are the days when the carbonate petrographers were fully content in distinguishing cement A and cement B“ (Schroeder 1988). For a long time these terms were used in distinguishing first stage cement, generally regarded as early diagenetic, and second stage cement, considered to be late diagenetic in origin (Graf and Lamar 1950). Precise petrographical thin-section analyses and cathodoluminescence studies demonstrate that a simple two-fold classification gives only a very rough and often incorrect picture. Diagenetic stages are indicated by the superposition of different mineral phases, or by differences within cement minerals (Braithwaite 1993). Another pitfall in recognizing small-scale diagenetic successions is spatial variation within carbonate units (Pl. 35). Spatial variations increase in the course of diagenesis over time, resulting in a highly complex pattern of cementation, dissolution, and porosity development. The petrogenetogram, a time-space flow diagram ascribes vary-
7.5 Diagenetic Textures The most important processes occurring in subsurface carbonates are compaction and pressure solution. Burial of noncemented carbonate sediments under an increasing burden results in compaction and then in pressure solution. The specific criteria of these processes can be recognized on different scales including thin sections. During mechanical compaction, strain is concentrated at grain contacts. Grains may behave in either a brittle or a ductile fashion. Early cementation protects grains against compaction and pressure solution. Mechanical compaction eventually can continue on to chemical compaction if the grains begin to dissolve at their contacts. Chemical compaction in lithified limestones leads to pressure solution generating dissolution seams and stylolites. References on mechanical and chemical compaction are listed in Box 7.8.
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Compaction
Box 7.8. Important references on mechanical and chemical compaction of carbonate rocks. Overviews: Chilingarian 1989; Chilingarian and Wolf 1976; Coogan and Manus 1975; Hutcheon 1989; Choquette and James 1989; Shinn et al. 1976; Wanless 1983; Wolf and Chilingarian 1976 Experimental compaction: Bhattacharyya and Friedman 1983; Deelman 1975; Ebhardt 1968; Fruth et al.1966; Golightly and Hyde 1988; Molenaar and Venmans 1993; Neilson 1990; Pitman and Larese 1991; Robertson 1967; Shinn and Robbin 1983; Shinn et al. 1980 Compaction models: Bayer 1978; Carrio-Schafhauser et al.1990; Ortoleva et al. 1993; Ricken 1987, 1992; Stefaniuk and Mackowski 2000; Terzaghi 1940 Compaction studies: As-Saruri and Dietrich 1996; Bathurst 1987; Borre and Fabricius 1998; Chanda et al. 1977; Kendall 2000; Klotz 1990; Lasemi et al. 1990; Myers 1980; Myers and Hill 1983; Pray 1960; Saller 1996; Scholz 1973; Shinn et al. 1983; Trurnit 1980; Wanless 1979; Westphal and Munnecke 1997 Compaction measurement: Coogan 1970; Coogan and Manus 1975; Kahn 1956; Masson 1951; Nicolaides and Wallace 1997; Taylor 1950 Decompaction: Doglioni and Goldhammer 1990; Gnoli 2002; Lang 1989; Martire and Clari 1994; Meder 1987; Monaco et al. 1996; Perrier and Quiblier 1975; Steiger 1981; Wetzel and Aigner 1989 Stylolite studies: Bäuerle et al. 2000; Buxton and Sibley 1981; Friedel 1995; Friedman et al. 1981; Garrison and Kennedy 1977; Gillett 1983; Huber 1987; Luczynski 2001; Nelson 1981; Railsback 1993a, 1993b, 1993c, 1993d; Salameh and Zacher 1982; Steiger 1981 Stylolite models: DeBoer 1977; Deelman 1975; Merino et al. 1983; Park and Schot 1968; Weyl 1959 Compaction, stylolites and hydrocarbon reservoirs: Abu Dhabi National Research Foundation 1984; Amthor et al. 1988; Dunnington 1967; Hunt and Fitchen 1999; Leythhaeuser et al. 1995; Scholle 1977; Shinn et al. 1984; Smith 2000. Classification and terminology: Logan 1984; Logan and Semeniuk 1976; Wardlaw 1979.
7.5.1 Mechanical Processes: Compaction ‘Limestone compaction is an enigma’ (Shinn et al. 1977) Mechanical compaction is associated with dehydration, porosity reduction, and significant reduction of sediment thickness which may be reduced to one quarter of the original thickness. The initial compaction is characterized by three stages: (1) Deposition of sedimentary grains, dehydration, changes in packing density. Reduction of lime mud to 80–75% of the original porosity. (2) Compaction and dehydration of carbonate mud. The porosity is reduced to about 40%. Reorientation of grains. (3) Plastic deformation of grains. Micrite may be preserved in protected areas. Processes during stages 1 and 2 can occur within a burial depth of just one meter. Compaction of half of the original thickness and a loss of 50–60% of the initial porosity often needs a burial depth of at least a hundred meters. Dewatering is the main process in the mechanical compaction of muddy shelf carbonates and grain-supported sands, followed by pressure solution in deeper burial environments. In contrast, porosity loss of pelagic oozes is caused at burial depth in the upper tens of meters by a mechanical rearrangement of grains followed by grain crushing and pressure solution.
Fig. 7.16. Description of grain contact types according to Taylor (1950). Important types are point contacts, tangential contacts, concavo-convex contacts, and sutured contacts. Point contacts indicate initial compaction, tangential contacts increasing compaction, and sutured and concavo-convex contacts point to effects of pressure solution at grain contacts. Measurements taken along traverses assist in quantifying different grain contact types; the traverses should be vertical to the sedimentary bedding.
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Box 7.9. Compaction criteria seen in thin sections. Mud-supported micritic limestones (mudstones, wackestones) • elongated and merged peloids forming a clotted structure, • closely compressed shells, • shell breakage. Note that shells in experimentally compacted lime muds may not show breakage (Shinn et al. 1977; Shinn and Robbin 1983). Compaction effects on allochems decrease as the matrix percent increases (Fruth et al. 1966), • enrichment of skeletal grains in defined patches, • degree of burrow deformation (Ricken 1986), • distorted fenestral voids and desiccation structures, • the abundance of microfenestrae (Lasemi et al. 1990) decreased in experiments with increasing pressure. They describe the porosity evolution in micritic limestones and can be studied only in SEM. Microfenestrae have a maximum diameter of 1.5–15 µm. • thinned and wispy laminations, • irregular stringers of organic matter draping over rigid grains, forming wispy seams, • deformation of thin-walled organic microfossils within the sediment (Westphal and Munnecke 1997). Early cemented limestones contain spherical to slightly deformed microfossils, mechanically compacted carbonates exhibit flattened microfossils. Grain-supported limestones (grainstones, packstones) • plastic deformation of peloids (Pl. 36/2) and cortoids, • collapsing and telescoping of grains, • mechanical rearrangement of grains (Pl. 36/1, 3), • grain rotation indicated by rotated geopetal fabrics, • overpacking, • truncation of grains by adjacent grains, • spalled ooids (Pl. 36/5) and shells, e.g. foraminifera (Pl. 36/1), • broken and welded fossils (Pl. 36/1, 3), • broken micrite envelopes, • grain flattening (Pl. 36/5, 6), • curviplanar parallel grain contacts (Pl. 36/5), • stylolites.
Compaction criteria in thin sections Box 7.9 lists common compaction criteria for limestones. The intensity of the compaction of carbonate muds and carbonate sands is different as is the possibility of recognizing all the effects in thin sections. Description and measurement of compaction Measurements of compaction in thin sections require point-counter determinations of grain volume percentage, or the minus-cement porosity expressed by the frequency of cement or grain contact types. The resulting numbers are compared with inferred average porosities of modern carbonate sediments of similar composition. Grain contacts should be described following precise definitions (Taylor 1950; Fig. 7.16). Some cau-
Compaction-Decompaction
tion is necessary because the original shape of the grains may determine the type of grain contacts and not compaction effects. Several methods have been proposed for quantifying the degree of compaction: • Measuring the increase in grain volume percentage. This can be measured directly in thin sections of grainstones using point counter methods and expressed by Packing Density (Kahn 1956; Coogan 1970). Packing Density is the grain volume percentage as measured in thin section. A Compaction Index based on Packing Density was proposed by Coogan (1970) for oolitic grainstones, grainstones with aggregate grains, and skeletal grainstones. The Compaction Index is a measure of the degree of grain compaction, given as a percentage. It is derived from the grain volume measurement and adjusted for the original volume percentage of the uncompacted sediment according to grain type. • Measuring the increase in grain closeness. This is reflected by the increased number of grains that touch each other as compared to the original uncompacted sediment. It is based on grain boundary measurements, requiring measurement of grain-to-grain, grain-to-cement, or grain-to-pore contacts along traverse lines and is expressed by the Packing Index (Masson 1951; Coogan 1970). The Packing Index is determined from the ratio of the number of grains touching each other to the total number of grains in thin sections. The rationale behind the Packing Index is the idea that compacted grainstones should have more grains touching each other than uncompacted grainstones. This index is more quickly determined than packing density, but is less useful because it depends on the total number of grains counted, and is affected by grain shape and interparticle cementation, which may change grain contact configurations. • Measuring porosity and cements. In thin sections of grainstones, the degree of compaction is estimated by measuring the percentage of interparticle porosity and interparticle cements and comparing it with modern sediments of similar composition (Nicolaides and Wallace 1996). • Measuring the diameters of deformed burrows and comparing it with non-deformed circular burrow diameters (Ricken 1987; Gaillard and Jautee 1987). Open burrows of organisms living within the sediment will be deformed, whereas burrows filled with early sediment or cement will not be. Open burrows are deformed similarly to the neighboring sediment. • Measuring deformed fossils with originally circular cross sections, e.g. nautilid cephalopods.
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Controls on Compaction
Compaction and decompaction The degree of compaction can be calculated according to the Carbonate Compaction Law (Ricken 1987, 1992). The calculation requires quantitative data on the CaCO3 content and insoluble residue contents of the sample, rock porosity measurements, and compaction values derived from deformed burrows. An estimate of compaction percentages is necessary for estimating the precompaction thickness and the geometry of carbonate units by decompaction (Ricken 1985). Compaction values are used for decompaction calculations that can restore the original thickness of the limestone sequences. The evaluation of original thickness is necessary in understanding paleorelief geometries and the spatial relationships between reefs, platforms and basins (e.g. Lang 1989). Low compaction is common in reef boundstones, is medium in mudstones, high in packstones and marls, and often high in grainstones. Decompaction is commonly measured by evaluating of spheres deformed into ellipsoids. This technique uses deformed trace fossils that originally had a circular cross section (Wetzel and Aigner 1989), ooids (Coogan 1970) as well as originally spherical microfossils (Martire and Clari 1994; Monaco et al. 1996). Another common method for calculating thickness changes during burial compaction is based on porosity-depth curves (Perrier and Quiblier 1975; Doglioni and Goldhammer 1990). Algorithms relating compaction and decompaction of carbonates were developed by Goldhammer (1997). Controls on compaction Many dynamic, inherited, and also inhibiting variables influence the degree of compaction: Dynamic factors • overburden • subsurface temperature • pressure • duration of burial stress • pore pressure • pore-water chemistry: Pore water is of prime importance in chemical compaction but not in mechanical compaction (Weyl 1959). The greater the degree of undersaturation with respect to CaCO3, the lesser the amount of overburden needed to cause interparticle and intraparticle pressure solution. The greater the degree of oversaturation, the greater the amount of overburden needed to cause pressure solution. Inherited factors • sedimentary fabric • skeletal architecture: Microcrystalline skeletal grains
•
•
• •
(e.g. bryozoans) are more susceptible to mechanical and chemical diagenesis than single-crystal echinoderm grains (Meyers 1980) skeletal mineralogy: Apparently without major influence on inter- and intraparticle chemical compaction clay content: Clay-bearing rocks are often more susceptible to strong compaction than clay-free rocks. In limestone/marl sequences limestones are not or less compacted, but marls exhibit strong compaction (Munnecke 1990). The precise effects of clays (e.g. decreasing permeability to cementing pore waters, enhanced diffusion of dissolved CaCO3) require further studies grain size and sorting grain shape.
Inhibiting factors • preburial cementation; precompaction diagenesis of interparticle cements may inhibit compaction • preburial dolomitization • clay content (see inherited factors above). Compaction of carbonate muds and sands Compaction data on sedimentary carbonates have been gathered by Goldhammer (1997). In contrast to muddy siliciclastic beds, which are commonly highly compacted to fissile shales, muddy carbonates preserve sedimentary structures much better, because carbonate muds are far less compacted than clay-rich muds and because precompaction cementation is widespread, at least in shallow-marine carbonate. Modern shallow-marine lime mud, consisting of loosely arranged aggregates of micron-sized aragonite needles, has porosities in the range of 70–80% (Enos and Sawatsky 1981; Shinn and Robbin 1983). Squeezing experiments prove that these muds can not be condensed by mechanical compaction below a porosity of about 25%. Initial porosity loss is dominantly a result of dewatering. Grain stabilization is reached under burial conditions of 100 m or less. A fully compacted lime mud may then have the porosity of a compacted sand, but differs in the very low permeability. Mechanical compaction is particularly important in mud-dominated sediments as proved by studies of pelagic oozes and chalks (Scholle 1971; Neugebauer 1973; Schlanger and Douglas 1974; Scholle et al. 1983). Compaction of carbonate sands is favored by deep burial and high effective strength, a low geothermal gradient, a low rate of loading, low pore pressure, long burial under stress, and water-wet grain surfaces. Further factors are small average grain size, moderately to well sorted grains, close primary packing, stable min-
314
eralogy, and minimal early cementation and dolomitization. Grain-supported carbonates may be resistant to compaction down to 700 m.
7.5.2 Chemical Processes: Pressure Solution and Stylolitization Following mechanical compaction many sediments are subject to chemical compaction, expressed by pressure solution and the formation of stylolites and solu-
Pressure Solution and Stylolitization
tion seams that are often associated with fracturing structures (Logan 1984). Pressure solution is caused by load and/or tectonic stress. In more strongly cemented rocks solution takes place along extensive surfaces, forming stylolites. Dissolution often starts along a bedding surface. Mechanisms by which pressure solution occurs are dissolution in a thin solution film as result of compressive stress inside grain-to-grain boundaries (Weyl 1959; Rutter 1983) or dissolution at or just outside the rims of grain contacts resulting in an undercutting (Bathurst 1975; Tada and Siever 1986).
Plate 36 Mechanical Compaction and Deformation Structures Mechanical compaction is a process caused by sediment overburden and resulting in a general reduction of porosity and rock volume. The thickness of the overburden necessary to produce compaction structures is controversial. Overburden results in the mechanical failure of grains. Compaction is commonly followed by pressure solution recorded by stylolites and solution seams (see Pl. 37). The degree of mechanical compaction is inhibited or reduced by early precompaction cementation (–> 2) and can be studied in grainstones by cement/ grain ratios (Meyers 1980). 1 Mechanical and chemical compaction resulting in the formation of a diagenetic packstone exhibiting a condensed fabric after Logan and Semeniuk (1976). Overburden stress at grain contacts is accommodated by grain deformation in the form of ductile squeezing and grain fracturing (black arrows). Lime mud within the foraminifera is protected from compaction. Pressure solution is indicated by nonparallel multigrain sutured seams (microstylolites). The diagenetic packstone consists of worn, redeposited, and strongly compacted fusulinid foraminifera (F) and platy algae (A). Some matrix has been lost during compaction and a welded mass of skeletal grains remained. The sequential history of fusulinids includes: (1) Infilling of the tests with gray micrite (GM) and red micrite (RM) prior to compaction. Globular tests of the fusulinids suffered stronger breakage than spindle-shaped shells (top right). Deep burial environment. Early Permian (Rattendorf Formation): Carnic Alps, Austria. 2 Differential compaction. Ooid-peloid grainstone. The sample consists of alternating layers with ooids (center and uppermost layer), and layers with peloids, crustacean coprolites (Favreina, F) and ooids (below and above the ooid layer). Compaction is indicated by plastic bending (arrows) of peloids and Favreina grains. Compaction within the ooid layer was inhibited by early cementation and is only indicated by the close packing of the grains. The interparticle cement postdates compaction of grains, pointing to shallow subsurface burial cementation. Early Cretaceous (Purbeck facies, Berriasian): Subsurface, Bavaria, Germany. 3 Crushed fusulinid test, the victim of physical compaction caused by bioclast-to-lithoclast (L) contact. The whorls are broken and rearranged (arrows). Differences in the rigidity of lithoclasts and the fossil caused breakage and splittering of whorls associated with burial fracturing. Note the absence of stylolitization. The limestone is a sedimentary microbreccia generated by the erosion and deposition of platform-derived material. Early Permian (Tressdorf limestone): Carnic Alps, Austria. 4 Tectonically deformed limestone with branching scleractinian corals. Preservation stages of corals reflects several of the diagenetic and deformation stages shown in Fig. 7.15: Corallites with well-preserved septa (1); calcite within partially or totally dissolved corallites (2); infilling of red micritic sediment (RM) in hollow corallites (3). Corals were affected by shearing (4) and fracturing of the rock. The diagenetic pattern of aragonitic corals is often related to growth types: The open framework of branching corals favors dissolution, whereas massive corals are completely replaced by neomorphic spar. Cretaceous: Building stone of the Hittite wall, Bogazköy, central Anatolia. 5 Strong compaction. Ooid grainstone with distinctive compaction features: Flattened grains (FG) and parallel grain contacts (PGC). The nuclei of the ooids are dissolved. Deformation (arrows) of partially dissolved ooids suggests that compaction took place during burial. Late Permian (Main Dolomite, Zechstein): Subsurface, Debki well, northern Poland. 6 Shallow-burial compaction. Diagenetic pisoid grainstone with contorted grains. The concentration of calcite cement below the grains (arrows) points to precipitation in the meteoric-vadose zone subsequent to burial compaction. Another explanation is indicated by the laterally continuous cement garlands underneath the grains, which may mark stylolites filled with burial cements. Late Permian (Main Dolomite, Zechstein): Subsurface, Miloszewo well, northern Poland. –> 4: Flügel and Flügel-Kahler 1997; 5 and 6: Peryt 1986.
Plate 36: Compaction and Deformation
315
316
Pressure solution can increase the dissolution of calcite and is believed to be a significant source for the formation of porosity-occluding subsurface cements in deep burial settings. On the other hand, pressure solution may create conduits for fluids and open migration paths. Fracturing and pressure solution are prime factors for reservoir rocks in the Middle East.
Chemical Compaction
Stylolitization contributes to bulk volume reduction, resulting in a marked drop in the original thickness of carbonate units. Stylolitization modifies the primary sedimentary fabric of limestones dramatically and their description, therefore, requires a specific terminology focussing on the fabric and structure of the altered rocks.
Plate 37 Chemical Compaction: Pressure Solution (Stylolites and Seams) Chemical compaction is exemplified by pressure solution resulting in stylolites and solution seams formed under burial conditions. Stylolites are irregular, suture-like contacts produced by differential vertical movement under pressure accompanied by solution. They are marked by irregular and interlocking penetration on two sides: Columns, pits and tooth-like projections on one side fit into their counterparts on the other side (–> 2). Concentrations of insoluble residues along stylolite surfaces are common (–> 5). Solution seams are isolated or swarm-like partings characterized by thin seams, often with accumulations of insoluble residues. Pressure solution processes involve a strong reduction in rock bulk volume with a resultant loss in porosity caused by the occlusion of pores by late diagenetic subsurface cements (Wong and Oldershaw 1981). Volume loss caused by chemical compaction is facies-dependent. Stylolitization changes original depositional textures and produces new diagenetic fabrics. A specific terminology is necessary for describing the type of interfaces between adjacent particles, the structures produced by pressure solution and shear fracture, the fabrics developed by pressure solution, and the resulting rock type (Fig. 7.18). 1 High amplitude columnar stylolites in a peloidal bioclastic wackestone. Stylolite interfaces are partly covered by reactate calcite (arrow). SEM . Late Cretaceous (Mishrif Formation): Middle East. 2 Peaked high amplitude stylolites along a bedding plane separating laminated fine-grained peloid grainstone (LPG) and non-laminated medium-grained peloidal grainstone (PL). Both microfacies types originated in intertidal environments. The lower one is a microbial mat with abundant bacterial peloids (see Pl. 8/5). The upper microfacies corresponds to another mat formed by organically bound peloids and oncoids (O). The limestone was not mechanically compacted prior to stylolitization owing to early cementation by calcite. Stylolitization has dissolved a high amount of the limestone. Early Cretaceous (Purbeck facies, Berriasian): Subsurface of the Molasse zone, Bavaria, Germany. 3 Bedding-parallel irregular anastomosing sutured seams (microstylolite set) in lime mudstone. Note dark insoluble residue (stylocumulate, SC) along pressure-solution surfaces and calcite, representing styloreactate (material that has crystallized at stylolite interfaces, SR). Pressure solution took place subsequent to leaching of evaporite minerals (black arrows). Late Permian (Zechstein): Harz, Germany. 4 Pressure solution postdates compaction. Grain-to-grain pressure solution of conspicuously large, tangentially structured fractured ooids (TSO) exhibiting jagged surfaces. Nonparallel reticulate grain contacts. Note dark stylocumulate (clay residue, calcite fragments and dolomite). Dolomitization occurs within the stylocumulate, but also in ooids as selective replacement. Arrows point to dolomite crystals. Cambrian: Takkanawa Fall, Yoho Valley, British Columbia, Canada. 5 Widely spaced high-amplitude columnar stylolites in a burrowed ostracod wackestone. The dark stylocumulate consists of clay and pyrite. The limestone has suffered no compaction prior to deep-burial pressure solution. Silurian: Valentintörl, Central Carnic Alps, Austria. 6 Non-sutured wispy parallel sets of dissolution seams producing a stylolaminated fabric. This fabric tends to develop when the clay content is high, larger than at least 10%. Arrows point to individual seam sets. Note that very similar features have been produced by mechanical experimental compaction (Shinn and Robbin 1983). Lime mudstone. Jurassic: Peloponnesus, Greece. 7 Stylobreccia. Sutured reticulate multi-grain fabric (Wanless) or condensed fabric (Logan and Semeniuk). Idens (bodies that behave as homogeneous entities under physical or chemical conditions) are densely packed and bound by irregular, anastomosing microstylolites (white arrows). The black arrow points to dissolution of a miliolid foraminiferal shell. RC: Reactate calcite formed by pressure solution. Subsurface of the Molasse zone, Bavaria, Germany. 8 ‘Horse tail’ structure characterized by lateral splitting of seams (arrows). Bedded non-sutured wispy seams (Wanless). Irregular anastomosing sets (Logan and Semeniuk). Estimates of the minimum local extent of chemical compaction can be made by determining the difference in thickness between a solution seam and the laterally equivalent wedge of diagenetically laminate carbonate (horsetails: Kendall 2000). Cretaceous: Egypt.
Plate 37: Stylolites and Seams
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Fig. 7.17. Stylolites on a microfacies scale. Stylolites are serrated boundaries between units. The boundary usually has an accumulation of clay, oxides, and/or organic matter. Lowermost Cretaceous carbonate ramp: Subsurface, southern Bavaria, Germany. Scale is 1 mm.
Criteria of pressure solution in carbonate rocks Chemical compaction (Lloyd 1977) is recorded by • dissolution seams (horse tails, wispy stylolites, microstylolites) characterized by surfaces that can be followed across several grains; common in fine-grained and clay-rich limestones, • stylolites characterized by greater amplitudes than in dissolution seams; common in more coarse-grained limestones or at lithologic boundaries, • stylolites often cross the whole rock, intersecting grains, matrix and cement, • stylolites often have a zigzag form, due to differences in the solubility of different components (e.g.
Pressure Solution
skeletal grains may be less soluble than micrite matrix), • the less soluble parts of three-dimensional ‘fingers’, extending into the more soluble parts, • amplitudes of stylolites may range from < 1mm to > 1 m. The amplitudes indicate the minimum thickness of the dissolved material, • insoluble material (often clay, hydrocarbons) becomes concentrated along the stylolite surface. This material does not appear to reprecipitate in the immediate vicinity of its origin, implying transport by fluids, • pressure solution-derived cements, • inter- or intragranular dissolution at grain contacts causing sutured boundaries. Terminology and classification of carbonate rocks affected by chemical compaction Diagenetic features of carbonate rocks, even if they are striking, are only rarely recorded in names of limestone types. Exceptions are the modification of the Dunham classification proposed by Wright (1992; Sect. 8.4.1) and the consideration of pressure solution structures in the classification of Logan and Semeniuk (1976). Fig. 7.18 summarizes the main criteria used by Logan and Semeniuk and by Wanless (1979).
>>> Fig. 7.18. Terminology of pressure solution features. Most of these categories can be derived from thin-section studies (see Pl. 37). A - Styles of pressure-solution features after Wanless (1979) modified by Choquette and James (1987). The non-sutured type is characterized by microstylolites, microstylolite swarms, clay seams and grain contact sutures; it forms in limestones that contain a significant amount of insoluble material (clay, silt, organic matter). The sutured type is typified by stylolites and grain-contact sutures and forms in limestones with very low amounts of insoluble residual material. Bedding-parallel sutured seams correspond to stylolites of different amplitudes, which may occur as single structures or in swarms. Reticulate systems of anastomosing microstylolites (fitted fabrics, see B) and non-sutured seams occur in limestones made up of rigid entities in grain-to-grain contact or surrounded by more soluble sediment. B – Classification of pressure solution features by Logan and Semeniuk (1976). The classification concentrates on the morphology of stylolites (columnar, Pl. 37/1, 5; high amplitude - Pl. 37/2 and low amplitude - Pl. 37/3 - types, irregular, hummocky or smooth) and differences in stylolite sets (parallel, irregular anastomosing, Pl. 37/3; conjugate - Pl. 37/6), and focuses on fabric types caused by pressure solution and structural types that can be recognized on micro- to macroscales. Note that in rare cases stylolites can also be formed by mechanical compaction without pressure solution (Braithwaite 1986). Fabric types are differentiated according to the behavior of idens (bodies that behave as statistically homogeneous entities under physical and chemical changes). Iden-supported fabric: Idens form a self-supporting frame and are in contact at grain points. Condensed fabric: Idens are very densely packed, grains contact over a surface. Fitted fabric: Idens produce a puzzle-like fabric. They are surrounded by solution seams and are in contact with their neighbors all along their margins. Stylocumulate/reaction support: Larger idens are supported by fine-grained stylocumulate (insoluble residue accumulated along a pressure-solution surface). Diagenetic structures: Stylobedding: Pseudo-bedding caused by parallel pressure solution that may be parallel to the original horizontal layering. Stylolaminated: Laminated appearance due to swarms of parallel solution seams or low-amplitude stylolites. Stylonodular: Nodules and lenses of limestone idens separated by broad stylolite swarms. The original micritic matrix is more or less dissolved. Stylomottled: Characterized by patchy enrichment of insoluble stylocumulate and reactate (material which has grown at solution interfaces, e.g. calcite, dolomite, or quartz), may be also caused by patchy enrichment of marls due to inhomogeneous pressure solution. Stylobreccoid: Originates from selective pressure solution (Pl. 37/7).
Pressure Solution: Terminology
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Fig. 7.18. Terminology of pressure solution features. Most of these categories can be derived from thin-section studies (see Pl. 37). Explanation on the left.
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The terms introduced by Logan and Semeniuk for the description of rocks suffering from pressure solution were included by Logan (1984) in a more rigorous classification that also considers deformation structures caused by pressure response to burial and/or tectonism. Carbonate rocks affected by compressional stress are accommodated by compaction, pressure solution, and shear fracture. The main response to tensional stress is the formation of strain cavities - represented by veins, vugs, and inter-fragment voids within breccias. Compressional and tensional stress is often associated with dolomitization, silicification, and other mineralization processes. Controls on chemical compaction Chemical compaction is controlled in a complex manner and depends on • burial depth with attendant lithostatic load. Pressure solution becomes effective at various depths. Many authors favor burial of several hundreds to thousands of meters, but stylolite formation has also been recorded after only some tens to hundreds of meters. The paradigma of 900 m overload necessary for the formation of stylolites (Dunnington 1967) is out of date. • tectonic stress leading to tectonically controlled formation of stylolites (Bathurst 1984), • vertical stress caused by overburden pressure resulting in preferential solution of the upper and lower surfaces of grains, via solution films, • the chemistry of pore fluids. Mg-poor meteoric pore waters will enhance pressure solution, • the formation of stylolites can be retarded and prevented by the presence of oil or organic films between and on grains, • pore fluid pressure in overpressing settings may retard or stop pressure solution, • original mineralogy controls the early compactional history. Aragonite mineralogy at times of burial (shallow-marine carbonate muds) will enhance chemical compaction; primary calcite mineralogy (e.g. deep-marine carbonate oozes and chalk) will retard pressure solution, • insoluble residues (clay, organic matter) seems to enhance pressure solution in fine-grained carbonates (Oldershaw and Scoffin 1967, Sassen et al. 1987), • grain size and grain surface properties, • early dolomitization can inhibit or preclude stylolitization, • clay minerals; the presence or absence of minor amounts of clay minerals appear to be a major control on whether or not pressure solution occurs in carbonate rocks (de Boer 1977).
Pressure Soultion: Significance
Box 7.10. Significance of compaction and pressure solution. Thickness and paleotopography of sedimentary units: Stylolitization results in a considerable reduction of the thickness of a sequence. Differential compaction may control seafloor relief and reef geometry (Anderson et al. 1996) and reduces the thickness of adjacent facies units, e.g. reef and platform margins (Matyszkiewicz 1999). Burial cements: Pressure solution increases the dissolution of calcite (Maliva and Siever, 1989). The supply of CaCO3 from pressure solution along stylolites is a major source of late diagenetic cement and may or may not (Kreutzberger and Peacor 1988) affect also phyllosilicates within carbonate rocks. Dolomitization: Transformation of clay minerals by burial diagenesis could be a source of magnesium and dolomitization (McHargue and Price 1982; Sternbach and Friedman 1984). Economic importance: Compaction and stylolitization reduce porosity depending on the degree of precompaction cementation and depending on the textural rock type. Stylolites may enhance and inhibit reservoir quality (Abu Dhabi National Reservoir Research 1984). Stylolite surfaces may be important permeability barriers, influencing the flow in aquifers and petroleum reservoirs. Grain-to-grain pressure solution is the most detrimental process for permeability (Budd 2002), followed by cementation. However, stylolites can act as conduits for fluids (Braithwaite 1989), and permeability may be created by veining activating the connectivity of stylolites (Smith 2000).
Pressure solution and rock types Pressure solution in grain-supported carbonate rocks with minor cementation starts with solution at grainto-grain contacts. Solution affects the more soluble grains causing a ranking of grain solution. Pressure solution subsequent to cementation results in the formation of a fabric that distributes pressure laterally, resulting in distinct stylolites or in a fitted fabric (Buxton and Sibley 1981). Grain-rich carbonates can be stylolitized under shallow burial of not more than 30 m (Railsback 1993). Lime mudstones, wackestones, packstones, and weakly cemented grainstones exhibit preferential solution seams that may indicate diagenetic bedding processes (Ricken 1986). Solution seams can be pre- and post-burial in origin (Bose et al. 1996). Solution resistance of grains? Differences in solution resistance of different grain and matrix types have been observed by various authors (e.g. Trurnit 1968; Logan and Semeniuk 1976; Steiger 1981; Huber 1987). Sandstone and dolomite extraclasts are more resistant
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Alteration and Recrystallization
Box 7.11. Terms introduced by Folk (1965) in order to differentiate neomorphic processes. Aggrading neomorphism: A kind of coalescive neomorphism or recrystallization in which the crystal size increases and finer crystal mosaics are replaced by coarser crystal mosaics of the same mineral or its polymorph without the intermediate formation of visible porosity (Bathurst 1975). Occurs chiefly in fine-grained limestones and results in microsparitic patches, laminae and beds. See Morse and Mackenzie (1990) for a discussion of aggrading neomorphism. Pl. 38/2, 3. Coalescive neomorphism: Larger crystals grow at the expense of smaller crystals or small crystals grow within a larger crystal. Degrading neomorphism: A kind of coalescive neomorphism in which the crystal size decreases; smaller crystals grow within a larger crystal. Not common, but may occur under stressed or low-metamorphic conditions.
to solution than carbonate grains. Carbonate grains exhibit strong variabilities with regard to pressure solution potential. The only feature common to all case studies is that echinoderm grains seem to be the most resistant grains, and micritic particles (e.g. peloids) and argillaceous micritic matrix are the least resistant parts of carbonate rocks.
7.5.3 Significance of Compaction and Pressure Solution
quartz), and recrystallization (changes in crystal size, shape and crystal lattice without a change in mineralogy). The last term is often used rather loosely, since it is not always possible to demonstrate whether or not a mineralogical change has occurred. Neomorphic processes lead to crystal enlargement or diminution (Box 7.11). The transformation of aragonite and High-Mg calcite grains and mud to Low-Mg calcite is one of the most important processes in carbonate diagenesis because it controls the ultimate petrophysical properties of limestones and their geochemical composition (particularly stable isotopes and Sr content: Al-Asam and Veizer 1986a, 1986b; Banner 1995; Maliva 1998). Replacement of aragonite by calcite involves gradual dissolution of the original mineral and precipitation of calcite, so that relicts of original shell or cement structures may be preserved in the neomorphic calcite.
7.6 1 Recrystallized Carbonate Rocks: What to do? Recrystallized limestones are predominantly characterized by changes in the size, shape and arrangement of crystals, and by the obliteration of the original depositional textures and constituents. Recrystallization of micritic limestones may be controlled by the clay content of limestones; a clay content > 2% seems to inhibit or prevent crystal coarsening (Bausch 1968). Enlargement of micrite may create microspar (see
Box 7.10 lists important effects of the mechanical and chemical compaction of carbonate rocks during burial as well as the influence of dolomitization. It also discusses the economic importance of compaction and pressure solution.
7.6 Neomorphic Processes: Alteration and Recrystallization Carbonate sediments and carbonate rocks are affected by diagenetic processes that alter mineralogical composition and/or crystals and crystal fabrics. The term neomorphism (Folk 1965) relates to transformations of minerals taking place in the presence of water and includes processes of replacement (the dissolution of one mineral and simultaneous formation of another mineral, e.g. silicification of carbonates), inversion (the replacement of a mineral by its polymorph; e.g. aragonite –> calcite (calcitization); opal-A –> opal-CT –>
Fig. 7.19. Recrystallized coral limestone. Arrows point to centers of former corallites of a colonial coral. Note faint radial structures representing the recrystallized coral septa. Aggrading neomorphism has replaced cement crystals between coral calices by a medium to coarse-grained inequigranular, hypidiotopic calcite fabric. Crystals were enlarged at the expense of other crystals (note left side of the figure in particular). Jurassic: Argentina. Scale is 5 mm.
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Recrystallized Limestones and Marbles
Plate 38 Recrystallized Limestones and Metamorphic Carbonates (Marbles) The original depositional microfacies and early diagenetic textures of limestones are often altered (–> 1) or destroyed by recrystallization and processes acting during burial diagenesis and thermal history in relation to the ups and downs of the rocks. Recrystallized carbonate rocks (–> 2-3) should not be disregarded in facies analyses, even though depositional criteria are obliterated or completely destroyed. Recrystallization depends on clay content (> 2 % may inhibit recrystallization) and can create microspar (see Pl. 6/2) and pseudospar, the latter with crystals measuring tens to even hundreds of microns in diameter. Burial diagenesis and changes in temperature, pressure, and volatile content will favor the formation of carbonate rocks with crystal sizes up to several tens of millimeters. Some of these rocks already correspond to marbles, showing a fine- to coarse-grained granoblastic texture. Marble (–> 4-8; Sect. 7.9) is a metamorphic rock consisting predominantly of fine- to coarse-grained, dynamically recrystallized calcite and/or dolomite, usually with a granoblastic, crystalline texture. Differentiation of marbles in thin section requires the description of mineralogical composition (carbonate and non-carbonate minerals), grain shape, fabric (size and mutual relations of crystals), deformation features like frequency, bending and orientation of twin lamellae with respect to a reference plane (–> 4), and granulation textures (caused by mechanical breakage along gliding planes). Recrystallized limestones 1 Tectonically elongated fusulinid foraminifera (Quasifusulina). Low-grade tectonic deformation can remodel the shape of fossils. The tests were originally deposited parallel to each other and later stretched along their length axis. Fossils of known geometric parameters deformed in this way are important tools in reconstructing shear processes. Permian: Bükk Mountains, northern Hungary. 2 Neomorphized limestone. Note the differences in crystal size and preservation of primarily calcitic micrite clasts (MC) and the aragonitic dasyclad alga (DA). The skeleton of the alga has been replaced by calcite. Arrows point to twin lamellae developed within the segmented algal skeleton. The former micritic matrix is altered by aggrading recrystallization. Early Carboniferous: Anhui, southern China. 3 Neomorphized recrystallized limestone. Dark areas are relicts of former micritic grains (1). Former cement crystals in interparticle (2) and intraparticle pores suffered aggrading neomorphism, indicated by the coarsening of calcite crystals. The outline of the structure marked by (3) suggests a foraminiferal shell whose central part has been replaced by coarse calcite. Common diagnostic features of the neomorphic sparite (pseudospar) are crystals of conspicuously different sizes, predominantly curved interfaces between crystals, and large crystals in association with smaller crystals. Early Carboniferous: Anhui, southern China. Marbles 4 Inequigranular coarsely crystalline calcite marble with silicate impurities. Xenotopic and poikilotopic fabric (larger calcite crystals enclose smaller silicate mineral crystals: black arrows). White arrows point to an ‘enfacial junction’ (a joining point of three intercrystalline boundaries where one of the three angles is 180°). The frequency of this type of triple junctions is used in distinguishing neomorphic and metamorphic spar and carbonate cement. Note the abundance of thin twin lamellae indicating deformation at relatively low temperatures (Ferrill 1998). As shown by the varying directions of the twin lamellae, and strong differences in grain size, the marble is a protomylonite, caused by multiple deformations. Paleozoic: Roman quarry Tentschach, Carinthia, Austria. 5 Inequigranular medium to coarsely crystalline marble, a product of intensive strain recrystallization. Xenotopic fabric. Note smaller crystals (size about 0.16 mm, arrows) between larger crystals (maximum size 1.0 mm), indicating grain diminution. Marble used in Roman buildings. Hergla, Gulf of Hammamet, eastern Tunisia. 6 Equigranular finely crystalline marble consisting of anhedral calcite grains, probably resulting from cataclastic granulation. Arrow points to a former shell now consisting of twinned and bent calcite crystals. Twinning is often indicated by a conspicuous strippy pattern. Ghosts of fossils (especially shells, echinoderms) may be preserved even in highly metamorphic carbonates. Jurassic (Chemtou marble): Roman quarries, Chemtou, Tunisia. 7 Inequigranular coarsely to very coarsely crystalline marble. Xenotopic fabric consisting of calcite crystals exhibiting no crystal faces bounding the mineral grains (anhedral crystal shape). Size between 0.9 and 2.5 mm. Note different directions of twin lamellae indicating dynamic recrystallization. Mesozoic marble: Sterzing, South Tyrol, northern Italy. 8 Same thin section as –> 7, crossed nicols. Note extensive twin development. Twin planes are closely spaced. The black arrow points to multiple set twin lamellae, the white arrow to bent and displaced twin lamellae. –> 1: Kahler 1988
Plate 38: Recrystallized Limestones and Marbles
323
324
Sparite
Fig. 7.20. Crystallization fabrics of recrystallized carbonates and marbles can be described according to equant or nonequant crystal size and crystal fabric depending on crystal shape (see Box 7.12 for definitions). From Friedman (1965).
Sect. 4.1.3) and continuing diagenesis may produce pseudospar with crystals measuring tens to even hundreds of microns in diameter (Sect. 7.7). Burial diagenesis and changes in temperature, pressure, and volatile content will favor the formation of carbonate rocks with crystal sizes up to several tens of millimeters. Some of these rocks already correspond to marbles, characterized by fine- to coarse-grained granoblastic textures (Fig. 7.20). Recrystallized carbonate rocks are not loved by geologists interested in facies analysis. These rocks show
Box 7.12. Descriptive terms for crystallization textures and fabrics after Friedman (1964). Compare Fig. 7.20. The scheme can also be used to describe calcite and dolomite marbles. Crystallization textures: Refer to the shape of mineral crystals and the type of crystal faces at crystal boundaries. Descriptive terms are Anhedral: Characterized by the absence of crystal facing bounding the mineral grains. Pl. 38/3. Subhedral: Characterized by partly developed crystal faces. Euhedral: Characterized by crystals that are bounded by crystal faces. Crystallization fabrics: Refer to the size and mutual relations of crystals that can be differentiated into Equigranular fabrics: Consist of crystals of approximately the same size (e.g. Pl. 38/6). Fabrics subdivided into Xenotopic: Predominantly anhedral crystals. Hypidiotopic: Predominantly subhedral crystals. Idiotopic: Predominantly euhedral crystals. Inequigranular fabrics: Consist of crystals of various sizes (e.g. Pl. 38/4, 5, 7). Fabrics are subdivided into xenotopic hypidiotopic idiotopic. Each of these inequigranular fabrics is in turn subdivided into Porphyrotopic: Larger crystals are enclosed in a fine-grained matrix. Poikilotopic: Crystals are of more than one size; larger crystals enclose smaller crystals of another mineral.
only a few characteristics that can be used in paleoenvironmental interpretations and they are poor in criteria allowing a reproducible classification. What to do? If no relicts of the primary depositional textures are preserved, a classification can be based only on the attributes of the diagenetic calcite. These features include crystal sizes, crystal shapes, mutual relationships between crystals, types of boundaries between crystals, and crystal foundations.
7.6.2 How to Describe Recrystallized Carbonate Rocks? Textures and fabrics of recrystallized limestones and dolomites as well as carbonate cements can be described using the classification developed by Friedman (1964, 1965) and summarized in Box 7.12. Hibbard (1994) drew attention to the limited value of the terms anhedral, subhedral, and euhedral where crystal morphology is considered an expression of the conditions under which crystals and rocks are formed.
7.7 Sparite: Recrystallization Product or Carbonate Cement? The non-genetical term sparite (Fig. 7.21) designates a fabric consisting of calcite spar. Genetically, sparite is carbonate cement (corresponding to orthosparite) or a product of recrystallization. Distinguishing the products of aggrading neomorphism from sparite representing carbonate cement is often difficult. Criteria have been discussed by Folk (1965) and Bathurst (1975). Sparite representing cement occurs within pores. Boundaries between sparite and pore limits are sharp. Crystal boundaries are straight and have enfacial junctions. Neomorphic fabrics often exhibit irregular crystals with curved and embayed boundaries, variable crystal size with remnants of micrite, and the presence of grains
Sparite
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Fig. 7.21. Sparite (sparry calcite) represents carbonate cement or products of neomorphic changes. The former have been termed orthosparite, the latter pseudosparite. The thin-section photograph displays various early and late diagenetic stages of a lithoclast/bioclast rudstone deposited near a reef. The numbers indicate syn- and postdepositional diagenetic alterations affecting the skeleton of a dasyclad green alga (Macroporella sp.) and its infillings as well as the carbonate precipitated in interparticle pore spaces. 1: Early diagenesis starts with the micritization of the walls between the algal branches (preserved as pores 2). 2: Infilling of algal pores with micrite (now preserved as microspar) and small peloids. 3: Microbial encrustation of the algal surface. 4: Transformation of the originally aragonitic pore walls to calcite. 5: Infilling of the algal interior by sedimentary grains, some of which are covered by thin marine cement crusts composed of bladed calcite crystals. 6: These grains are incorporated within originally micritic microbial crusts formed in the algal interior and now preserved as microspar. 7: The alga is overgrown by a complex biogenic crust. 8: The interior of the alga is partly filled with larger bioclasts. 9: The boundary between the infilling and the remaining pore space is marked by an irregular microbial crust. 10: Large interparticle spaces between (dark) reef-derived lithoclasts are occluded by abundant shallow-marine radiaxial calcite cements (orthosparite). 11: Late diagenetic processes are responsible for the formation of coarse calcite spar (pseudosparite), probably of burial origin. 12: Recrystallization causes coarsening of the walls of algal pores. 13: Tectonic shear is indicated by breaking and dislocation of the radiaxial cement crusts and by 14: a thin calcite-filled veinlet crossing cements and alga. 15: Neomorphic alterations enlarge crystals within the calcite cements. Sparite-producing diagenesis took place under marine near-subsurface conditions (points 4, 5, 10) and burial conditions (points 11, 15). Middle Permian: Straza near Bled, Slovenia. Scale is 2 mm.
apparently ‘floating’ within sparite. The boundary between sparitic and micritic areas is often poorly defined and can crosscut micritic grains. Larger crystals are often irregularly intergrown. A common neomorphic fabric is microspar characterized by a crystal size between about 5 and 30 µm (see Sect. 4.1.3). Coarse neomorphic fabrics originating from aggrading recrystallization and consisting of calcite crystals larger than 30 µm or so are called pseudosparite (Folk 1965). Metasparite (Kovacs 1987) designates a similar fabric that exhibits indications of metamorphic processes (e.g. granoblastic texture, complex twinning).
7.8 Dolomitization and Dedolomitization The first part of this chapter deals with the descriptive criteria of dolomites seen in microfacies thin sections. The lack of undisputed modern analogues for ancient dolomites is the core of the ‘dolomite problem’ and the reason for several competing dolomitization models that are briefly summarized in the second part of this chapter. The third part discusses dedolomitization phenonema.
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Dolomite and Dedolomite
Plate 39 Dolomite Fabrics, Dolomite Types (Saddle Dolomite) and Dedolomitization The basic terminology of dolomite fabrics used in thin sections considers crystal shape, the mutual relationships of crystals and uniform or non-uniform size distribution (–> 1-3). Saddle dolomite (–> 5) or ‘baroque dolomite’ is a specific dolomite type, commonly interpreted as having formed under deep burial or hydrothermal conditions. Dedolomitization refers to the early or late diagenetic replacement of dolomite by calcite (–> 4, 6-8). Criteria of dolomite fabrics The fabric of dolomite rocks varies between two end-members: Limestones with scattered dolomite rhombs (–> 2) and sucrosic mosaic dolomites (–> 3). 1 Criteria: Inequigranular; non-rhombic crystals; tightly packed anhedral crystals; xenotopic fabric; intercrystalline boundaries usually curved, lobate, irregular; preserved crystal junctions rare; grain size between 0.4 and 1.5 mm. Classification: Friedman – Decimicron- to millimeter-sized inequigranular, xenotopic dolomite: Gregg and Sibley – Polymodal nonplanar matrix replacement dolomite; Randazzo and Zachos – Inaequigranular mosaic dolomite. Replacement dolomite. Crystal size scale (Folk) – Coarsely to very coarsely crystalline. Black spots are artifacts. Interpretation: Late burial replacement dolomitization is indicated by the amoeboid fabric and the serrate crystal boundaries. 2 The sample is characterized by clear euhedral and subhedral dolomite crystals (0.2–0.4 mm) exhibiting cloudy centers and clear rims. Cloudy centers can be caused by calcite relicts, non-carbonate or fluid inclusions. The dolomite crystals exhibit straight compromise boundaries. Descriptive criteria: Equigranular; rhombic crystals; crystal-supported; idiotopic fabric. Intercrystalline areas are still microcrystalline calcite. Fenestral voids (FV) are filled with recrystallized fibrous calcite cement. Classification: Friedman – Decimicron-sized idiotopic dolomite. Gregg and Sibley – Unimodal planar partial matrix replacement dolomite. Randazzo and Zachos – Equigranular sieve mosaic dolomite. Size scale – Medium to coarsely crystalline. Interpretation: Dolomite crystals postdate shallow-marine calcite cements of fenestral tidal limestones. Missing evidence of compaction prior to dolomite crystal growth and the restriction of the crystals to fenestral boundaries indicate replacement dolomitization of tidal sediment in a shallow burial setting. 3 Criteria: Inequigranular; non-rhombic crystals; tightly packed anhedral and subhedral crystals; hypidiotopic fabric; intercrystalline boundaries lobate and straight; some crystal-face junctions preserved; grain size between 0.1 and 0.4 mm. Classification: Friedman – Centimicron-sized inequigranular, hypidiotopic dolomite; Gregg and Sibley – Polymodal nonplanar matrix replacement dolomite; Randazzo and Zachos – Inequigranular mosaic dolomite. Size scale – medium to coarsely crystalline. These fine- to medium-grained mosaic dolomites formed by replacement of fine-grained calcite matrix. The hypidiotopic fabric indicates recrystallization of burial dolomites. –> 1–3: Triassic: Roman mosaic stones, Hemmaberg near Globasnitz, Carinthia, Austria. Dedolomitization (–> 4, 6–8) and Saddle Dolomite (–> 5) 4 Rhombohedral dolomite crystals have been dissolved. Molds were filled with equigranular microcrystalline Low-Mg calcite. Preferred solution of the core zones of dolomite rhombohedrons is a common phenomenon, because these zones often are richer in Ca than the periphery of the crystals (Randazzo and Cook 1987). SEM microphotograph. Late Jurassic: Southern Franconian Alb, Germany. 5 Late diagenetic ‘saddle dolomite’ cement crystals within a vug characterized by curved crystal faces (black arrow). The zonar dolomite has been internally corroded by low-saline formation-water migrating along rhombohedral cleavage planes and causing burial dedolomitization and the formation of intracrystalline pores (white arrows). VP: Vuggy porosity. Fluid inclusion analysis provides direct information about fluid movement and the timing of hydrocarbon migration. Saddle dolomites are characterized by large, mm-sized crystals, curved crystal faces and cleavages, undulose extinction, commonly ferroan composition, often fluid inclusions that make the crystals appear cloudy in thin sections and pearly in hand specimens. Saddle dolomites are commonly associated with sulphide mineralization, hydrothermal activity, and the presence of hydrocarbons at relatively deep burial depths. 6 Burial dedolomitization of a late diagenetic dolomite. Calcitization is restricted to the inner Ca-rich zones of former zonar dolomite crystals (white arrows). 7 Near-surface early meteoric-vadose dedolomitization: Solution-enlarged molds of the former dolomite crystals are filled with meteoric calcite cements (arrows). 8 Near-surface early meteoric-vadose dedolomitization. Crystals molds are geopetally infilled by gray vadose crystal silt (arrows) and granular meteoric cement. –> 5–8: Jurassic (Kimmeridgian): Subsurface, TB-3 Saulgau well, southwestern Germany. –> 1–3: E. Flügel and Ch. Flügel 1997; 4–7: Meder 1987
Plate 39: Dolomite and Dedolomite
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Reviews summarizing the state of the art of the dolomite problem and the enigma concerning the uneven distribution of dolomite rocks during the Precambrian and the Phanerozoic were given by Zenger et al. (1980), Land (1983), Machel and Mountjoy (1986), Hardie (1987), Given and Wilkinson (1987), Shukla and Baker (1988), Morrow (1990a, 1990b), Purser et al. (1994), Sun (1994), Burns et al. (2000), and Warren (2000).
7.8.1 Descriptive Criteria and Terminology of Dolomite Fabrics Similar to limestones, dolomite rocks can be grouped according to mineralogical and chemical composition, and to textural criteria. Various terms are in use for describing the amount of dolomite contents in carbonate rocks. The definition of these terms is not uniform and sometimes vague: • The use of the term dolomitized may (a) simply indicate that dolomitization has converted limestones wholly or partly to dolomite rocks or dolomitic limestone, or more specifically, (b) that the rock contains more than 10% replacement dolomite (Folk 1959). • Dolomitic refers to (a) a dolomite-bearing or dolomite-containing rock that contains different amounts of the mineral dolomite. Dolomitic limestones yield a conspicuous amount of the mineral dolomite, but calcite is more important. The opposite is a calcareous dolomite (see Fig. 3.1). The ill-defined term magnesian limestone designates a rock composed of mixtures of calcite and dolomite or a limestone with some MgO but no dolomite. Dolostones are classified genetically into (1) syngenetic dolostone (penecontemporaneously formed within the depositional environment), (2) diagenetic dolostone (formed by replacement of carbonate sediments or limestones during of following consolidation), and (3) epigenetic dolostone (formed by localized replacement along postdepositional faults and fractures).
7.8.1.1 Thin-Section Description and Terminology of Dolomite Rocks In thin sections, dolomites may or may not exhibit their original depositional and compositional textures. Calcareous dolostones and dolostones having preserved the original depositional components and textures are classified according to textural composition, crystal textures, and crystal fabrics. The shape and the size of do-
Dolomite Fabrics
lomite crystals as well as their mutual relationships are important for classifying those dolomite rocks whose original textures have been obscured or obliterated (Friedman and Sander 1967). Most dolomite nomenclature systems are descriptive and based on compositional grouping modified by textural terms (Folk 1959; Leighton and Pendexter 1962; Powers 1962; Bissell and Chilingar 1967). A ‘coarsely crystalline oolitic dolomite’ (Folk 1959), for example, is a dolomite characterized by ooids within a matrix consisting of dolomite crystals 0.25 to 1 mm in size. The basic terminology used in describing dolomite rocks was developed by Friedman (1965; see Box 7.12) and expanded by Randazzo and Zachos (1983) and Sibley and Gregg (1987). The use of crystal size scales allows more extensive differentiation of dolomite fabrics. Different crystal size limits are used by Friedman (1965), Randazzo and Zachos (1984) and Wright (1992). The latter coined the terms dolomicrostone (crystal size < 4 µm), dolomicrosparstone (crystal size 4–10 µm) and dolosparstone (crystal size > 10 µm). Crystal size distributions carry information on nucleation and growth (Sibley et al. 1993). The nomenclature proposed by Randazzo and Zachos (Fig. 7.22A) offers a possibility for recognizing dolostone types originating from ‘homogeneous’ and ‘heterogeneous’ dolomitization. Practical use of this classification demands answers to the following questions: • Crystal size: Very fine, fine, medium, or coarse crystalline? • Crystal size texture: Equigranular, inequigranular (unimodal, polymodal) or extremely fine (aphanotopic)? • Fabric: peloidal, mosaic, porphyrotopic, poikilotopic? • Crystal distribution: Tightly packed, loosely packed, isolated patches, crystals isolated or floating? • Crystal shape: Anhedral, euhedral or subhedral? Following the Randazzo and Zachos classification an ‘inequigranular, idiotopic, floating-rhomb porphyrotopic dolomite fabric’ is a fabric composed of predominantly euhedral, isolated crystals of different size, floating within a fine-grained matrix. An ‘equigranular, microxenotopic, sutured mosaic fabric’ is a fabric consisting of fine, tightly packed, predominantly anhedral crystals of equant size. Further examples for the use of both classifications are shown on Pl. 39 and Pl. 40.
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Dolomite Fabrics
A
EQUIGRANULAR: Unimodal crystal size Peloidal
Mosaic
Sharp to diffuse clotting of crystals of uniform size
Sutured: Tightly packed anhedral crystals with no or little intercrystalline porosity
Sieve: Loosely packed anhedral to euhedral crystals; high moldic intercrystalline porosity
INEQUIGRANULAR: Multimodal crystal size Mosaic
Spotted Isolated and well-defined spots of fine to very fine crystals in a coarse mosaic groundmass
Porphyrotopic: Larger crystals enclosed in in a finer-grained groundmass
Fogged Irregular or diffuse areas of very fine crystals in a coarsemosaic groundmass
Contact-rhomb Loosely aggregated euor subhedral crystals in a fine-grained matrix
Floating rhomb Isolated euhedral or subhedral crystals in a fine-grained matrix
Poikilotopic
Finer crystals enclosed in larger crystals
APHANOTOPIC: Crystals <0.002 mm
B Nonplanar: closely packed anhedral crystals with mostly curved, lobate, serrated, or otherwise irregular intercrystalline boundaries. Preserved crystal-face junctions are rare and crystals often have undulatory extinction under crossed nicols.
Planar-e (euhedral): most dolomite crystals are euhedral rhombs; crystal-supported with intercrystalline area filled by another mineral or porous (as in sucrosic texture).
Planar-s (subhedral): most dolomite crystals are subhedral to anhedral with straight, compromise boundaries and many crystal-face junctions. Low porosity and/or low intercrystalline matrix.
Fig. 7.22. Classification of dolomite fabrics. A: The basic terminology of crystallization textures defined by Friedman (1965) is used in the descriptive classification of Randazzo and Zachos (1983). Friedman’s classification has been expanded by a finer distinction of subtypes, particularly of inequigranular fabrics. Size classes used are 0.256–0.016 mm, 0.016–0.002 mm (indicated by the prefix micro), and <0.002 mm (termed aphanotopic). B: Basic texture types are differentiated according to planar or non-planar crystals defined by crystal shapes and boundaries (After Sibley and Gregg 1987).
Sibley and Gregg (1987) proposed a practical classification of dolomite textures considering crystal size, growth effects expressed by planar or non-planar textures (Fig. 7.22B), and the degree of dolomitization of the grains, matrix, and voids (see Pl. 39 and Pl. 40). Dolomitization of grains is categorized into unreplaced, in molds, partially replaced and replaced. Matrix dolomitization may be unreplaced, partial or replaced. Voidfillings can be unreplaced, partially replaced or replaced by dolomite. 7.8.1.2 Dolomite Cement The occurrence of dolomite cements (Pl. 40/5) precipitated in voids has received little attention, although
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dolomite cements have been reported from various diagenetic environments. Early diagenetic dolomite cements forming bladed rims around grains occur in inter- and supratidal (e.g. Müller and Tietz 1966; Land 1973; Mazzullo et al. 1987; Lasemi et al. 1989) as well as shallow subtidal settings, including reefs (e.g. Mitchell et al. 1987). Dolomite cements from burial settings were described by Dorobek et al. (1993) and Reinhold (1998).
7.8.1.3 Value of Dolomite Textures Textures seen in thin section of dolomites offer some hints to the origin of dolomites, but the assignment to the main models depends on geological and geochemical data. However, many ancient dolomites have undergone multistep episodes of textural and geochemical changes via dissolution-reprecipitation and mineral stabilization (e.g. Ghazban et al. 1992; Nielsen et al. 1994; Reinhold 1998) and many dolomites have been subject to repeated recrystallization. Petrographic evidence of recrystallized dolomite includes an increase in crystal size and increase in the number of non-planar over planar crystal interfaces. Further evidence is given by CL patterns and SEM observations. Geochemical characteristics include dolomite stoichiometry and ordering, depletion of stable isotopes and of Sr and Na, and enrichment in Fe and Mn.
7.8.2 Some Dolomitization Models Dolomites form under different conditions and in depositional environments. The simple distinction between ‘primary’ dolomites believed to be directly precipitated and ‘secondary’ dolomites interpreted as products of the replacement processes has become somewhat obsolete. Primary dolomites seem to be very rare and restricted to saline lakes and lagoons (see Fig. 7.23F). Most of the dolomite in the geological record appears to be of replacement origin. Many models have been developed in order to explain the origin of ancient dolomite rocks. Fig. 7.23 illustrates some of the currently discussed models. Numerical simulation indicates that some of these models may be more and others less reliable with regard to fluid flow patterns in carbonate platforms (Kaufman 1994): The seepage-reflux (Fig. 7.23B) model appears to be a viable dolomitization mechanism, but requires millions of years. Thermal convection of seawater through carbonate platforms (Fig. 7.23G) offers a high
Dolomitization Models
potential for dolomitization. Burial compactiondoes not seem to be an efficient mechanism for regional dolomitization of platforms.
7.8.2.1 Dolomites Associated with Evaporites Various dolomitization models can be seen in Fig. 7.23: A: The evaporative model is based on the comparison with modern evaporative and sabkha dolomite and explains ancient dolomites as supratidal in origin. Holocene dolomite is well documented from tidal settings in the Persian Gulf, the Bahamas and Florida: Supratidal dolomites are typically microcrystalline (1-5 µm) and non-stoichiometric, with weakly ordered crystal structure. These dolomites are found within muddy carbonate sediments or as surface crusts on supratidal flats. Dolomitization is explained as being formed by hypersaline brines derived from intense evaporation in sabkhas. Dolomite replaces preexisting metastable carbonate sediment. Dolomitizing solution is a brine with a high Mg/Ca ratio resulting from Ca removal through precipitation of aragonite, gypsum or anhydrite. Reevaluation of mass balance calculations sheds some doubt on the replacement origin of all sabkha dolomites and discusses the possibility of direct precipitation (Hardie 1987; Lasemi et al. 1989). References: Ambers and Petzold 1996; Butler 1969; Carballo et al. 1987; Chavetz and Rush 1994; Friedman 1980; Gebelein et al. 1980; Gunatilaka 1991; Illing et al. 1965; Lasemi et al. 1989; Mazzullo et al. 1987; McKenzie 1981; Patterson and Kinsman 1982; Shinn et al. 1965.
B: The seepage-reflux model is a popular model for dolomites associated with evaporites. It was developed from studies of a supratidal gypsum-precipitating lake on Bonaire (Netherlands Antilles), and favors a high Mg/Ca ratio and Mg2+-rich hypersaline fluids, permeating underlying carbonate sediments. The model involves the generation of dolomitizing fluids through evaporation of water or tidal flat pore water. The seepage-reflux model has been applied to shelf and lagoonal carbonates of the Permian Reef Complex of western Texas, Late Permian Zechstein dolomites of England, and the Early Cretaceous Edwards Formation in Texas. References: Adams and Rhodes 1960; Clark 1980; Deffeyes et al. 1965; Fisher and Rodda 1960; Kendall 1988; Lu and Meyers 1998; Smith 1981.
C: The evaporation-drawdown model tries to explain the dolomitized intertidal and subtidal facies, often formed beneath evaporites. Dolomitization is regarded as a response to sea-level changes (Magaritz and Peryt
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Dolomitization Models
1994). The model is frequently used to explain the dolomitization of cyclic carbonate-evaporite successions.
Fig. 7.23. Dolomitization models: The two basic questions in all the models developed for the formation of dolomite rocks are the source of the Mg2+ ions and the process involved in pumping dolomitizing fluids through the carbonate sediment. Currently, the interpretation of ancient dolomitization processes considers five models: evaporative (A, C) and seepage-reflux (B), mixing zone (D, E), burial (H) and seawater (I). A survey of papers published during the last twenty years shows that the sabkha evaporation model, various modifications of the mixing zone model, and the burial model are favored over other concepts, but seawater dolomitization models are increasingly being used. But hydrothermal dolomitization and dolomitization associated with organically- and microbially-derived processes (Gornay et al. 1999; Wacey and Wright 2001; Burns et al. 1988, Bernasconi 1994; Mazzullo 2000) are highly disputed models. Slightly modified after Tucker (1985).
D: The Coroong model is based on ephemeral alkaline lakes in South Australia lying behind a modern beach barrier which are fed by the sea and continental groundwater discharging from an unconfined seaward-flowing aquifer. The model stresses evaporation of recharging Mg2+-rich groundwaters. Fine-grained dolomite formed at different times in the annual lake cycle of flooding and desiccation. References: Botz and Von der Borch 1984; Muir et al. 1980; Rosen et al. 1989; Simms 1984; Von der Borch 1976; Von der Borch and Lock 1979; Warren 1988; Wright 1999.
7.8.2.2 Mixing-Water and Seawater Models Dolomites not associated with evaporites are often explained by mixing zone models or by seawater dolomitization models (Fig. 7.23):
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E and F: Meteoric-marine mixing zone models – Dorag model (Dorag =‘mixed blood’) involve mixing of meteoric and marine waters and postulate precipitation from coastal unconfined or deep confined aquifers. The meteoric-marine mixing zone model has been used to explain subtidal dolomites that were formed near-surface relatively early and prior to compaction, and that are not associated with evaporites. Mixing-zone dolomites should be located in the more landward parts of carbonate platforms. References: Badiozamani 1973; Burns and Rossinsky 1989; Cander 1994; Choquette and Steinen 1980; Dunham and Olson 1980; Gill et al. 1995; Humphrey 1988, 2000; Humphrey and Quinn 1989; Land 1973; Meyers et al. 1997; Narkiewicz 1979; Ruppel and Cander 1988; Ward and Halley 1985.
G: Reports of dolomites forming from marine or only slightly modified seawater in marine settings, including platforms, reefs and pelagic (chalks) environments, led to the development of seawater dolomitization models (G). The models emphasize that seawater itself may be able to dolomitize if there is an efficient mechanism for pumping the water through carbonate sediments. A mechanism for moving seawater through the sediment is tidal pumping, resulting in the formation of tidal dolomites (Carballo et al. 1987). Driving mechanisms for seawater circulation through atolls (Minoura 1992), reefs, and platform margins can be oceanic tides and oceanic currents (Saller 1984), the downward reflux of higher saline waters over platforms (Simms 1984) as well as thermal convection caused by higher heat flow through a volcanic basement (Aharon et al. 1987). Thermally driven convective flow processes (endo-upwelling) may operate within the upper part of volcanic foundations and overlying carbonates (atolls, platforms) and can eventually be responsible for carbonate dissolution and formation of massive replacement dolomites. References: Aharon et al. 1987; Flood et al. 1996; Kastner 1984; Kaufman 1994; Land 1985; Machel and Burtin 1994; Mattes and Mountjoy 1980; Mullins et al. 1985; Saller 1984; Tucker and Wright 1990.
H: Large-scale and prolonged circulation of seawater into carbonate platform margins is inferred in the Kohout convection model (H). Kohout convection occurs in response to a horizontal density gradient between cold marine waters adjacent to carbonate platforms and geothermally heated groundwater within the platforms. References: Kohout 1967; Mullins et al. 1985.
Dedolomitization
7.8.2.3 Subsurface Burial Dolomites The principal prerequisites of the burial dolomitization model (I) are sufficient sources of Mg2+, appropriate driving (transport) mechanisms, and favorable conditions for the precipitation of dolomite. The principal process invoked in this model is the compactional dewatering of basinal mudrocks and expulsion of Mg2+rich fluids from porewater during the transformation of clay minerals (conversion of smectite to illite) with increasing depth and rising temperature. Other sources for deep burial dolomites are pressure solution and possible metamorphic and hydrothermal fluids. Hydrothermal dolomitization (Cervato 1990; Boni et al. 2000) is the latest dolomite bandwagon. Many reef and platform carbonates have been extensively replaced by burial dolomites (e.g. Devonian of western Canada: Machel et al. 1994). Criteria of burial dolomitization are coarse crystals, saddle dolomite, the iron content of dolomite crystals, syngenetic or younger formation of dolomite together with stylolites, and specific isotope values. Often only the matrix is dolomitized (‘matrix dolomite’). Saddle dolomite (Pl. 39/5) or ‘baroque dolomite’ are coarse, milky-white, or brown dolomite crystals, generally a millimeter or larger in size, with curved saddlelike crystal faces due to rotating c-axes. The large crystals consist of subcrystals giving the crystal a steepened surface. Fluid inclusions are abundant. In thin sections saddle dolomite appears cloudy, exhibits undulose extinction, and displays growth layers. It occurs in moldic and vuggy pores, less commonly as a massive replacement of carbonates, and often in sulfatebearing carbonate host rocks associated with hydrocarbons and epigenetic sulfides (e.g. Mississippi Valley Type ore deposits). Saddle dolomite is commonly interpreted as having formed under deep burial or hydrothermal conditions from high-saline brines and under high temperatures, or as a by-product of thermochemical sulfate reduction. References: Barnaby and Read 1992; Dix 1993; Eren 1993; Gawthorpe 1987; Gillhaus 2000; Lee and Friedman 1987; Machel and Anderson 1989; Machel et al. 1994; Mattes and Mountjoy 1980; McHargue and Price 1982; Sternbach and Friedman 1984; Mountjoy 1991; Mountjoy and HalimDihardja 1991; Mountjoy et al. 1999; Reinhold 1998; Rosen and Holdren 1986; Sachan 1993; Schofield and Adams 1986; Sternbach and Friedman 1984, 1986.
7.8.3 Dedolomitization Dedolomitization is the diagenetic replacement of dolomite by calcite, particularly under the influence of
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Recognizing Dedolomitization
Box 7.13. Selected case studies of dedolomitization. Meteoric environment: Al-Hashimi and Hemingway 1973; Back et al. 1983; Bischoff et al. 1993; Black et al. 1983; Braun and Friedman 1970; Burger 1989; Cussey et al. 1977; Deike 1990; Evamy 1967; Folkman 1969; Frank 1981, Friedman 1970; Fritz 1966; Goldberg 1967; Jorgensen 1988; Kargel et al. 1996; Lang 1964; Meder 1987; Qingan et al. 1979; Shearman et al. 1961; Török 1997; Ulmer 1984; Warrak 1974; Wirsing 1986. Lacustrine environment: Arenas et al. 1999; Canaveras et al. 1996. Shallow burial environment: Al-Hashimi and Hemingway 1973; Bausch et al. 1986; Black et al. 1983; Dunham and Olson 1980; James et al. 1993; Kenny 1992; Kyung Sij Woo and Moore 1996; Lucia 1961; Meder 1987; Purser 1985; Sellwood et al. 1984. Deep burial environmemt: Budai et al. 1984; Land and Prezbindowski 1981; Mattes and Mountjoy 1980; Stoessel et al. 1987.
meteoric water and porewaters of different composition, and often resulting in the formation of secondary porosity. This process affects marine, lacustrine, and terrestrial carbonates and occurs in meteoric and burial diagenetic environments (Box 7.13). Early diagenetic dedolomitization may result from (a) the instability of Ca-dolomite within crystal cores facilitating the replacement of the core by calcite, (b) near-surface recrystallization under meteoric conditions (characterized by rhombohedrons displaying microcrystalline calcite with zonar hematite zones), and (c) meteoric dissolution of dolomite rhombohedrons within the micrite and infilling of the crystal molds by granular meteoric cements (Pl. 40/ 4) and geopetal internal sediment. Late diagenetic dedolomitization controlled by salinity variations of burial pore waters can result from (a) the diagenetic instability of Ca-dolomite (similar to the above mentioned early diagenetic dedolomite), and (b) corrosion of zonar rhombohedron along cleavage planes, formation of intracrystalline pores and later growth of syntaxial calcite within these pores (Pl. 40/5). The alteration of dolomite to calcite was first reported by the Swiss geologist A. von Morlot (1847) based on solution experiments and field studies in Styria. He coined the term dedolomite and suggested that groundwater leaching of preexisting evaporite deposits containing gypsum or anhydrite would be enriched with respect to calcium and sulfate, thus thereby resulting in the calcitization of former dolomite. The term dedolomitization is now used along with calcitization to describe the replacement of dolomite by calcite. Staining with Alizarine Red S shows the distribution of calcite and dolomite within the dolomite rhombs very clearly.
7.8.3.1 Textural Criteria for Recognizing Dedolomitization • Brown and reddish rock colors: Liberation of Fe2+ from Fe-rich dolomite may form hematite occurring as thin coatings on the crystals or ferric iron precipitates. Many dedolomites are heavily stained with iron oxides and hydroxides and exhibit brown or reddish rock colors. • Weathering along crystal faces of a calcitized dolomite may produce a loose sandy sediment. • Calcite pseudomorphs after sucrosic xenomorphic, nonplanar dolomite. • Syntaxial calcite rims bordering rhombohedral crystals. • Relicts of dissolved dolomites preserved inside the pseudomorphs or at their boundaries. Dedolomitization may be centripetal or centrifugal (center to periphery: Shearman et al. 1961). • Association of dedolomite with evaporite minerals: Dedolomite is often associated with pseudomorphs after evaporitic minerals, because the solution of evaporite minerals favors calcitization.
7.8.3.2 Origin of Dedolomite Two general mechanisms have been proposed - a reaction of dolomite with calcium sulfate solutions (Morlot 1847; Shearman et al. 1961; Evamy 1963; Goldberg 1967) or the alteration of ferroan dolomite by oxygenated meteoric water. The sulfate ions needed for the reaction are believed to be derived from the oxidation of pyrite or from gypsiferous solutions. Nearsurface dedolomitization is often related to dissolution of gypsum and dolomite in vadose or phreatic meteoric waters. Meteoric waters falling on outcropping gypsum generate a solution with a very high Ca/Mg ratio. In some cases, dedolomite may be indicative of subaerial exposure. Gypsum dissolution drives the precipitation of calcite, thus consuming carbonate ions released from dolomite. This may lead to karstification. Prerequisites for the formation of near-surface dedolomite is a solution high in dissolved calcium and low in magnesium, low CO2 partial pressure, and temperatures < 50 °C. Dedolomitization by ground water in carbonate aquifers may occur on a regional scale (e.g. Back et al. 1986; Deike 1990). CO2-bearing waters contribute to the dissolution of dolomite, which sometimes leads to the formation of dolomite sands under recent subaerial conditions (Fritz 1966). Continuous replacement of dolomite by calcite via solution/precipitation
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is indicated by xenotopic fabrics. Complete dolomite solution and subsequent cementation is shown by dolomolds. Reaction rates and the mode of dedolomitization in experimental studies appear to depend on solution concentrations, and the temperature and crystal size of the dolomite (De Groot 1967; Kastner 1984; Stoessel et al. 1987). Experiments point to a relatively slow process. In contrast, dedolomitization of dolomitic marbles, resulting in the formation of calcite crusts, seems to be a rather rapid process as shown by the study of ancient sculptures (Doehne et al. 1992).
7.8.3.3 Significance of Dedolomitization Recognizing subaerial exposure and unconformities: Most authors consider dedolomitization to be a nearsurface process related to weathering. Laterally con-
Marbles
tinuous dedolomitized horizons, therefore, may indicate the presence of unconformities (Brown and Friedman 1970). Porosity development: Dedolomitization can increase moldic and intercrystalline porosity, forming reservoir rocks (Purser 1985). Open dolomolds may be indicative of subaerial exposure and freshwater effects.
7.9 Metamorphic Carbonates and Marbles The deep burial diagenetic environment grades into the zones of metamorphic control. Metamorphic processes occur at various depths, under different pressure and thermal conditions, and within a time range of millions and more years. There is no clear, generally accepted distinction between diagenesis and metamorphism,
Plate 40 Selective Dolomitization and Dedolomitization of Platform Carbonates Dolomitization, particularly multigenerational burial dolomitization, often causes obliteration of primary constituents and textures. Many dolomites, however, exhibit selective dolomitization patterns (Murray 1964), indicated by the dolomitization of only the fine-grained matrix (–> 2, 5) or only of individual grains (–> 6, 7). 1 Heterogeneous dolomitization affecting only parts of the carbonate rock. Larger and smaller isolated euhedral dolomite rhombohedrons (‘floating rhombs’) occur within a microcrystalline calcite matrix. Using the Sibley and Gregg classification, the sample would be described as partial dolomite exhibiting unimodal and euhedral planar crystals. Some crystals are basically filled with calcite internal sediment (arrows), indicating meteoric dedolomitization (see Pl. 39/7). Lagoonal facies. Late Jurassic: Yumakle, southern Turkey. 2 Fabric-selective dolomitization. Distinctly zoned euhedral dolomite crystals occur only within the fine-grained matrix and at the periphery of an auloporid coral. Zonation is accentuated by Fe-rich zones (appearing black) within the dolomite crystals pointing to burial dolomitization. Matrix dolomitization by unimodal and euhedral planar crystals. Mediumcrystalline bioclastic dolomite. Early Devonian (Emsian): Erfoud, Morocco. 3 Partial burial dolomitization. Fossils (echinoderms, E; styliolinids, S) are not dolomitized. Arrows point to a calcitic core of dolomite rhombs. Note micrite relicts within the recrystallized matrix. Early Devonian (Emsian): Erfoud, Morocco. 4 Early diagenetic dedolomite. Molds of former dolomite rhombs are filled with vadose meteoric calcite. Fenestral bindstone. Lagoonal facies. Early Jurassic: Korfu Island, Greece. 5 Selectively dolomitized ooid grainstone. Replacement dolomitization is restricted to intergranular pores. Note the irregular boundary between tangentially-structured ooids and the dolomite crystals forming circumgranular rims (arrows) consisting of dolomite crystals similar to the dolomite crystals occluding the remaining pore space. The circumgranular rims represent isopachous dolomite cement, probably formed in a shallow burial setting. Late Triassic (Carnian): Central Carinthia, Austria. 6 Selectively dolomitized ooid grainstone. Note the distinct difference in comparison with –> 5. Only ooids are dolomitized but not the echinoderms (E). The ooids are mimically replaced, e.g. they exhibit still traces of their concentric microstructure. Cambrian: Takkanawa Fall, Yoho Valley, British Columbia, Canada. 7 Selective dolomitization restricted to micritic clasts (MC) and the micrite infilling of the gastropod (arrow). Restriction is probably controlled by the High-Mg mineralogy of the microcrystalline calcite. Skeletal grains are echinoderms (E) and gastropods. Back-reef facies. Shallow-burial replacement dolomitization. Late Devonian: Canning Basin, western Australia. 8 Dolomitized ooids. The different intensity in dolomitization of the ooid layers and nuclei possibly reflects differences in their composition. Back-reef facies. Late Devonian (Frasnian): Canning Basin, western Australia.
Plate 40: Dolomitization and Dedolomitization
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Marbles
Box 7.14. Descriptive criteria of marbles in thin sections. Visual differentiation of marble samples takes into account criteria related to the petrogenesis of marbles, such as shape, size distribution, and spatial distribution.
Fig. 7.24. Carrara marble is one of the best known marbles, appreciated by structural geologists interested in deformation tests, loved by artists like Michelangelo, and already used in the antique for sculptures and as building stones. The brightness and the translucity of the marble is due to the strikingly homogeneous composition and texture. Calcite crystals are very regular in shape, and consistent in size, and exhibit regular boundaries. The marble developed from a Jurassic limestone, which was affected during the Tertiary by folding, metamorphism, and heating processes. Carrara, Alpi Apuani, Italy. Scale is 1 mm.
except that metamorphism occurs at higher temperatures and pressures than diagenetic processes. The ultimate products of increasing strain and temperature are carbonate marbles (Pl. 38/4-8). The fabric of marbles seen in thin sections (Fig. 7.24) is predominantly the result of the deformation and recrystallization history of the rock. Lithotypes of marbles are caused by tectono-metamorphic deformation of sedimentary lithologies and ductile processes leading to tectonic fabrics (e.g. schistosity and fractures). Dynamic recrystallization of the protolith (the precursor rock) changes the size, shape, boundaries, and the crystallographic orientation of calcite crystals. Increasing recrystallization leads to complete obliteration of sedimentary textures and the formation of pavement mosaics of completely interlocking crystals. The main effects of metamorphic alterations of limestones are an increase in the matrix crystal size, followed by a progressive loss of the original texture, and, ultimately, total obliteration of fossils and other grains. Increasingly more thin-section studies are being done of marble exhibiting relicts of the texture, composition and even the microfacies of the precursor rock. Methods The methods used for investigating marbles depends on the questions that need to be answered and requires a multi-method approach applying different micro-
Texture Crystal size: Average size, maximum size, and size ranges. At least 100 crystals should be determined (use line-intercept method or digital image analysis, Lumbreras and Serrat 1996). Maximum diameter and diameter ranges are used in distinguishing marbles of different provenance. The increase in the mean diameter may be caused by thermal control (Schmid 1997; Taguchi and Aoki 1998). Equi- or non-equidimenional fabric (see Fig. 7.19)? Uni- or bimodal size distribution? Bimodal distribution may indicate growth of crystals at the expense of adjacent crystals. Crystal boundaries: Straight or gently curved, or irregular and serrate? Boundaries reflect grain growth history. Crystal shape and fabric can be described by using the terms proposed by Friedman (1965) or those introduced by Sibley and Gregg (see Fig. 7.22). Microstructural fabrics Some of the criteria that describe conditions of crystal growth are twin and glide surfaces, annealing and segregation of a crystal into subcrystals, and alternation of undulose extinction. Calcite twinning and growth of twins (Pl. 38 /4, 7, 8) is related to the formation of dislocations and to the glide of dislocations within slip planes. Twinning types are used to estimate paleostress history, deformation temperature, and relative timing of tectonic events (Wenk et al. 1983; Simonsen and Friedman 1992; Railsback 1993). Twinning types are differentiated according to the abundance, thickness and orientation of twin lamellae (thin, thick, straight, or sutured twins; Burghard 1993, with many references). Thin twin lamellae record small strains in marbles deformed under little cover. Curved thick twin lamellae and patchy twin lamellae indicate large deformation. The appearance of twin lamellae changes as a function of deformation temperature. Crystal fabric analyses allow marble types (Schmid et al. 1999) and deformation styles to be distinguished. Accessory minerals Common accessory minerals in calcite and dolomite marbles are quartz, clay minerals, mica, pyrite, and graphite. Fossils Fossils are very rare in marble thin sections except for scarce crinoidal fragments or ghosts of shells (Pl. 38/6). Insoluble residues of some marbles may yield microfossils, particularly pyritized, originally siliceous radiolarians and sponge spicules (Kiessling 1992).
Basics: Diagenesis, Porosity, Dolomitization
scopic, chemical, and isotopic techniques. Archaeologists concerned with the problem of where antique marbles originated from will use all the data (e.g. texture, isotopic signals, geochemical criteria) that distinguish marbles from different regions and localities (e.g. Herz and Waelkens 1988; see Sect.19.6). Geologists trying to decipher the tectonic and low-grade metamorphism history of marbles will look for stress and strain markers (e.g. Molli and Heilbronner 1999) and indicators of paleothermal levels (e.g. illite crystallinity, vitrinite reflectance, thermal alteration indices based on fossils. Buggisch 1986; Königshof 1992; Krumm 1992). Microfacies-oriented people will look for relicts of premetamorphic features, such as sedimentary structures, fossils, and organic matter that can be interpreted as derivates of former organisms. Lamination and graphite banding of marbles, expressed by differences in color, carbon content, and differences in dominating crystal sizes, can reflect original depositional structures as, demonstrated e.g. by loferite structures in Styrian marbles (Lelkes et al. 1999). Descriptive criteria of marbles seen in thin sections are summarized in Box 7.14. Basics: Diagenesis, porosity and dolomitization Carbonate diagenesis: Overviews Bathurst, R.G.C. (1989): Early diagenesis in carbonate sediments. – In: Parker, A., Sellwood, B.W. (eds.): Sediment diagenesis – 345-377, Dordrecht (Reidel) Berner, R.A. (1980): Early diagenesis as a theoretical approach. – 241 pp., Princeton (Princeton University Press) Garrison, R.E. (1981): Diagenesis of oceanic carbonate sediments: A review of the DSDP perspective. – Soc. Econ. Paleont. Min. Spec. Publ., 32,181-207 Macillreath, I.A., Morrow, D.W. (eds., 1990): Diagenesis. – Geoscience Canada Reprint Series, 4, 338 pp. Marshall, J.D. (1987): Diagenesis of sedimentary sequences. – Geol. Soc. America, Spec. Publ., 36, 368 pp. Moore, C.H. (1989): Carbonate diagenesis and porosity.– Developments in Sedimentology, 46, 338 pp. Morse, J.W., Mackenzie, F.T. (1990): Geochemistry of sedimentary carbonates. – Developments in Sedimentology, 48, 707 pp. Parker, A., Sellwood, B.W. (eds., 1983): Sediment diagenesis. – 427 pp., Dordrecht (Reidel) Tucker, M.E., Bathurst, R.G.C. (eds., 1990): Carbonate diagenesis. – Int. Assoc. Sedimentologists Reprint Series, 1, 312 pp. Mineralogical constraints and processes Fernandez-Doau, L., Putnis, A., Prieto, M., Putnis, C.V. (1996): The role of magnesium in the crystallization of calcite and aragonite in a porous medium. – J. Sed. Research, A66, 482-491 Folk, R.L. (1974): The natural history of calcium carbonate: effect of magnesium content and salinity. – J. Sed. Petrol., 44, 40-53 Given, R.K., Wilkinson, B.H. (1985): Kinetic control of morphology, composition and mineralogy of abiotic sedimen-
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tary carbonates. – J. Sed. Petrol., 55, 109-119 Reeder, R.E. (ed., 1983): Carbonates: mineralogy and chemistry. – Reviews in Mineralogy, 11, 394 pp. Richter, D.K. (1984): Zur Zusammensetzung und Diagenese natürlicher Mg-Calcite. – Bochumer Geologische und Geotechnische Arbeiten, 15, 310 pp. Walter, L.M., Morse, J.W. (1984): Magnesian calcite stabilities: a reevaluation. – Geochim. Cosmochim. Acta, 48, 1059-1069 Further reading: K034, K035 Diagenetic environments Choquette, P.W., James, N.P. (1987): Diagenesis, 12, Diagenesis in limestones, 3, The deep burial environment. – Geoscience Canada, 14, 3-35 James, N.P., Choquette, P.W. (1990): Limestones - the meteoric diagenetic environment. – In: Macillreath, I.A., Morrow, D.W. (eds.): Diagenesis. – Geoscience Canada, 11, 161-194 James, N.P., Choquette, P.W. (1990): Limestones - the burial diagenetic environments. – In: Macillreath, I.A., Morrow, D.W. (eds.): Diagenesis. – Geoscience Canada Reprint Series, 4, 75-111 Longman, M.W. (1980): Carbonate diagenetic textures from nearsurface diagenetic environments. – Amer. Ass. Petrol. Geol. Bull., 64, 461-487 Moore, C. (2001): Carbonate reservoirs. – 460 pp., Amsterdam (Elsevier) Schroeder, J.H., Purser, B.H. (eds., 1986): Reef diagenesis. – 455 pp., Berlin (Springer) Wanless, H.R. (1989): Burial diagenesis in limestones. – In: Parker, A., Sellwood, B.W. (eds.): Sediment diagenesis. – 379-417, Dordrecht (Reidel) Further reading: K034, K036, K037, K045, K050 Porosity Anselmetti, F.S., Luthi, S., Eberli, G.F. (1998): Quantitative characterization of carbonate pore systems by digital image analysis. – Amer. Ass. Petrol. Geol. Bull., 82, 1815-1836 Archie, G.E. (1952): Classification of carbonate reservoir rocks and petrophysical considerations. – Amer. Ass. Petrol. Geol. Bull., 36, 278-298 Bathurst, R.G.S. (1986): Carbonate diagenesis and reservoir development: Conservation, destruction and creation of pores. – In: Bathurst, R.G.C., Land, L.S. (eds.): Carbonate deposition environments. Part 5. Diagenesis 1. – Colorado School of Mines Quarterly, 81, 1-25 Chilingarian, G.V., Mazzullo, S., Rieke, A., Deminguez, G.C., Samaniego, F.Y. (1992, eds.): Carbonate reservoir characterization: a geologic-engineering analysis, part I. – 640 pp., Amsterdam (Elsevier) Choquette, P.W. and Pray, L. (1970): Geologic nomenclature and classification of porosity in sedimentary carbonates. – Amer. Ass. Petrol. Geol. Bull., 54, 207-250 Enos, P. (1988): Evolution of pore space in the Poza Rica trend (mid-Cretaceous), Mexico. – Sedimentology, 35, 287-335 Enos, P., Savatzky, N.H. (1981): Pore networks in Holocene carbonate sediments. – J. Sed. Petrol., 51, 961-985 Friedman, G.M., Ali, S.A. (eds., 1981): Diagenesis of carbonate rocks. Cement-porosity relationships. – Soc. Econ. Paleont. Miner., Reprint Series, 10, 295 pp. Kopaska-Merkel, D.C., Mann, S.D. (1993): Classification of lithified carbonates using ternary plots of pore facies: ex-
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amples from the Jurassic Smackover Formation. – In: Rezak,R., Lavoie, D.L. (eds.): Carbonate microfabrics. – Frontiers in Sedimentary Geology, 265-277 Moore, C. (2001): Carbonate reservoirs. – 460 pp., Amsterdam (Elsevier) Roehl, P.O., Choquette, P.W. (eds., 1985): Carbonate petroleum reservoirs. – 622 pp., Berlin (Springer) Schroeder, J. (1988): Spatial variations in the porosity development of carbonate sediments and rocks. – Facies, 18, 181-204 Further reading: K034, K073 to K076 Carbonate cements Bricker, O.P. (ed., 1973.): Carbonate cements. – John Hopkins University Studies in Geology, 19, 376 pp. Harris, P.M., Kendall, C.G., Lerche, J. (1985): Carbonate cementation: a brief review. – In: Schneidermann, M., Harris, P.M. (eds.): Carbonate cements. – Soc. Econ. Paleont. Miner., Spec. Publ., 36, 79-95 Koch, R., Zinkernagel, U. (1994): Zur Zementation in Kalksteinen. – Zentralblatt für Geologie und Paläontologie, Teil I, 1994, 1353-1398 Meyers, W.J. (1991): Calcite cement stratigraphy: an overview. - In: Barker, C.E. and Kopp, O.C. (eds.): Luminescence microscopy and spectroscopy; qualitative and quantitative application. – Soc. Econ. Paleont. Miner., Short Course, 25, 133-148 Sandberg, P. (1985): Aragonite cements and their occurrence in ancient limestones. – In: Schneidermann, N., Harris, P.M. (eds.): Carbonate cements. – Soc. Econ. Paleont. Miner., Spec. Publ., 36, 33-58 Schneidermann, N., Harris, P.M. (1985): Carbonate cements. – Soc. Econ. Paleont. Miner., Spec. Publ., 36, 397 pp. Further reading: K039–K041, K048 Diagenetic textures Bathurst, R.G.C. (1987): Diagenetically enhanced bedding in argillaceous platform limestones: stratified cementation and selective compaction. – Sedimentology, 34, 749778 Chilingarian, G.V. (1989): Compactional diagenesis. – In: Parker, A., Sellwood, B.W. (eds.): Sediment diagenesis, 57-166 Hutcheon, I.E. (ed., 1989): Burial diagenesis. Short Course Handbook. – Mineralogical Association of Canada, 15, 499 pp. Logan, B.W. (1984): Pressure responses (deformation) in carbonate sediments and rocks – analysis and application, Canning Basin. – In: Purcell, P.G. (ed.): The Canning Basin, W.A. – Proceedings of the Geological Society of Western Australia/Petroleum Exploration Society of Australia, 235-251, Perth Logan, B.W., Semeniuk, V. (1976): Dynamic metamorphism; process and products in Devonian carbonate rocks; Canning basin, western Australia. – Geological Society of Australia, Spec. Publ., 6, 138 pp. Ricken, W. (1987): The carbonate compaction law; a new tool. – Sedimentology, 34, 571-584 Shinn, E.A., Robbin, D.M., Claypool, G.E. (1984): Compaction of modern sediments: implications for generations and expulsion of hydrocarbons. – In: Palacas, J.G. (ed.): Petroleum geochemistry and source rock potential of carbonate rocks. – Amer. Ass. Petrol. Geol., Studies in Geology, 18, 197-203
Basics: Cements, Dolomitization, Marbles
Wanless, H.R. (1979): Limestone response to stress: Pressure solution and dolomitization. – J. Sed. Petrol., 49, 437-462 Further reading: K046, K048, K100 Neomorphism Folk, R.L. (1965): Some aspects of recrystallization in ancient limestones. – In: Pray, L.C., Murray, R.C. (eds.): Dolomitization and limestone diagenesis. – Soc. Econ. Paleont. Miner., Spec. Publ., 13, 14-48 Friedman, G.M. (1965): Terminology of recrystallization textures and fabrics in sedimentary rocks. – J. Sed. Petrol., 35, 643-655 Further reading: K034, K035, K059 Dolomitization and dedolomitization Allan, J.R., Wiggins, W.D. (1993): Dolomite reservoirs: Geochemical techniques for evaluating origin and distribution. – Continuing Education Courses Notes, Amer. Ass. Petrol. Geol., 36, 167 pp. Land, L.S. (1983): Dolomitization. – Amer. Ass. Petrol. Geol., Continuing Education Courses Notes, 24, 1-20 Machel, H.G, Mountjoy, E.W.(1986): Chemistry and environments of dolomitization - a reappraisal. – Earth Sciences Revolutions, 23, 175-222 Purser, B., Tucker, M., Zenger, D. (eds., 1994): Dolomites. A volume in honour of Dolomieu. – Spec. Publ. Int. Ass. Sed., 21, 451 pp., Oxford (Blackwell) Randazzo, A.F., Zachos, N.G. (1983): Classification and description of dolomitic fabrics of rocks from Floran aquifer, U.S.A. – Sedimentary Geology, 37,151-162 Sellwood, B.W., Scott, J., James, B., Evans, R., Marshall, J. (1987): Regional significance of dedolomitization in Great Oolite reservoir facies of S. England. – In: Brooks, J., Glennie, K. (eds.): Petroleum Geology of N.W. Europe. – 129-137, London (Graham and Trotman) Shukla, C., Baker, D.V. (eds., 1988): Sedimentology and geochemistry of dolostones. – Soc. Econ. Paleont. Miner., Spec. Publ., 43, 266 pp. Sibley, D.F., Gregg, J.M. (1987): Classification of dolomitic rock textures. – J. Sed. Petrol, 57, 957-965 Warren, J. (2000): Dolomite: occurrence, evolution and economically important associations. – Earth-Science Reviews, 52, 1-81 Zenger, D.H., Dunham, J.B., Ethington, R.L. (eds., 1980): Concepts and models of dolomitization. – Soc. Econ. Paleont. Miner., Spec. Publ., 28, 320 pp. Further reading: K060–K063 Metamorphic carbonates and marbles Burkhard, M. (1993): Calcite twins, their geometry, appearance and significance as stress-strain indicators of tectonic regime: a review. – Journal of Structural Geology, 15, 315368 Herz, N., Waelkens, M. (eds., 1980): Classical marble: geochemistry, technology, trade. – NATO Advanced Science Institute Series, E, 153, 482 pp. Molli, G., Conti, P., Giorgetti,G., Meccheri, M., Oesterling, N. (2000): Microfabric study on the deformational and thermal history of the Alpi Apuane marbles (Carrara marbles), Italy. – J. of Structural Geology, 22, 1809-1825 Siegesmund, S., Weiss, T., Vollbrecht A.. Ullemeyer, K. (1999): Marble as natural building stones: rock fabric, physical and mechanical properties. – Zeitschrift der deutschen geologischen Gesellschaft, 150, 237-257 Further reading: K034, K051 to K057
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8 Classification – A Name for Your Sample
There are too many papers that simply conclude that the carbonate rocks of a given study can be arranged according to the classification of one or more authors. All too often these workers have failed to recognize that a classification is simply a tool for organizing information, not the source of a conclusion (Blatt et al. 1972). Classifying and naming carbonate rocks is a rather dry pursuit, but nonetheless essential to microfacies analysis when interpreting rock properties and paleoenvironments. Classifications should include objective, measurable and reproducible features that allow genetic inferences to be made from descriptive criteria. All the limestone classifications commonly used in facies analyses are based on textural and compositional criteria. The classifications proposed by Dunham (1962) and Folk (1962) have proved most practical. Modifications suggested by Embry and Klovan (1971), Wright (1992) and Strohmenger and Wirsing (1991) are of help. Naming and classifying reef limestones, nonmarine carbonates as well as recrystallized carbonate rocks call for specific concepts. Major problems involved in the classification of limestones are the ‘low-energy micrite’ paradigm, the confusion between depositional and diagenetic textures, and the underestimation of the complexity of biological controls on the formation of reef carbonates. Demands arising from the industrial use of carbonate rocks as raw materials led relatively early to classifications focused almost entirely on the chemical and mineralogical composition. Most textbooks written before World War II discuss only these classifications. Classifications based on environmental conditions and depositional settings began with Pia (1933). During their work on Arabian reservoir rocks, Bramkamp and Powers (1958) were probably the first to stress the importance of distinguishing between carbonate sediments deposited in quiet-water environments and those formed as current-washed deposits. The distinction between ‘low energy’ and ‘high energy’ carbonates has become a ‘leitmotif’ of almost all later classifications, whether
designated as ‘genetic’ or ‘descriptive’. The concept tries to consider differences in the kinetic energy (caused by wave or current action) that might have existed within the water column either at the interface of deposition or a meter or two above it. Geologists like this rather simple dual distinction between high and low energy sediments!
8.1 Basic Concepts There are different approaches to carbonate classifications, including (1) the choice between textural and nontextural classification concepts, (2) the preference for strongly descriptive or genetic classifications, (3) the divergent consideration of various aspects of deposition, biotic contribution and diagenesis, and 4) the differentiation between biologically controlled limestones formed in place, and limestones produced by the accumulation of grains. (1) Non-Textural versus Textural Classifications Non-Textural Classifications Mineralogical and chemical composition: The mineralogical composition of carbonate rocks (rocks containing more than 50% by mass of carbonate minerals) provides a relatively simple means for subdivisions (see Sect. 3.1.1.1). The distinction between calcium carbonate minerals (predominantly calcite) and dolomite forms a basic element in almost all chemical carbonate rock classifications. The term limestone is generally used for rocks composed predominantly of calcite (>50%), the term dolomite (or dolostone) characterizes rocks consisting predominantly of the mineral dolomite (>50%). The industrial use of carbonate rocks, in particular requires categories describing intergradations caused by differences in the relative amounts of calcite and dolomite and by various amounts of non-carbonate minerals. Differences in chemical composition are expressed as Ca/Mg weight ratios or as percentages of calcite and dolomite or non-carbonate constituents. Dif-
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_8, © Springer-Verlag Berlin Heidelberg 2010
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Autochthonous Carbonates
ferent limits are used to define subgroups within the calcite/dolomite/non-carbonate triangle. Subdividing carbonate rocks according to their chemical composition (e.g. CaO, MgO, SiO2, and P2O5) is essential for the economic use of limestones and dolostones (e.g. Oates 1998). Mineralogical-chemical classifications are discussed in Sect. 18.1. Physical properties: Reservoir rock studies need classifications which consider petrophysical properties such as porosity and permeability. These classifications (Choquette and Pray 1970; Mazzullo et al. 1992; Kopaska-Merkel and Mann 1993) differ significantly from conventional limestone classifications because they focus on genetic complex criteria (type, pore geometry and distribution) which are controlled by the total of the factors reflecting the history of the rock (see Sect. 7.3).
subjective. Most of the currently used limestone classifications are descriptive-genetic and emphasize the descriptive data that provide evidence of environmental and depositional controls. The term eventstone, introduced by Hüssner (1985), is strictly genetic. It designates limestones that were deposited relatively quickly as compared with the background sedimentation. Eventstones include turbidites and tempestites and exhibit various depositional fabric types. The name should not be used in a descriptive way (see Sect. 12.1.2). A distinctly genetic classification was proposed by Plumley et al. (1962). The classification is based on the inferred degree of water agitation. The authors distinguish five major limestone categories, with a grading spectrum ranging from quiet water to strongly agitated water deposits (see Sect. 12.1.1.2).
Textural classifications Nearly all limestone classification systems used today were developed during the 1950s and 1960s in response to the needs of oil companies who had discovered major hydrocarbon reservoirs in carbonate rocks. The long-standing bestselling book Classification of Carbonate Rocks (Ham 1962) published by the American Association of Petroleum Geologists, includes the two standard classifications of Folk and Dunham as well as less frequently used systems (Leighton and Pendexter 1962; Powers 1962; Thomas 1962). Folk’s Practical classification of Limestones and Dunham’s Classification of carbonate rocks according to depositional structure have become largely accepted and widely used systems. Both classifications were modified during the last few decades: New grain types were added to Folk’s system and quantitative boundaries changed (Flügel 1982; Strohmenger and Wirsing 1991), additional terms were introduced to Dunham’s classification allowing a more detailed differentiation of reef carbonates (Embry and Klovan 1971; Tsien 1981; Cuffey 1985) and diagenetically altered limestones (Wright 1992). Some of these proposals were criticized, others, particularly those dealing with names coined for reef carbonates, were apparently misunderstood because of the different geological and biological approach in reef studies.
(3) Depositional, biological and diagenetic aspects Classifications are differently focused on depositional, biological-paleontological and diagenetic criteria seen in carbonate rocks. Most classifications using texture and fabric claim that the groupings reflect differences in depositional processes. In contrast, classifications of reef carbonates demand an understanding of biological controls and of the interplay between organisms, sedimentation and cementation. Diagenesis modifies and obliterates primary depositional textures, and creates new fabrics that require specific nomenclatorial treatment (see Pl. 47).
(2) Descriptive versus genetic views Because a geologist must commonly classify and name a sample before its origin can be determined, descriptive classifications are preferable, at least during field work. Genetic classifications require an understanding of the complicated processes responsible for the formation of a given rock type and are, hence, more
(4) Autochthonous and allochthonous carbonates Most limestone classifications differentiate between two major groups: (1) limestones originating from the activity of benthic sessile organisms and (2) limestones composed of grains, lime mud and/or carbonate cement. The first group contains autochthonous carbonates. The second group has been called allochthonous carbonates (or ‘allochemical’ carbonates sensu Folk 1959). Characteristic autochthonous carbonates are reef limestones and carbonates whose formation is influenced by microbes (microbialites). Note that the term ‘allochthonous carbonates’ is also used in the context of massflow deposits (see Sect. 15.7.2).
8.2 Reef Limestones and Microbial Carbonates (Autochthonous Carbonates) Starting with Grabau (1904), all major classifications have considered limestones which were organically bound at the time of deposition as an independent group.
Autochthonous Carbonates
341
Fig. 8.1. Autochthonous carbonates. A – Reef rock types (Embry and Klovan 1971). The boundstone category is subdivided into bafflestone, bindstone, and framestone (Pl. 41 and Pl. 42; see Box 8.1 for definitions). The use of these names requires a decision on whether or not the fossil organisms bound the sediment during deposition and if so, in what manner. Framestone, bindstone, and bafflestone, therefore, are highly interpretative rock types. Floatstone and rudstone designate limestones with transported grains larger than 2 mm (Pl. 45). B – Roles of organisms in the construction of reef carbonates. Note that framestone fabrics can originate from supporting, secondary stabilizing, and cementing. Bindstone fabrics are formed by binding, stabilizing, and microbial micrite. After Tsien (1994). C – Dependence of reef rock types on environmental constraints of benthic communities. Reefal fabrics can be placed into four groups, which are united by the morphology of the dominant members of the reef-building community: Laminar forms that stabilize sediment may form bindstones, cruststones and coverstones; mutual encrusting, domal, and massive branching morphologies may give rise to framestones and branchstones. Cup-shaped, stick or branched forms that accumulate sediment can produce bafflestones and branchstones. Note that ecologically disparate communities may produce convergent rock fabrics. The distribution of reef biota is highly dependent on nutrient availability, which has profound effects on the trophic structure and growth forms of reefbuilders, and hence, on the resultant carbonate sediment and its fabrics. An important message from this model is that different hydrodynamic energy levels are not the only and often not the most important causes of different reef rock fabrics. After Wood (1993).
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Both, Folk (1959) and Dunham (1962) proposed names for limestones whose formations were dependent on sessile benthic organisms. Folk uses the term biolithite as a sackname for autochthonous reef rocks exhibiting corals, algae and other organisms in growth position. The name should be discarded in favor of the other sackname boundstone introduced by Dunham. The author’s examples include coral limestones with corals in living position, stromatolite limestones, and limestones exhibiting dense intergrowth of various encrusting organisms. The difference between geological and biological thinking: Dunham did not specify how the organisms
Autochthonous Carbonates: Terminology
contribute to rock-building processes. Problems arose as subsequent modifications of the Dunham classification tried to take a more stringent look at the inferred role of organisms in reef-building: Geologists are interested in distinguishing reef rock types by organic structures. Biologists and paleontologists examine the significance of reef builders for the buildup of reefs (Fig. 8.1). Subdivision of autochthonous carbonates More and more names for smaller and smaller differences? Beginning with the well-founded subdivision of boundstones into the genetically defined categories
Box 8.1. Genetic-descriptive terms used for limestones built predominantly by sessile organisms (autochthonous lime stones: reef carbonates and microbialites). The author of the original definition is marked by an asterisk. Bafflestone – Carbonate rocks containing in-situ stalk-shaped and branching fossils that trapped sediment during deposition by acting as bafflers, slowing down water movement and allowing sediment to settle (*Embry and Klovan 1971). Minimally transported members of the baffler guild predominate in the samples. Pl. 41/1, Pl. 116/3 Bindstone – (a) Carbonate rocks containing in situ laminar, tabular, or encrusting fossils which encrusted and bound sediment during deposition (*Embry and Klovan 1971). In bindstones the in-situ fossils do not form a three dimensional framework as they do in framestone. (b) Carbonate rocks dominated by encrusting organisms that encrust or bind debris together, forming a framework (Tsien 1981). (c) Carbonate rock containing in-situ fossils that encrust and bind sediment during deposition. The matrix and not the in-situ fossils form the supporting framework of the rock; fossils may represent as little as 15% of the rock volume (Fagerstrom 1987). Pl. 6/5, Pl. 8/1-2, Pl. 123/3 Biocementstone – Carbonate rocks formed by algae and bacteria which produce micrite cement and trap, stabilize, and support lime mud (*Tsien 1981, 1984). Bioconstructed limestones – Limestones characterized by ‘colonial framework builder organisms with interstitial material’ (Carozzi 1989). Biolithite – A ‘sackname’. Limestone constructed by organisms that grew in place, forming a rigid framework that bound sediment and remained in place (* Folk 1959). The name should not be used for limestones containing broken off fragments of reef organisms. Branchstone – Reef carbonate rocks, characterized by branching colonies consisting of densely-spaced, touching skeletal elements with only minor sediment in between; e.g. some modern Acropora thickets (* Cuffey 1985). Compare Fig. 8.1C. Cementstone – (a) Reef limestones composed of >50% spar of early and late diagenetic as well as neomorphic origin (*Fagerstrom 1987). The cement volume may exceed that of organic framework or internal sediment. Pl. 50/6, Pl. 145/1. (b) Limestones composed almost totally of fibrous cement (commonly replaced and/or recrystallized), in which grains or in-situ biogenic material does not constitute a framework (*Wright 1992). Describes a diagenetic fabric! Coverstone – Carbonate rocks dominated by in-situ tabular or lamellar organisms covering, protecting and stabilizing broken debris (*Tsien 1981, 1984). Includes a part of ‘bindstone’, sensu Embry and Klovan (1971). Pl. 78/4 Cruststone – Reef limestone consisting of densely packed organic crusts growing one upon the other (*Cuffey 1985). Pl. 51/7 Framestone – (a) Carbonate rocks containing in-situ massive sedentary fossils that built a rigid, three-dimensional supporting framework at the time of deposition (*Embry and Klovan 1971). (b) Reef frameworks dominated by the constructor guild (Fagerstrom 1987). Pl. 42/1, Pl. 79/1, Pl. 116/2, Pl. 149/2 Globstone – Reef carbonate rocks, consisting of colonies attached one upon the other (* Cuffey 1985). Lettucestone – Reef carbonate rocks, built up by calcified organisms characterized by lettuce-like growth forms (e.g. phylloid algae), aiding in the formation of a framework and intra- and interskeletal trapping of sediment. The descriptive name recalls the similarity of the growth form to that of composite plants, e.g. cabbages (*Cuffey 1985). Platestone – A limestone composed of an in-situ accumulation of algal plates, e.g. of phylloid algae (* Davies and Nassichuk 1990). Pl. 58/1 Shellstone – Carbonate rock characterized by shells (e.g. bivalves or brachiopods) cemented upon one another (e.g. oyster banks). Most shells are in contact; there is no or only little sediment between the shells (* Cuffey 1985). Pl. 87/4.
Examples for Framestone and Biocementstone
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Fig. 8.2. Framestone. More precise names are ‘branched coral framestone’ or ‘Retiophyllia framestone’ Densely spaced scleractinian coral colonies are still in life position. The corals form extended thickets. The coral limestone exactly fits the original definition of a framestone (in-situ sedentary fossils that built a three-dimensional framework) and also corresponds to the definition of framestones based on a guild character (dominance of the constructor guild). Note the difference in comparison with Pl. 41, also showing a framestone. Late Triassic: Adnet near Salzburg, Austria. Sawed quarry wall. The Adnet reef is a spectacular example of Triassic reefs (Bernecker et al. 1999). Width of the picture 80 cm.
of bindstone, framestone and bafflestone (Embry and Klovan 1971; see Pls. 41 and 42; Fig. 8.1), and followed by a wealth of special terms proposed in order to characterize different reef fabrics, reef rocks became more and more ambiguous and perplexing. Current reef research make it clear that the Embry and Klovan’s classification does not cover all the modes in which sessile organisms contribute to the formation of reefs. Hence, there is a great temptation to coin rock names describing all possible modes of reef building. Recognizing the complex manner in which organisms contribute to the construction of organic frameworks. Tsien (1981) and Cuffey (1985) proposed additional rock names, especially for ‘reef limestones’ (see Box 8.1). Tsien’s biocementstone and coverstone are useful and handy terms. They describe biological processes different from those addressed by Embry and Klovan: • Biocementstone (Fig. 8.3) describes a fine-grained texture, predominantly controlled by microbes and algae that colonize reef debris and produce microcrystalline cements as metabolic by-products. Biocementation is an important factor in the formation of mud
Fig. 8.3. Biocementstone. Many reef limestones are characterized by the abundance of micritic or sparry carbonate cements and the intimate association of cements with microbiota, e.g. calcimicrobes, sponges. The thin section shows sponge-like fossils (black) bound by submarine cements. Early Cambrian (Krol Formation): Lesser Himalaya, India. Scale is 5 mm.
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Bafflestones and Bindstones
Plate 41 Classification of Autochthonous Carbonates: Bafflestones and Bindstones Many textures used in the classifications of limestones are strongly biologically controlled. The term boundstone introduced by Dunham (1962) designates ‘carbonate rocks showing signs of being bound during deposition’ (e.g. reef limestones or stromatolites). This rather comprehensive term was later split up into several new terms, all of which demand a decision of the user about the potential control of sessile organisms on the formation of specific rock textures. Studying Devonian reef carbonates, Embry and Klovan (1971) proposed three terms to replace the name boundstone (see Fig. 8.5). The terms are widely used by geologists and palecologists interested in ancient reefs, despite considerable problems regarding the genetic meaning. The term boundstone should only be used if no decision on the mode of binding during deposition can be made. Framestones (see Pl. 41; Fig. 8.2) are characterized by massive in-situ fossils that constructed a rigid threedimensional framework during deposition. The in-situ fossils therefore form the supporting framework of the rock, with matrix material occurring in interstices of the fossils. Bindstones (–> 2) contain in-situ, tabular or lamellar fossils that encrusted and bound sediment during deposition. In bindstones the matrix, not the in-situ fossils, form the supporting framework of the rock. Bafflestones (–> 1) contain in-situ stalk-shaped fossils which trapped sediment during deposition by acting as baffles (i.e. reducing the rate of flow of water, thus causing deposition). The ingredients for recognizing a bafflestone are the presence of a large number of in-situ stalk-shaped fossils, and a good imagination on the part of the geologist. Plates 41 and 42 exemplify the application of the three terms and demonstrate the advantage of using a few more sophisticated reef rock terms. 1 Bafflestone. The picture displays a vertical to slightly oblique section of a tabulate coral colony belonging to the group Auloporida and the genus Multithecopora Sokolov. Species of this genus are characterized by colonies composed of cylindrical very thick-walled branching corallites that are not in contact and arranged parallel to each other. The corallites of the erect colonies of Multithecopora syrinx (Etheridge) are arranged subparallel to each other or grow in various directions. The corals occur in situ within mini-mounds found in a shelf sequence formed during a relatively sea-level highstand (Flügel and Krainer 1992). Maximum height of the mounds is 30 cm, maximum width 90 cm. Except for the tabulate corals, the mounds yield only a few foraminifera and ostracods. The fine-grained sediment of the mound is represented by (a) homogeneous micrite (M) within and between the auloporid colonies and in the center of the mounds, and (b) inhomogeneous calcareous siltite (S) between corallites and colonies, reflecting the allochthonous background sedimentation, also indicated by siltstone beds overlying and underlying the mounds. The micrite was formed prior to the siltite, which increases in frequency towards the mound margin. The delicate upwards branching growth form of the coral implies that these organisms were able to filter out carbonate mud, owing to baffling of gentle currents. A slight increase in water energy as indicated by changes in the orientation of corals enabled the siltite sediment to be trapped between the corals leading to a more rapid accumulation of the mound sediment as compared with the adjacent area. The name bafflestone is justified because of (a) the occurrence of stalk-shaped fossils in life position (as shown by the geopetals (GP), and (b) abundant fine-grained sediment baffled and trapped by the organisms. SMF 7-BAFFLESTONE. Middle Carboniferous (Kasimovian): Carnic Alps, Italy. 2 Bindstone. Vertical section of an irregularly laminated fenestral limestone. The limestone is part of a peritidal cyclic sequence consisting of subtidal, intertidal (this example) and supratidal members. The texture is characterized by alternating peloidal layers (PL), thin clotted micrite layers (ML) and intercalated horizontally elongated, laminate fenestral pores (F). The peloidal layers consist of fine and coarser micrite grains which are grain- or mud-supported. Mud-support occurs in association with micrite layers. Smaller micrite grains are well-rounded, larger grains are subangular. Some larger particles are composite grains (CG) and may represent ‘black pebbles’ (see Sect. 4.2.8). Spar-filled interspaces between smaller grains correspond to tiny tubes representing relics of microbial and algal filaments (white arrow). Many sparfilled fenestrae exhibit flat bottoms and slightly digitated roofs. Isolated grains within the fenestrae (black arrows) indicate a partial breakdown of the roofs. The vertical intrusion of a large clast (C) as well as the irregular shape of some fenestrae (F) indicate the soft consistence of the sediment. Despite the absence of in-situ lamellar fossils, this limestone is a bindstone because the grains were bound together by microbial activity concentrated in the micritic layers. The binding effect was enhanced by synsedimentary cementation in intergranular pores. SMF 21-BINDSTONE. Late Triassic (Pantokrator limestone, Norian): Didymi Mountains, Peloponnesus, Central Greece.
Plate 41: Autochthonous Carbonates: Bafflestone and Bindstone
345
Autochthonous Carbonates: Framestone
346
mounds. It is common in protected environments and cryptic habitats. • Coverstone refers to a texture characterized by insitu sedentary tabular or lamellar reef organisms covering, protecting, and stabilizing broken debris (Pl. 78/ 4). This type was included by Embry and Klovan in the bindstone category together with reef carbonates characterized by encrusting organisms that bind debris together, forming a secondary reef framework. For the needs of a field geologist, Cuffey’s (1985) expansion of the differences in the relationship of reef organisms and texture may appear to be too rigorous. The terms, however, have the advantage of reminding us of the complexity of reef-building, which could not be described by the use of more widely defined terms such as ‘boundstone’ or ‘framestone’. Cuffey’s terms, therefore, will be especially valuable for studies dealing with the biological controls on reef building.
ern reefs are gorgonian fans, alcyonarian whips, or the bladelike coenostea of hydrocorallines. Sea grass is also involved in baffling and trapping processes by its blades or roots. However, geologists may be baffled by bafflers. Ancient organisms believed to have acted as bafflers comprise sponges, bryozoans, bivalves and algae as well as some corals (Fagerstrom 1987). It is virtually impossible to demonstrate that the original organisms baffled sediment; therefore, baffling is inferred from functional morphology (erect and dominant upward growth; flexible, semirigid or rigid skeletons) and close intercolonial and intracolonial spacing. But both criteria may be misleading (Flügel and Krainer 1992). If organisms are attributed to the baffler guild, additional criteria should be considered (e.g. composition of the fine-grained sediment between and around the ‘potential bafflers’, see Pl. 41/1; role of extrinsic factors).
Do rock names assist in reliable environmental interpretations ? The question can be discussed by using the term bafflestone, which implies the process of baffling. ‘Baffling’ refers to current reduction and gravityinduced dropping of sediment grains to the substrate. In reef limestones baffling organisms are rarely found in their growth position. Examples of bafflers in mod-
Controls on growth forms used in limestone classifications: Growth forms of organisms used to define textures of reef carbonate types are commonly thought to be related to water energy. But what controls the organic growth types used in textural carbonate classifications? Common reef fabrics may be placed into four groups, characterized by the morphology of the domi-
Plate 42 Classification of Autochthonous Carbonates: Framestones Framestones (a subdivision of boundstones) are defined by the occurrence of sessile benthic fossils that are densely spaced and preserved in life position. The morphology and distribution of the fossils should fit into an imaginary three-dimensional organic framework. Organisms contributing to the formation of framestones have changed during the Phanerozoic and include corals (Fig. 8.2), coralline sponges (this example), stromatoporoids, rudist bivalves, and calcareous red algae. Framestone. The thin section of a reef limestone, cut obliquely to the growth direction of most of the sessile organisms, exhibits fossils, cement-filled interspaces and areas with fossils and micrite. Most of the fossils are coralline sponges, except for a few mollusk shells (MS) and crinoid columnals (C). Sponges include various ‘sclerosponges’ (S), inozoid sponges (IS) characterized by the absence of segmentation, as well as segmented sphinctozoid sponges (SS). The growth forms of these sponges correspond somewhat to densely spaced bundles of cylinders united by organic encrustations. Encrusters are sclerosponges, Archaeolithoporella (A) and Tubiphytes (T) - two microfossils of unknown systematic position (see Pl. 98/8). Isolated sponges as well as the concentrations of fossils were initially bound by microbial crusts and marine carbonate cement. The originally aragonitic cements and the biogenic encrustations formed a rigid, highly porous framework structure (accentuated by white spot lines). Large voids between the organic framework were partly occluded by marine-phreatic, fibrous isopachous cements (IC) and botryoidal cements (BC) as well as sparry calcite mosaics (SCM). The remaining intergranular voids were occluded by large calcite crystals formed in a meteoric-vadose zone during exposure of the reef. The composition and diagenesis of these reef rocks are similar to Late Permian reefs in Texas and New Mexico (cf. Pl. 145), both of which are very important hydrocarbon reservoirs. The name framestone is substantiated by the existence of an organic framework made by in-situ fossils and forming a rigid structure that was strengthened by coeval marine cementation. A massive growth form, as indicated in the original definition of the name, is not compulsory, because small densely growing sessile organisms can build frameworks as well. This sample comes from one of the very last reefs formed prior to the huge Permian-Triassic mass extinction. SMF 7FRAMESTONE. Uppermost Permian (Changsingian): Jiantinba, western Hubei, South China. Crossed nicols. Picture courtesy of Fan Jiasong (Beijing)
Plate 42: Autochthonous Carbonates: Framestone
347
Dunham and Folk Classifications
348
nant members of the community (Fig. 8.1). In general, the distribution of growth forms of benthic shallowwater organisms reflect patterns of functional significance. Many authors underline the adaption of growth habits of reef organisms to water energy, e.g. strength, type, and duration of currents and waves. Specific growth forms are used as indicators of the existence of particular or dominating energy stages (see Sect. 12.1.1). There is, however, growing evidence that nutrient levels and resulting predator patterns as well as substrate conditions are major and probably dominating controls on the biogenic fabrics that are used in more sophisticated subdivisions of reef limestones (Wood 1995).
vestigations of cores, and in laboratory studies. The use of Folk’s classification is more restricted to laboratory studies based on thin sections or peels.
8.3.1 Prerequisites Before classifying anything, it is necessary to determine what constituents occur (grain categories, matrix and cement types). The second fundamental principle concerns the question of whether the constituent grains are grain- or mudsupported (see Pl. 43).
8.3.2 Original and Expanded Dunham Classification
8.3 Classifications Based on Depositional Texture The most widely used classifications are those of Dunham (1962) and Folk (1959, 1962). Dunham’s classification can equally well be applied in the field, in in-
Figure 8.4 shows the original version of the Dunham classification, Fig. 8.5 the expanded and revised versions (Embry and Klovan 1971; Wright 1992). The criteria of the rock categories are depicted and discussed in Pl. 43 and Pl. 44.
Carbonates Dunham (1962) Groundmass: Fine carbonate matrix
sparry cement +spar
Matrix-supported Grains: < 10% > 10% MUDSTONE
Folk (1959, 1962) Allochems: < 1% 1-10% fossiliferous MICRITE
Bioconstruction
Grain-supported
WACKESTONE
10-50% sparse
PACKSTONE
packed BIOMICRITE
GRAINSTONE
> 50% poorly washed BIOSPARITE
BOUNDSTONE
BIOLITHITE
Terrigenous Matrix-supported Sand: < 10% 10-25% sandy MUDSTONE
Grain-supported > 25% WACKE
SUBWACKE SANDSTONE
ARENITE
Fig. 8.4. Fossiliferous limestones classification after Dunham (1962) and Folk (1959, 1962). Both classifications distinguish allochthonous limestones (mudstone, wackestone, packstone, grainstone) and autochthonous limestones (here called boundstone or biolithite). Limestones whose components were deposited as discrete grains are grouped according to mud-support or grain-support and the abundance of grains. The Dunham classification stresses the depositional fabric, the Folk classification tries to evaluate hydrodynamic conditions. Both classifications consider the dominating groundmass types. Note the divergent categorization of limestones with abundant, densely packed grains and a fine-grained matrix or mixed fine-grained/ sparry groundmass (packstone).
Dunham Classification of Limestones
349
CLASSIFICATION OF LIMESTONES (DUNHAM 1962) DEPOSITION TEXTURE RECOGNIZABLE
Original components not bound together during deposition Contains mud Lacks mud (particles of clay and fine silt size) and is Mud-supported Grain-supported grainsupported
less the 10% grains
more than 10% grains
MUDSTONE
WACKESTONE
PACKSTONE
DEPOSITIONAL TEXTURE NOT RECOGNIZABLE
Original components were bound together during deposition as shown by intergrown or lamination contrary to gravity, sediment-floored cavities that are roofed over by organic or questionable organic matter and are too large to be interstices
GRAINSTONE
CRYSTALLINE CARBONATE
(Subdivide according classification designed to bear on physical texture or diagenesis)
BOUNDSTONE
EXPANDED CLASSIFICATION (EMBRY and KLOVAN 1971) ALLOCHTHONOUS LIMESTONE ORIGINAL COMPONENTS NOT ORGANICALLY ORIGINAL BOUND DURING DEPOSITION
Less than 10% > 2 mm components contains lime mud (< 0.03 mm) Mud supported less than greater than 10% grains 10% grains ( > 0.03 mm and < 2 mm) MUDSTONE
WACKESTONE
no lime mud
Greater than 10% > 2 mm components
Grain-supported
PACKSTONE
AUTOCHTHONOUS LIMESTONE COMPONENTS ORGANICALLY BOUND DURING DEPOSITION
GRAINSTONE
Matrixsupported
> 2 mm component supported
FLOATSTONE
RUDSTONE
by organisms which build a rigid framework
encrust and bind
act as bafflers
BOUNDSTONE FRAMESTONE BINDSTONE BAFFLESTONE
REVISED CLASSIFICATION (WRIGHT 1992) DEPOSITIONAL
Mixed supported clay and silt grains) < 10% > 10% grains grains grains
CALCI-MUDSTONE
BIOLOGICAL
Grain-supported with matrix
no matrix
WACKEPACKGRAIN-STONE -STONE -STONE FLOATSTONE RUDSTONE ------------------------------
DIAGENETIC
In situ organisms rigid organisms dominant FRAME-STONE
encrusting organisms binding acted to organisms baffle BOUND-STONE
Grains > 2 mm
BAFFLE-STONE
Non-obliterative main component in cement CEMENT-STONE
many grain most grain contacts ascontacts microare microstylolites stylolites CONDENSED FITTED GRAINSTONE
Obliterative crystals > 10 µm
SPARSTONE Crystals < 10 µm MICRO-SPARSTONE
Fig. 8.5. Original, expanded, and revised Dunham classification. Embry and Klovan (1971) introduced the size aspect and distinguished grains smaller or larger than 2 mm. The new categories floatstones and rudstones correspond to wackestones and grainstones and packstones, respectively. Other changes concern the subdivision of the boundstone category (see Sect. 8.2). Wright (1992) suggested terms describing diagenetic changes of depositional fabrics. The author emphasizes that the fine-grained matrix commonly called ‘mud’ is not necessarily identical with micrite, but corresponds to a matrix consisting of clay-sized and silt-sized constituents. Note that in this scheme the term mudstone is replaced by calcimudstone and bindstone by boundstone owing to slight semantic problems.
8.3.2.1 Concepts The original classification includes five textural classes. Two major groups are distinguished – (1) carbonates whose original components were organically bound together during deposition (boundstones), and (2) carbonates whose original components were not organically bound. The second group is subdivided according to mud-support (mudstone and wackestone) or grain-support (packstone and grainstone). The subdivision of mud-supported rocks is based on the < or > 10% grain bulk boundary. These rock names can be combined with grain type names in various ways: Ostracod mudstone (mudstone containing < 10 total rock volume % ostracods), crinoid packstone (packstone containing a high percentage of closely-packed crin-
oids), ooid grainstone (grainstone containing predominantly ooids). Dunham proposed the notation ‘lime’ or ‘dolomite’ for characterizing the mineralogy, but only the reference to dolomite rocks (e.g. dolomite mudstone) appears necessary. The classification is highly flexible. The rock names should be complemented by all relevant field or laboratory observations. These can be sedimentary structures (e.g. medium-bedded, cyclic bioclast grainstones) and specific textural criteria (Stromatactis-bearing microbial mudstone). Limestones with small-scale mixtures of fabric types may be described by combinations of names: Peloid wacke/packstone corresponds to a sample with > 10% peloids and exhibiting wackestone and, to a lesser degree, packstone textures. Such close occurrence of wackestone and packstone textures may be caused by burrowing.
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8.3.2.2 Discussion Dunham’s rock names are generally accepted despite some problems in the difficulty of unequivocally differentiating of the textural categories and their inferred environments. The greatest problem is that the classification strongly suggests a ‘depositional’ character of the textures. It must be kept in mind, however, that limestones exhibiting mudstone, wackestone, packstone and even grainstone fabrics could also be products of diagenetic processes, such as micritization, cryptocrystalline calcite cementation or neomorphism.
Grain-Support and Mud-Support
What do mudstones and wackestones indicate about paleoenvironmental conditions? Mudstones are per definition ‘muddy carbonate rocks containing less than 10 per cent grains’. Grains and grain-producing organisms are rare, possibly reflecting unfavorable environmental conditions. The sparsity of grains, however, could be a matter of scale and method: Many pelagic mudstones correspond to densely packed wackestones or even to packstones if studied with the aid of ultra thin-sections or the stereoscan electron microscope. Linking the deposition of mudstones and wackestones to low-energy settings may be a misinterpretation, be-
Plate 43 Allochthonous Carbonates. Basic Depositional Fabrics: Grain-Support and Mud-Support Major criteria for classifying allochthonous limestones are the depositional fabric, the abundance of various grain categories as discussed in Chapt. 4, and the type of fine-grained matrix and carbonate cements associated with the grains. The principal feature of the classification proposed by Dunham (1962) is the distinction between grain-support (particle-support) and mud-support (matrix-support) describing the basic depositional fabrics: Grain-support (–> 1) is defined as a fabric with little or no mud and abundant sand-sized grains that are in three-dimensional contact and able to support each other. The fabric is comparable to that of grains in a pile of sand. Mudsupport (–> 2) is defined as a fabric whose sand-sized grains (at least 10% of the total rock volume) ‘float’, i.e. are embedded in, and supported by the muddy matrix. The fabric resembles pebbles within a mud puddle. Distinguishing grain-supported packstone from mud-supported wackestone may be difficult, because thin sections of both rock types can show grains ‘floating’ within a matrix and apparently not in contact. Because of the two-dimensional aspect of thin sections, irregularly shaped grains of a grain-supported fabric often do not appear to be in contact along every surface. Hence the point contact of grains is not necessarily a decisive factor in defining a ‘packstone’. The degree of grain packing, grain density, and number of grains can be measured in thin sections using various methods (see Sect. 6.1.1.1). However, the number of grains alone is not conclusive for distinguishing between grain- and mud-support because even 20-30% of irregularly shaped large grains may form a self-supporting framework. Helpful but not unequivocal criteria in recognizing grain-support are: (a) lack of carbonate mud, (b) floored interparticle pores, (c) embowed grains, (d) close and overly close packing, and (e) abundant interparticle cement. The grain-/mud-support concept is fundamental in interpreting depositional controls (e.g. hydrodynamic conditions). The distinction between grain-supported and mud-supported fabrics reflects major differences in primary porosity types (Sect. 7.3) and, in turn, the potential of limestones as reservoir rocks (Sect. 17.1.5.2). 1 Grain-support fabric, indicated by absence of mud, close packing of grains, and abundance of carbonate cement in interparticle pores. Most particles are aggregate grains, followed by some micritized bioclasts (foraminifera, F; mollusks, M), very small gastropods (G) and peloids. The grains are equally sized and very well sorted (standard deviation 0.35), see Fig. 6.3. The calcite cement in the interparticle pores appears white. The fine-grained lithoclastic grainstone forms a capping bed overlying a rudist patch reef. The capping beds correspond to mobile midshelf grain flats. SMF 17. MidCretaceous (Edwards Limestone, Comanche Formation, Albian): Belton, central Texas. 2 Mud-support fabric indicated by grains ‘floating’ in lime mud. Skeletal grains are echinoderms (E), thin-shelled bivalves (filaments, FI), double-valved bivalves (B) and foraminifera. Note that geopetally filled cavities (black arrows) form a distinct horizon, possibly indicating shallow-burial dissolution. Aragonitic shells are dissolved and infilled with sediment (ammonite mold: white arrow). Original calcitic grains (echinoderms, filaments) are well preserved. The sample would be classified as echinoderm floatstone according to the echinoderm fragments, which are larger than 2 mm and comprise more than 10% of the rock. This name would fail to document the conspicuous bimodality of the sample and specific composition of the matrix. The total rock fabric is better described as ‘crinoid floatstone with a fine-grained skeletal wackestone matrix rich in filaments’. The red nodular limestones originated in a moderately deep pelagic basin. Late Triassic (Hallstatt limestone, Carnian): Rötelstein, Styria/Austria.
Plate 43: Grain-Support and Mud-Support
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cause binding organisms (e.g. microbes) may also fix mud within higher-energy settings. What is a ‘packstone’? The term packstone is ambiguous with regard to its genetic implications. Packstones are characterized by grain-support as well as by mud matrix, indicating different energy levels during deposition. How can this contradiction be explained? How can the matrix be introduced into the grain-supported fabric? Do all packstones represent depositional textures? • Mud from low-energy deposits may filter into empty pores of underlying, grain-supported high-energy deposits or settle out of the water flowing through pores.
Dunham’s Classification
Evidence of this can be seen in geopetal fabrics. These packstones may exhibit a micritic matrix occurring together with sparite formed as cement within the remaining open pores. • Precipitation of microcrystalline carbonate cement within intergranular pores may result in a packstone fabric. This is common in calcretes or beachrocks but may also occur in marine carbonates. • Compaction of wackestones, facilitated by dewatering may produce a ‘diagenetic packstone’ fabric (Shinn and Robbin 1983). • Uncemented grains may be welded together during pressure dissolution resulting in a fabric that is commonly classified as ‘packstone’, even though ‘matrix’ is very rare (Pl. 47/1).
Plate 44 Dunham’s Classification of Carbonate Rocks: Mudstone, Wackestone and Packstone Mud- and grain-supported limestones (Pl. 43) are subdivided into textural categories according to the abundance of grains (Fig. 8.4). The complete rock name requires the combination of the name of the category with adjectives referring to grain kinds. Texture, mineralogy and frequency criteria can be included in the rock name. Mudstones are muddy carbonate rocks containing less than 10% grains measured as grain-bulk percent. The name is more or less synonymous with calcilutite. Wright (1992) replaced the name mudstone by calcimudstone because mudstone is also used for rocks consisting of silicate silt and clay material. This term refers to ‘lithified material composed of greater than 90% by volume, silt and clay-grade calcite’. It is not synonymous with ‘micrite’. The usual interpretation of mudstones is that they represent deposition of fine-grained sediment under lowenergy conditions allowing carbonate mud to settle in calm and quiet waters. However, it is now clear, that finegrained carbonates can be also formed by precipitation of microcrystalline carbonate cements (Reid et al. 1990). Examples are known from microbial carbonates (see Pl. 8), terrestrial carbonates (caliche, Pl. 128/6) as well as seep and vent carbonates (Pl. 148). Mudstone is a useful term in the field and also for thin-section studies. It should be noted, however, that the ultrastructure of many mudstones corresponds to extremely fine-grained bioclastic or peloidal packstones and wackestones. Wackestones are mud-supported carbonate rocks containing more than 10% grains. High variations in abundance and type of grains call for a breakdown of this category into informal subcategories (e.g. packed or sparse wackestone, > or < 50% grains; –> 2 and Pl. 118/1). Packstones are grain-supported muddy carbonate rocks exhibiting features pointing to deposition in agitated waters (grain-support) and criteria pointing to quiet-water deposition. Packstones have different origins including infiltrating of previously deposited mud-free grainstones, prolific production of grains in calm waters, mixing of sediment by burrowers, partial leaching of mud or compaction and dewatering of original wackestones (Shinn and Robbin 1983). A rather common feature of packstones in thin sections is the co-occurrence of mud (micrite) and irregularly distributed areas with sparry calcite. 1 Mudstone. Radiolaria-spicula mudstone (MS). Note the homogeneous texture of the matrix. The matrix is a fine-bioclastic micrite, corresponding to type 7 in Fig. 4.1. A different fabric is shown near the bottom of the picture (packed radiolaria-spicula wackestone; WS). The sample comes from well-bedded lithographic limestones formed in restricted parts of an open shelf environment. SMF 1. Late Jurassic (Vaca Muerta Formation, Tithonian): Zapala, Neuquén Basin, Argentina. 2 Wackestone. Packed echinoderm wackestone. Most grains are crinoid plates. White arrows point to ophiurid elements, the black arrow to a foraminifera. The echinoderm grains are very worn. Straight elements are bivalve shell debris. The echinoderm limestone is part of a transgressive unit formed after drowning of Triassic reefs (Böhm et al. 1999). SMF 9. Early Jurassic (Hettangian): Adnet near Salzburg, Austria. 3 Packstone. Bioclastic coral-algal packstone composed of corals (C), solenoporacean red algae (Pseudochaetetes asvapatii Pia; PS), bivalve shells, and gastropods (G). Arrows point to sediment-binding squamariacean red algae (Pseudolithothamnium album Pfender). The infilling of mud and fine skeletal debris in interstices between larger grains indicates grain-support. SMF 8. Early Tertiary (Kammbühel limestone, Paleocene): Mooshuben near Mariazell, Styria/Austria.
Plate 44: Mudstone, Wackestone, Packstone
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Are grainstone fabrics infallible indicators of highenergy conditions near the depositional interface? Many grainstones represent deposition in current or wave-influenced settings where mud-grade sediment does not accumulate. It should be kept in mind that not all grainstones indicate the existence of high-energy conditions and that some ‘grainstones’ may be the result of specific depositional or diagenetic processes: • Some grainstone fabrics are the product of a local accumulation of grains that have settled out of a flow of a mud-grain mixture while the mud is washed away. • Other grainstones develop from the winnowing of mud from previously deposited grainy/muddy sediment. • A too rapid production and accumulation of grains will prevent the contamination by mud.
Dunham’s Classification
• The desiccation of subaerially exposed carbonate muds can produce ‘diagenetic grainstones’. This process, called ‘grainification’ (Mazzullo and Birdwell 1989; Wright 1990) is common in peritidal, paludal, and pedogenic carbonates. • Diagenetic grainstones are also produced by pressure solution during burial diagenesis, resulting in ‘condensed grainstones’ and ‘fitted grainstones’ (Wright 1992). Floatstone and rudstone fabrics may be pitfalls in environmental interpretation: • A coarse-grained texture can be better seen in larger rock samples and thin sections than in small ones, due to the variability in grain sizes of transported bioclasts.
Plate 45 Dunham’s Classification of Carbonate Rocks: Grainstone, Floatstone and Rudstone The terms floatstone and rudstone were proposed for equivalents of wackestones, packstones and grainstones differing from these rock types in their larger mm- to cm-sized grains. Distinguishing floatstone/wackestone and rudstone/grainstone is particularly meaningful for investigating reef limestones and allochthonous limestones deposited on slopes. Grainstones are grain-supported and mud-free carbonate rocks and consist of skeletal and non-skeletal carbonate grains. The absence of mud has various causes: Deposition of grains in high-energy environments (e.g. in intertidal and shallow subtidal environments), rapid accumulation of grains allowing no coeval mud sedimentation (e.g. turbidites), deposition in current-controlled environments, or winnowing of mud from previously deposited grain/mud mixtures. Grainstones are highly variable with regard to grain type, shape, size, and sorting. Significance: Grainstones are common rocks in platform and ramp carbonates. They are usually related to the locus of wave energy absorption such as shorelines, shoals or shelf breaks. Here, grainstones form thick accumulations at the outer shelf margins or in inner ramp settings. ‘Bank-margin sands’ occur in tidal bar belts, tidal deltas, marine sand belts, back-reef areas, on beaches and in subaerial dunes. Note that accumulations of grainstones can also originate below the wave sweep base owing to current effects. Owing to their high interparticle porosities and the possibility of additional secondary moldic porosity caused by dissolution of grains (Pl. 29/3, 4), grainstones are of particular interest as reservoir rocks. Floatstones are matrix-supported carbonate rocks yielding more than 10% grains larger than 2 mm. This 2 mm boundary creates difficulties in applying the term to fossiliferous limestones or limestones with oncoids because the size of skeletal grains and oncoids can only reflect growth stages. The use of a strict terminology would put a gastropod limestone yielding adult snails into the floatstone category and limestones containing the same but juvenile snails into the wackestone category. Note that floatstone and rudstone were originally proposed for reef carbonates and carbonate breccias. The ‘matrix’ of floatstones does not necessarily correspond to micrite, but often exhibits fine-grained textures that must be described separately (–> 2). Rudstones are grain-supported carbonates rocks containing more than 10% grains larger than 2 mm. Many breccias belong in this category. Rudstones and floatstones must be further characterized by compositional and textural criteria (–> 3). Formation of rudstones needs erosion and transport. Erosion can be triggered by shallowwater settings allowing destruction by storms. Slopes are common depositional settings of rudstones. 1 Grainstone. Medium-sorted oolitic grainstone. Note the difference in ooid sizes, indicating some transport. SMF 15-C. Late Triassic (Carnian): Obir, Carinthia/ Austria. 2 Floatstone. Bioclastic coral (C)-gastropod (G) floatstone with a fine-grained calcimudstone matrix. The originally aragonitic corals are strongly recrystallized. Arrow points to still preserved septae. SMF 8. Late Triassic (Kössen Formation): Allgäu, southern Germany. 3 Rudstone. Poorly sorted lithoclastic rudstone. The large coated bioclasts are sclerosponges (S). Lithoclasts (L) correspond to reef debris deposited on reef slopes and redeposited as part of a megabreccia. Interparticle pores are filled with submarine and burial carbonate cements. SMF 6. Late Permian: Sosio, western Sicily, Italy.
Plate 45: Grainstone, Floatstone, Rudstone
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• The terms have been proposed to characterize transported and resedimented fossils as illustrated by fragments of stromatoporoids. The original definition used was not intended for classifying limestones with large components that owe their size to biological growth processes (e.g. oncoids or colonial fossils). A distinction in the nomenclature of limestones with double-valved bivalves which are either smaller or larger than 2 mm, will separate different biological growth stages whose size differences are not controlled by transport. This problem is commonly overlooked in the current use of the terms floatstone and rudstone. This is not of great importance as long as the texture types are not interpreted with respect to possible water energy levels or other environmental factors. • Coarse-grained bioclastic particles such as phylloid algal plates or Halimeda blades can accumulate, producing a rudstone texture. This term, however, is inappropriate, because it infers transport and allochthonous sedimentation. The algal elements, however, are essentially in-place accumulations. Therefore, the term platestone (Davies and Nassichuk 1990) is more appropriate.
8.3.2 Original and Expanded Folk Classification Folk’s classification is a useful and practical method for describing carbonate rocks, particularly in thin sections and peels.
8.3.2.1 Concepts The basic philosophy of Folk’s classification is that carbonate rocks are similar to siliciclastic rocks in their mode of deposition, because their textures are both controlled largely by the water energy in the depositional area. In calm water with only sluggish currents, carbonate mud (represented as micrite) with or without grains is predominantly deposited. By contrast, vigorous water energy hinders deposition of fine-grained material, thus favoring the sedimentation of winnowed sands with large amounts of pore space that is later filled with sparry calcite (sparite). The most important environmental break is between limestones with a lime-mud matrix and those with calcite cement, because this should reflect the point where water energy becomes turbulent enough to wash out (winnow) the lime mud, keep it in suspension and carry it into lower energy zones (Fig. 8.6).
Folk’s Classification
Folk (1959) distinguished three end members: (1) discrete carbonate grains (‘allochems’), (2) microcrystalline calcite matrix, and (3) sparry calcite, regarded as pore-filling cement. Using the relative proportions of these three end members, the author distinguished ‘sparry allochemical rocks’ (consisting chiefly of grains cemented by sparry calcite cement), ‘microcrystalline allochemical rocks’ (consisting of microcrystalline matrix and various amounts of grains) and ‘microcrystalline rocks’ (consisting of microcrystalline calcite and few or no grains). Further subdivision of these major groups is based on the ratio of grains to sparite to micrite as well as the type of grains. Quantitative boundaries include grain bulk percentages (> or < 10% grains with differentiation between 1–10% and < 1%) as well as proportions of special grain types related to the total of all grains. The Folk classification was designed to cover limestones, partly dolomitized limestones, ‘primary’ dolomites and replacement dolomites. Limestones were subdivided into eleven lithotypes, eight of them comprising allochemical rocks, two orthochemical rocks and one autochthonous reef rock. The kind of grains (only intraclasts, ooids, fossils and pellets were differentiated) and the ratio of grains, sparite and micrite are used in classifying and naming rock types. A strict use of Folk’s classification system requires consideration of several points when establishing rock names: Non-genetic meaning of grain names: Despite its genetic definition, Folk (1962) regarded the term intraclast for purposes of nomenclature ‘as a broad class term without specifying the precise origin’. Using the term in a rock name, says nothing per se about the origin of the grain or about the environment. The term may refer to mud clasts reworked and transported by tidal currents or hurricanes as well as synsedimentary cemented grain aggregates (grapestones), or merely fragments of weakly lithified sediment that have been retectured by burrowing. Similar to intraclasts, the use of the term pellet has no genetic implications. The distinction of pellets (i.e. peloids, see Sect. 4.2.2) and intraclasts is conventional and purely descriptive. Strict ranking of grains (allochems): Basically, Folk’s rock names consist of two prefixes. The first (abbreviated by the prefixes intra-, oo-, bio- or pel-) refers to the grain type, the second to the groundmass (sparite or micrite), see inset on Pl. 46). But many limestones exhibit various amounts of different grain types. Folk introduced a ranking system and regarded intraclasts as the most important carbonate grains ‘because of their implication of shallow water, lowered wave
Folk’s Classification
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Fig. 8.6. Concept and textural spectrum of the Folk classification. Modified from Folk (1962). Note the increasing textural maturity from left to right. Dismicrite is omitted in this figure.
base, or possible tectonic uplift’. Ooids are the next important group, followed by fossils and peloids. This ranking implies the knowledge of quantitative data: In ‘intraclastic limestones’ intraclasts must comprise more than 25% of all grains or in ‘oolitic limestones’ there must be more than 25% ooids regardless of the percentages of other grain types, which may be more frequent than ooids. However, more than one name-defining grain category may be important for rock naming. ‘Fossiliferous oosparite’, for example, would indicate that at least 25% of the grains are ooids, but fossils also occur in significant quantities. ‘Fossil-bearing oosparite’ would indicate that fossils occur, but that they are not frequent. These notations, of course, are not very precise, because both examples could contain between 1 and 74% fossils whose existence is lost in the name ! Grain sizes: Folk distinguished allochemical rocks with larger and smaller grains, using a median size limit of > or <1 mm in order to separate, for example ‘biosparrudite’ from ‘biosparite’. This boundary differs from the common rudite/arenite boundary at 2 mm. A more sensitive subdivision (Folk 1962), reflecting different depositional environments, is based on the relative proportion of grains and carbonate muds, the sorting of grains, and the rounding of grains. This expansion of the original classification of 1959 describes eight depositional stages, characterized by increasing textural maturity (Fig. 8.6): 1. Pure micrite with almost no grains and dismicrite; 2. Micrite with up to 10% grains, e.g. fossiliferous micrite or peloid-bearing micrite; 3. ‘Sparse’ biomicrite (or intramicrite, pelmicrite) with 10–50% grains. Particles are scattered throughout the matrix and are loosely packed; 4. ‘Packed’ biomicrite (or intramicrite, etc.) with more than 50% grains.
Particles appear densely packed; 5. Poorly washed biosparite (or intrasparite) with about subequal amounts of micrite and sparite; 6. Unsorted biosparite, etc. The poorly sorted grains are cemented by sparry calcite; micrite is rare; 7. Sorted biosparite. The grains are wellsorted but only slightly abraded and rounded; 8. Rounded biosparite. The majority of grains is wellrounded. These ideal textural groups reflect a shallowing-upward sequence from ‘deeper’ low-energy sediments (1 to 4) via a transitional zone (5) characterized by changing water energy to ‘shallow’ high-energy sediments (6 to 8). Textural inversions (transport of carbonate grains to areas which are not the site of their origin) could strongly modify this scheme. ‘Oomicrites’ may be the result of redeposition of high-energy ooids within low-energy environments (Pl. 46/3); bioturbation may cause a mixture of textural types (Pl. 19). Folk’s classification is open to all individual qualifications that add adjectives or field and laboratory criteria to the basic rock name. These may include: • features observed in the field such as rock color, sedimentary structures and physical properties (e.g. ‘white, porous cross-bedded oosparite’, ‘greenish, friable burrowed micrite’); • terrigenous admixtures (e.g. ‘sandy glauconitic foraminiferal biosparite’, ‘argillaceous nodular intramicrite’); • textural criteria such as grain size, sorting and rounding (e.g. ‘poorly sorted, densely packed biomicrite’, ‘coarse-grained rounded biosparite’). Modified and expanded classification (Strohmenger and Wirsing 1991) Supplements to and modifications of the classification pertain to the grain types and the boundary limits
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Plate 46
Folk’s Classification
Folk’s Classification of Carbonate Rocks
The classification proposed by Folk (1959) for non-reef carbonates is based on the ratio grains : micrite : sparite. Required for using this classification is (a) separation of carbonates with and without a micrite matrix – see figure on this page, (b) differentiation of fundamental grain types, and (c) frequency and proportion of grain types. The names of grains of the rock name are arranged in a specific sequence (intraclasts, ooids, skeletal grains/peloids). Although the classification is widely used, the names are often defined in a way which does not follow the strict rules proposed by Folk starting with a prefix of the less frequent grain type and increasing to the most abundant type or vice versa). The extension of Folk’s classification by Strohmenger and Wirsing adds oncoids and extraclasts as separate categories to the rock names, the change in frequency boundaries for grains, and the incorporation of all grains in the rock name as long as they comprise > 20% (or more than 40 % for skeletal grains/peloids). The modification is useful, because it results in reproducable rock names allowing statistical treatment. 1 Micrite. Corresponding names are mudstone and calcimudstone. Folk’s original definition characterizes ‘microcrystalline calcite ooze’ composed of 1-4 µm-sized calcite crystals, ‘generally subtranslucent with a faint brownish cast in thin section’ and dull and opaque in hand specimen. These criteria are typical for many lithographic limestones and also for the type of these extremely fine-grained rocks, the Solnhofen limestone shown in this figure. Lithographic limestones represent a particular type of micritic carbonate. They are formed under a wide range of environmental conditions, including lagoons (e.g. Solnhofen,), open shelf and also lacustrine settings (Oost et al. 1994; Swinburne and Hemleben 1994). The Solnhofen limestone is famous for its unusually well-preserved whole body fossils and the fact, that the fine- and even-grained rock has long been used in lithography, buildings, and in experimental work on deformation fabrics of carbonates affected by high strain rates (Casey et al. 1998). Some of the Solnhofen micrite consists of coccoliths, some of calciodinoflagellates embedded in minimicrite, and still others (this sample) are composed of very fine bioclastic detritus (arrows point to radiolarians). Fourier spectral analysis of millimeter-scale lamina couplets consisting of a carbonate-rich lamina and a relatively carbonate-poor lamina indicates a short-time cyclic climatic control on the deposition of this carbonate mud (Park and Fürsich 2001). Late Jurassic (Tithonian): Solnhofen, Bavaria, Germany. 2 Dismicrite. Corresponding name: fenestral bindstone (see Pl. 41/2). ‘Disturbed micrite’ is a sackname for micritic limestones containing irregular patches of sparry calcite representing spar-filled openings of diverse origin (see 5.1.5). Note geopetal infilling of some voids (arrows). Intertidal part of a loferite sequence (see Pl. 140). SMF 21. Late Triassic (Norian): Gaissau, Northern Alps, Austria. 3 Oomicrite. Corresponding name: oolitic lithoclast wackestone. Grains are ooids and intraclasts. Taking into account the frequency of ooids (> 25% of the grains) and the micritic matrix, the rock is called oomicrite. Because this name excludes the rather conspicuous intraclasts (I), some authors would prefer the name oo-intra-micrite or even, loosely, intra-oomicrite, emphasizing reworking of the sediment as indicated by the large intraclasts. The sample indicates transport and redeposition of ooid shoals in protected ramp settings (see Pl. 136). Lowermost Cretaceous: Subsurface, Kinsau well, southern Bavaria, Germany. 4 Biomicrite. Corresponding name: Red-algae bearing echinoderm packstone or Komia-bearing echinoderm packstone. The reference to Komia is important because of the facies-diagnostic value of Late Paleozoic red algae (see Pl. 108). All the grains are bioclasts, usually echinoderms associated with a few shells (S) and a red alga (Komia, K; see Pl. 56/6). There are >25% sandsized skeletal grains associated with a micritic matrix. Hence, the rock is a biomicrite. Packing is conspicuous, hence the name ‘packed echinoderm biomicrite’ is more appropriate (see Fig. 8.6). Pennsylvanian (Honnacker Trail Formation): Limestone Ridge, Goosenecks, San Juan, Utah. 5 Intrasparrudite. Corresponding name: Poorly sorted lithoclast rudstone. Grains are intraclasts consisting of pelmicrite, cemented by various generations of submarine calcite spar. Because the grain size (> 2 mm) is included in the rock name the prefix ‘rudite’ is a part of the name. Slope environment. Early Jurassic: Adnet near Salzburg, Austria. 6 Oosparite. Corresponding name: Poorly sorted ooid grainstone. Grains are large and small ooids (O; 65%), bioclasts (25% bryozoans, B; foraminifera, F; shells, S), and cortoids (10%; C). Strohmenger and Wirsing would classify this sample as oobiosparite (see Fig. 8.7). Early Cretaceous: Subsurface, Ostermünchen well, Bavaria, Germany. 7 Classification exercise. Try to find names for this sample, following the proposal of Folk and Dunham. For explanation see Appendix. Middle TriBasic concept of the Folk classification. assic (Anisian): Karawanken Mountains, Austria.
Plate 46: Examples for Folk’s Classification
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Modified Folk Classification
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Fig. 8.7. Modified Folk classification after Strohmenger and Wirsing (1991).
(Elf-Aquitaine 1975; Flügel 1982). In reality, Folk’s classification has often been applied in a rather loose manner without following the grain ranking concept (resulting in rather monstrous names like ‘peloncobiosparrudite‘!) or by creating names, not provided for within the original classification, by including several grain types in order of increasing importance within the rock name (Friedman and Sanders 1978). ‘Intrabiooncopelmicrite’, for example would describe a micritic carbonate rock whose grains are predominantly peloids, followed by oncoids, bioclasts and fewer intraclasts. In order to restore the reproducibility of Folk’s classification and to strengthen the value of rock names for paleoenvironmental interpretations, an expanded and modified textural classification has been proposed by Strohmenger and Wirsing (1991); see Fig. 8.7. Major modifications are: • the introduction of new grain categories: (1) oncoids (indicating special environmental conditions, e.g. differences in sedimentation rates, salinity, or substrate types; see Sect. 4.2.4.1) and (2) extraclasts (indicating
special depositional conditions, e.g. subaerial exposure) as two new grain groups, • adjustment of the percentage of individual grains from 25% to 20% for intraclasts, ooids and oncoids and 40% for fossils and peloids, • Incorporation of all the grains into the rock name as long as they comprise >20% (for intraclasts, ooids or oncoids) or >40% for fossils and peloids). Owing to the introduction of oncoids as a fifth grain type – they are in third position in the ranking list (intraclasts, ooids, oncoids, fossils and peloids) – the volumetric grain proportions have been changed, except for the fossil/peloid group. For Strohmenger and Wirsing the ranking list and the grain percentage must be respected, even if, as a consequence, the most frequent grain type is not indicated by the first prefix of the rock name. A strict use of the modified classification scheme results in the differentiation of a maximum of 31 possible rock names. The practical steps involved in applying the classification are shown in Fig. 8.7.
Specific Classifications
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8.3.2.2 Discussion
tors but also by the diagenetic history, he added five categories predominantly describing diagenetic rock types (Fig. 8.5): Three of these rock types are non-obliterative with regard to the primary depositional fabric (cementstone, condensed grainstone and fitted grainstone; Pl. 47/3). Two categories, differentiated by crystal size (using Folk’s 1962 scale), are obliterative because they have lost their depositional fabric as a result of complete obliterative recrystallization or replacement. Sparstones are composed of inequant sparry calcite crystals forming a blocky mosaic, with a crystal size larger than 10 µm. The term microsparstone applies to limestones consisting of calcite mosaics with crystals between 4 µm and 10 µm (Pl. 47/2).
Folk’s classification has many advantages but also some pitfalls in that not all the information included in microfacies features is revealed: • The combined rock names (prefix for grains, suffix for matrix) are simple to learn and the terms can be easily used in a descriptive way. • The rock names can be freely combined with a number of modifying and qualifying adjectives. • The recognition of well-defined rock types facilitates facies-oriented studies because the basic binary principle is oriented to major differences in the environmental controls of deposition. However, paleoenvironmental interpretations which cling too strongly to the idea of low versus high energy controls on the sedimentation, may go wrong. The possibility, that lowenergy conditions are only simulated by lime mud trickling down into empty pores of high-energy deposits or that the fine-grained ‘micritic matrix’ does not correspond to a low-energy lime mud but instead represents a cryptocrystalline cement, should always be taken into account. • A major shortcoming of the original classification is the loss of information about grain diversity because less frequent grain types are not indicated in the rock name. This difficulty can be overcome by using the modified classification of Strohmenger and Wirsing. • Limestones containing both groundmass types represent problems in all limestone classifications. In Dunham’s classification these types are incorporated within the packstone group. Folk (1962), stressing the wide textural spectrum of carbonates, proposed a transitional class (‘poorly washed limestones’), defined as having 1/3 to 2/3 spar and 2/3 to 1/3 micrite (Fig. 8.6). Both Folk’s and Dunham’s classifications can be sharpened by using additional criteria reflecting packing intensity (e.g. densely packed, loosely packed), the existence of large grains (e.g. rudite wackestone), the degree of winnowing (poorly or strongly winnowed), the most frequent grain type, and the amount of terrigenous particles.
8.4 Specific Classifications
8.4.2 Some Nonmarine Carbonates Need very Specific Names Many nonmarine carbonates are characterized by particular components, fabrics, and microfacies types, which are difficult to describe using the standard textural classifications developed for marine carbonate rocks. Some of these difficulties are caused by the strong impact of abiogenic factors on the formation of nonmarine carbonates such as speleothems, travertine, tufa and caliche. Particular rock names have been proposed for these limestones (Gebhardt 1988; Pedley 1990; Riding 1991; Wright 1992; Koban and Schweigert 1993). Lacustrine limestones are better suited for applying standard classification schemes, especially that of Dunham, because carbonates formed within lakes are more strongly controlled by biotic factors. Folk’s classification is less appropriate because of the strong emphasis on hydrodynamic aspects, which may have had less influence on the textures of lacustrine carbonates. Nonmarine carbonates formed at the margins of aquatic environments (e.g. lakes) may be strongly affected by pedogenic and meteoric processes overprinting primary textures. The Miocene carbonates of the Ries crater lake in Southern Germany are excellent objects for testing the efficiency of a classification, by combining descriptive criteria (following the Dunham scheme) and indications of overprinting (Arp 1995). Examples are shown in Pl. 48.
8.4.1 Diagenetic Changes in Depositional Textures
8.5 Classification of Wright (1992) revised both the Dunham (1962) clas- Mixed Siliciclastic-Carbonate Rocks
sification and its modification by Embry and Klovan (1971). Because the textures of most limestones are controlled not only by depositional and biological fac-
Sediments composed of mixtures of carbonate and siliciclastic material are common in near-coast environ-
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ments, inner shelf settings as well as at high latitudes. For a long time impure limestones with admixtures of siliciclastics have been neglected, in the context of both sediment classification as well as facies interpretation. A descriptive classification system (Fig. 8.8) was proposed by Mount (1985) using four components: (1) Siliciclastic sand (sand-sized quartz, feldspar etc.), (2) non-carbonate mud (mixtures of silt and clay), (3) carbonate grains or allochems (peloids, ooids, bioclasts and intraclasts >20 µm in size) and (4) carbonate mud (or micrite, defined as <20 µm in size). The name of the sediment type reflects both the dominant grain type and the most abundant antithetic component.
Classifying Mixed Siliciclastic-Carbonate Rocks
Fig. 8.8. Classification of mixed siliciclastic-carbonate rocks. Modified from Mount (1985).
Plate 47 Classification of Carbonates That are Marked by Diagenesis The textures of most limestones result from an interplay of depositional, biological, and diagenetic factors. Carbonates classifiable as calcimudstones, wackestones, packstones, or grainstones may be products of diagenetic processes that obliterate the depositional textures or create new non-obliterative textures (Fig. 8.5; Wright 1992). Obliterative diagenetic textures can be caused by recrystallization and/or replacement and result in a complete loss of the original fabrics and the formation of sparstone or microsparstone (–> 2). Sparstones consist of sparry calcite crystals (>10 µm in diameter) forming inequant blocky mosaics. Microsparstones are composed of calcite crystals within the range of 4–10 µm. Non-obliterative textures include diagenetic packstones (–> 1, 4) and condensed or fitted grainstones (–> 3) caused by modification of originally grain-supported textures by pressure solution during burial. However, pressure solution can also completely transform textures into new ones (–> 5, 6) and create new fabrics (see Pl. 37). Displacive and replacive growth of micrite cements in pedogenic carbonates can result in a complete textural inversion. Diagenetic and biological factors are responsible for the formation of cementstones. The term refers to reef limestones consisting largely of replaced but recognizable marine cements occurring in association with predominantly encrusting organisms (see Pl. 145). Conspicuous cementstones forming reefs on the paleoshelf occur in the Neoproterozoic Cap Carbonates, which worldwide cover the glacigene sediments of Early and Late Proterozoic age of the ‘snowball earth’ (James 2001). 1 Diagenetic echinoderm packstone. Pressure solution has transformed a depositional wackestone into a diagenetic packstone fabric (Pl. 144). Bioclasts are separated by microstylolites. Note the almost complete absence of matrix. Mississippian: Muleshoe Mound, Sacramento Mountains, New Mexico, U.S.A. 2 Microsparstone formed by meteoric recrystallization of a laminated micrite. Note the equant uniform crystals. Mississippian: Big Hill, Cairn, Canada. 3 Condensed grainstone caused by compaction. Grains are flattened and compacted micrite clasts. Clasts are surrounded by thin rims of burial calcite cement. Both rims and grains were dolomitized. The mechanical compaction also affected the cements (arrows). Early Cretaceous (Purbeck facies, Berriasian): Subsurface, Kinsau well, Bavaria, Germany. 4 Diagenetic stromatoporoid packstone. The original close packing of grains has been enhanced by pressure solution. Note stylolite seams (black arrow), grain penetration (GP) and irregular grain boundaries (white arrow). Most fossils are stromatoporoids (Stachyodes, S; Stromatoporella, ST) and thamnoporid tabulate corals (T). C: Rugose corals; B: Brachiopod shells. The matrix is a black argillaceous mudstone. Back-reef environment. Middle Devonian: Brilon, Sauerland, Germany. 5 Mudstone (M), grading into argillaceous lithoclast-bearing mudstone (LM), caused by pressure solution. Lithoclasts are relicts of former ‘tuberoids’. Stylolaminated fabric (see Sect. 7.5.2). Late Jurassic: Subsurface, Saulgau well, southern Germany. 6 Marly diagenetic lithoclast wackestone/floatstone, caused by continued pressure solution of a former mudstone (–> 5). The former micrite matrix was converted into a marly matrix with up to 60 % insoluble residue. Note concentration of stylolite swarms (arrows) near pressure-solution resistant components floating within the marly matrix. Stylolaminated to stylocumulate fabric. Same locality as –> 5. –> 5, 6: Huber 1987
Plate 47: Diagenetic Rock Types
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8.6 A Name for Your Samples: Some Practical Advice Which of the two major groups does your sample fit into? Autochthonous or allochthonous limestone? Autochthonous limestones: Use the classification proposed by Embry and Klovan (Fig. 8.1A and Fig. 8.5; Pl. 41 and Pl. 42), if you are sure that you can imagine
Practical Advices
the role of the sessile benthic organisms in the formation of the reef limestone or of the microbialites. Otherwise use the sackname boundstone. Avoid the name biolithite. • Integrate into the name the dominant or ecologically important fossil group, e.g. coral framestone, or more precisely, branched coral framestone, or, if you have good taxonomic data, Retiophyllia coral framestone (see Fig. 8.2).
Plate 48 Modified Textural Dunham Classification for Non-Marine and Palustrine Carbonates Many non-marine carbonates exhibit specific grains, fabrics and microfacies that are difficult to characterize by standard classifications developed for marine limestones. The Dunham classification, reflecting hydrodynamic and biological controls of textures, can be successfully applied to some lacustrine carbonates, e.g. to those deposited in profundal and infralittoral environments. However, carbonates formed at lake margins as well as terrestrial carbonates (caliche, tufa, travertines, speleothems) are strongly prone to pedogenesis, meteoric diagenesis, and inorganic and microbial carbonate precipitation. Changes in dissolution and reprecipitation by pCO2 play an important part. In studying lacustrine carbonates, Arp (1995) developed a modification of the Dunham classification that indicates pedogenic and early meteoric alterations by adding an interpretative attribute to the descriptive Dunham term. Attributes are transformed, neoformed or derived. The classification facilitates the description of ‘palustrine carbonates’ (shallow fresh-water and marsh deposits showing evidence of pedogenic modification related to subaqueous deposition and subaerial exposure). The term is taken from the French word for swampy or marshy. Diagnostic criteria of palustrine carbonates are an association of shallow-water, bioturbated mudstones and wackestones with a low-salinity biota of ostracods, gastropods and charophytes, intraclastic grainstones and packstones, and brecciated limestones displaying mottling, root tubules and vertically elongate fenestrae (pseudo-mikrokarst: Freytet and Plaziat 1982). Palustrine carbonates, originally described from the Late Cretaceous and Early Tertiary (Freytet 1984) have been reported from a variety of stratigraphic horizons (Carboniferous to Tertiary). The Florida Everglades are taken as a modern model of the development of palustrine conditions (Platt and Wright 1992). Examples on this plate are from the Upper Miocene lacustrine carbonates of the Nördlingen Ries impact crater (southern Germany). 1 Pedogenically transformed skeletal stromatolite (SS). A brownish pedogenic calcite crust (CC) fills up depressions. The vesicular fabric (VF) is caused by subaerial exposure within the supralittoral. Erect cyanobacteria filaments (F). Ehingen. 2 Pedogenically neoformed ‘ooid’ pack/floatstone. The pedogenic ooids (OO) were formed in situ around ostracods (O) and detrital grains. Cladophorites-clast (C) derived from adjacent algal bioherms. The formation of soils in swamps caused pedogenic circumcrustations around former lacustrine grains. Hainsfarth. 3 Pedogenically neoformed nodule wacke/packstone (top and bottom) with horizontal ‘craze planes’ (CP) distinctive for palustrine facies (cracks caused by desiccation, center). The inverse gradation was formed in-situ by alternation of dissolution and desiccation. V: Multiple, fibrous and microcrystalline pendant cements and internal sediment in voids. Origin: Alternating swamp pedogenesis and prolonged meteoric-vadose conditions. Palustrine facies. Ehingen. 4 Pedogenically derived intraclast grainstone. Brownish pedogenic particles with schlieren-like cortices (arrows). Reworking of pedogenically transformed lacustrine grains and redeposition within the wave-exposed eulittoral. Ehingen. 5 Meteorically transformed intraclast rudstone consisting of reworked algal laminites (AL). Meteoric dissolution is indicated by the reduction of particles to flake-like components (black arrow) and partial replacement of micrite by sparite (white arrows). Note fracturing of meteoric-phreatic cements (FC). CH: Encrusted chiropteran larva. Interpretation: Deposition of coarse carbonate sands within wave-exposed eulittoral followed by meteoric-vadose dissolution during emersions. Ehingen. 6 Meteorically transformed intraclast wackestone. The vesicular structure (VS) is caused by meteoric dissolution. Root cavities (RC) have been enlarged by dissolution and filled with drusy calcite. Supralittoral. Ehingen-Belzheim. 7 Meteorically neoformed cement-framestone. The framework consists of laminated, fibrous sinter cements (SC). The nucleation bases are preserved as a few remnants of microbial micrite (MM). Dendroid sinters (DS) are speleothems. Interpretation: Alternation of the primary thrombolite fabric by meteoric diagenesis and formation of a new fabric. Central part of a spring mound. Ehingen-Belzheim. –> 1–7: Arp 1995
Plate 48: Classification of Non-Marine and Palustrine Carbonates
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• Specify the composition of bindstones, e.g. fenestral bindstones, micritic bindstones, thrombolitic bindstones. Try to use the terms coverstone and cruststone (Box 8.1) to indicate the role of the organisms in binding. • For bafflestones and framestones include clues to the texture of the fine-grained matrix surrounding the reefbuilding organisms, e.g. coral bafflestone with finebioclastic wackestone matrix. Allochthonous limestones: Use the Dunham classification (Fig. 8.4; Pls. 44 and 45) first to separate your samples into large groups. Look at grain- or mud/matrix-support (Pl. 43). • Use the term calcimudstone (Fig. 8.4) in order to avoid confusion with argillaceous mudstone. Distinguish between pure calcimudstone with hardly any fossils or other grains and fossiliferous or peloidal calcimudstones. • Use the adjectives skeletal or bioclastic, peloidal, oolitic, oncolitic, and lithoclastic to describe the main constituents of your samples – e.g. oolitic grainstone (Pl. 45/1). If you can differentiate between intra- and extraclasts use these terms, otherwise use lithoclast.
Practical Advises
• Distinguish between sparse and packed wackestones considering abundance and packing of grains. Sparse wackestones are characterized by abundant matrix and relatively few grains, packed wackestones by high amounts of grains. • Try to distinguish packstones with fine-grained matrix from packstones with a mixed fine-grained and sparry matrix (Fig. 8.4). • For grainstones, include additional clues to sorting and grain size – e.g. very well-sorted equally sized oolitic grainstone. These informations are essential for the discussion of porosity evolution. • If you use the terms floatstone and rudstone (Fig. 8.5B) for larger grains, remember that the official 2 mm boundary does not reflect specific depositional conditions. The terms are useful if your sample consists of millimeter to centimeter-sized grains (e.g. calcirudites). • If you want a more specific and sophisticated separation of your samples, use the Folk classification (Pl. 46), which allows better treatment with multivariate methods in a further step. • Determine the abundance of the grain categories using point-counting or estimation methods (see Sect. 6.2). • Determine the type of groundmass (micrite, sparite).
Plate 49 Classification of Mixed Siliciclastic-Carbonate Rocks Mixed carbonate-siliciclastic deposits have been comprehensively treated by Doyle and Roberts (1988). Carbonate rocks containing between 10 and 50% of terrigenous admixtures must be distinguished in nomenclature from pure carbonates. A useful classification scheme was proposed by Mount (1985; Fig. 8.8). The plate shows several examples. A rock containing less siliciclastic than carbonate, more sand-sized siliciclastics than non-carbonate mud and more carbonate grains than micrite is termed a ‘sandy allochem limestone’ (–> 1). Prefixes and adjectives can be added to the rock names to include dolomitic sediments (e.g. dolo-allochemic sandstone), conglomeratic textures, and biogenic constituents. 1 Sandy allochem limestone. Nummulitid foraminiferal floatstone. The matrix consists of angular, fine-sand sized quartz grains (white) and glauconite grains, embedded within calcite micrite. Most nummulitids are represented by oblique and axial sections (see Kleiber 1991 for clues for determining axial sections). Near-coastal inner shelf setting. Early Tertiary (Lutetian, Eocene): Kressenberg, Bavaria, Germany. 2 Allochemic sandstone. Fossils are Nuia (N), mainly regarded as an alga. Note poorly sorted subangular and rounded quartz grains (Q) and carbonate lithoclasts (L). Cambro-Ordovician: Antarctica. 3 Muddy allochem limestone. Packstone. Most carbonate grains are worn and strongly altered foraminiferal tests (fusulinids, F). Note distinct burrowing (B). Burrows are compacted. Middle Carboniferous (Kasimovian): Carnic Alps, Italy. 4 Coarse-grained sandy allochem limestone. Carbonate grains are reworked barnacles (B). Note poor sorting of quartz grains. The limestone was deposited in a high-energy near-coastal temperate setting. Barnacles are highly resistant to abrasion. Late Tertiary (Early Miocene): Burgschleinitz near Eggenburg, Austria. 5 Sandy allochem micrite. Sandy mollusk shell packstone. Note gastropods (GA) and coated bivalve shells (BS). All the primary aragonitic mollusk shells are dissolved. Near-coast mixed siliciclastic-carbonate shelf. Middle Jurassic (Bajocian): Analamanga near Sakaraha, southwestern Madagascar. 6 Micritic sandstone. Note poor sorting of the detrital quartz grains. Carbonate grains are lithoclasts with ferruginous coatings (L), echinoderms (E), and coral fragments (C). The high number of angular quartz grains in combination with well-rounded quartz pebbles points to the relatively close proximity of a shallow marginal-marine area to a hinterland. Bioclasts were probably reworked in off-coast positions and redeposited within the siliciclastic sand. Late Jurassic (Oxfordian): Jubera, southwest of Pamplona, Spain. –> 5: Sample provided by M. Geiger, Bremen; –> 6: Benke 1981
Plate 49: Mixed Siliciclastic-Carbonate Rocks
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Be aware that micrite and sparite have a priori no genetic significance. The name sparite was designed for ‘sparry calcite cement’, but the common use also includes sparry calcite mosaics representing recrystallized fine-grained matrix. • While constructing the name of your sample, indicate clearly whether you are strictly following the increasing ranking order proposed by Folk or whether you prefer the reversed ranking. Imagine for a minute that about 20% of the particles of your micritic limestone are intraclasts, 75% bioclasts, and 15% peloids – would you give your sample the name intrapelbiomicrite or biointrapelmicrite, as many people do? • If the grains exceed sand-size ranges, use the syllable rud-, e.g. intrasparrudite (a limestone consisting of millimeter-sized or larger intraclasts occurring within a sparry groundmass, Pl. 46/5) • If you prefer a strict guideline leading you to a welldefined and reproducible name, use the expanded Folk classification proposed by Strohmenger and Wirsing (Fig. 8.7). • Utilize the open-minded character of the original Folk classification, which allows the rock name to be combined with additional descriptive characters – e.g. black, poorly laminated pelmicrite (a common lagoonal limestone), or red nodular pelagic biomicrite (a common open-marine deeper-water limestone). Because of the value of fossils as paleoenvironmental proxies, a specification of the fossil types is desirable – e.g. dasyclad green algal biosparite or pectinid mollusk biomicrudite. • Consider diagenetic alterations in depositional fabrics (Pl. 47); Fig. 8.5). The terms proposed by Wright (1992) are useful, particularly those which characterize alterations that have not completely obliterated the original texture. Try to differentiate between transformed and neoformed fabrics in pedogenic or meteoric modified carbonates (Pl. 48). • Mixed siliciclastic-carbonate rocks require specific names, e.g. those proposed by Mount (Pl. 49, or by Folk (1959), who uses adjectives (e.g. clayey, silty, sandy). • And read the quotation at the beginning of this chapter again! Basics: Classification of carbonate rocks Arp, G. (1995): Lacustrine bioherms, spring mounds and marginal carbonates of the Ries-impact crater (Miocene, Southern Germany. – Facies, 33, 35-90 Baisert, D. (1967): Analyse der Nomenklaturen sedimentärer Kalkgesteine. – Zeitschrift für angewandte Geologie, 13, 642-650 Chilingar, G.V. (1960): Notes on classification of carbonate
Basics: Limestone Classification
rocks on the basis of chemical composition. – J. Sed. Petrol., 30, 157-158 Cuffey, R.J. (1985): Expanded reef-rock textural classification and the geological history of bryozoan reefs. – Geology, 13, 307-310 Davies, G.R., Nassichuk, W.W. (1990): Submarine cements and fabrics in Carboniferous to Lower Permian reefal, shelf-margin and slope carbonates, northwestern Ellesmere Island, Canadian Archipelago. – Geological Survey of Canada, Bulletin, 399, 1-77 Dunham, R.J. (1962): Classification of carbonate rocks according to depositional texture. – In: Ham, W.E. (ed.): Classification of carbonate rocks. A symposium. – Amer. Ass. Petrol. Geol. Mem., 1, 108-171 Embry, A.F., Klovan, J.E. (1971): A late Devonian reef tract on northeastern Banks Island. N.W.T. – Bull. Canadian Petrol. Geol., 19, 730-781 Fagerstrom, J.A. (1987): The evolution of reef communities. – 600 pp., New York (Wiley) Folk, R.L. (1959): Practical classification of limestones. – Amer. Ass. Petrol. Geol. Bull., 43, 1-38 Folk, R.L. (1962): Spectral subdivision of limestone types. – – In: Ham, W.E. (ed.): Classification of carbonate rocks. A symposium. – Amer. Ass. Petrol. Geol. Mem., 1, 62-84 Ham, W.E. (ed., 1962): Classification of limestones. A symposium. – Ass. Petrol. Geol. Mem., 1, 279 pp. Leighton, M.W., Pendexter, C. (1962): Carbonate rock types. – In: Ham, W.E. (ed.): Classification of carbonate rocks. A symposium. – Amer. Ass. Petrol. Geol. Mem., 1, 33-61 Mazzullo, S.J., Chilingarian, G.V., Bissell, H.J. (1992): Carbonate rock classification. – In: Chilingarian, G.V., Mazzullo, S.J., Rieke, H.H. (eds.): Carbonate reservoir characterization: a geologic-engineering analysis, part I. – 59-108, Amsterdam (Elsevier) Mount, J. (1985): Mixed siliciclastic and carbonate sediments: a proposed first-order textural and compositional classification. – Sedimentology, 32, 435-442 Oates, J. A.H. (1998): Lime and limestone. Chemistry and technology, production and uses. – 474 pp., Weinheim (Wiley-VCH) Plumley, W.J., Risley, G.A., Graves, R.W., Kaley, M.E. (1962): Energy index for limestone interpretation and classification. – In: Ham, W.E. (ed.): Classification of carbonate rocks. A symposium. – Amer. Ass. Petrol. Geol. Mem., 1, 87-107 Strohmenger, C., Wirsing, G. (1991): A proposed extension of Folk’s textural classification of carbonate rocks. – Carbonates and Evaporites, 6, 23-28 Tsien, H. (1994): Contribution of reef building organisms in reef carbonate construction. – Courier Forschungsinstitut Senckenberg, 172, 95-102 Wood, R. (1993): Nutrients, predation and the history of reefbuilding. – Palaios, 8, 526-543 Wood, R. (1995): The changing biology of reef-building. – Palaios, 10, 517-529 Wright, V.P. (1992): A revised classification of limestones. – Sed. Geol., 76, 177-186 Further reading: K145
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9 Limestones are Biological Sediments
Most limestones are directly or indirectly influenced and controlled by biological processes. The present chapter deals with the formation of carbonates by microbes and benthic encrusting organisms, and with the destructive role of micro- and macroborers. Knowledge of constructive and degrading processes is essential in evaluating carbonate budgets.
9.1 Microbial Carbonates and Stromatolites Microbes encompass bacteria, fungi, small algae and protozoans. Bacteria comprise two major groups Archaebacteria and Eubacteria, now respectively called Archaea and Bacteria (including Cyanobacteria). The sedimentologically important Cyanobacteria (Sect. 10.2.1.1) are aerobic phototrophs, living in shallow-water and using sunlight as energy. Cyanobacterial calcification is associated with the photosynthetic uptake of CO2 and/or HCO3- that raises alkalinity (Pentecost and Riding 1986; Merz-Preiss 1999) and leads to calcification of the mucilaginous sheaths. Present day, intense cyanobacteria calcification appears to be essentially a freshwater phenomenon and is rare in modern subtidal environments, by contrast to ancient cyanobacteria which occupied tidal and subtidal environments. Many other bacteria are anaerobic heterotrophs and take their energy from the decomposition of organic material to inorganic components by redox processes. These bacteria can occupy lighted and shallow as well as dark and deep-water settings, and are responsible for ammonification, denitrification, sulfate reduction, anaerobic sulphide reduction and methanogenesis processes. These processes can lead to HCO3– concentration and increasing alkalinity favoring fine-grained CaCO3 precipitation (Knorre and Krumbein 2000) in the form of micrite. Microbial CaCO3 precipitation is triggered by various processes associated with (a) bacteria and small
algae, (b) extracellular polymeric substances (EPS; accumulating outside the cells to form a protective and adhesive matrix that attaches microbes to the substrate), (c) biofilms (submillimetric veneers of bacterial communities in an EPS matrix), (d) microbial mats (millimeter-sized complex layers composed of filamentous cyanobacterian algae and diatoms, and able to trap sediment), and (e) organomineralization (precipitation of CaCO3 in association with nonliving macromolecules independent of organic activity). See Riding (1991, 2000), Van Gemerden (1993), and Riding and Awramik (2000) for further explanations. The eminent role of bacteria and other microbes in the formation of carbonate rocks was summarized in Sect. 4.1.2 and is discussed in the following paragraphs which also cover controls and the terminology of microbial carbonates, and examine the importance of stromatolites.
9.1.1 Bacterial Contribution to Carbonate Precipitation Microbial precipitation of calcium carbonates played a vital role in the development of Proterozoic and Phanerozoic carbonate platforms and reefs. The importance of bacteria and cyanobacteria in the formation of finegrained carbonates in natural aquatic environments has long been a matter of discussion, starting with the research of Drew (1911, 1913) and Black (1933) on the action of denitrifying bacteria in tropical and in temperate seas, followed by observations on calcium carbonate precipitation in seawater and freshwater environments (Pentecost 1985). Reviews discussing this topic have been published by Cohen et al. (1984) and Jones (1985). Many researchers have demonstrated that life processes of marine bacteria and the decomposition of organic matter by bacteria cause physicochemical changes in the microenvironment that can result in calcium carbonate precipitation. This has been proved in laboratory experiments and observed in various modern carbonate-producing environments (soils, freshwa-
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_9, © Springer-Verlag Berlin Heidelberg 2010
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Recognizing Microbial Carbonates
ter and marine realms, particularly in lagoons; Castanier et al. 1989; Chafetz et al. 1991; Folk 1993; Arenas et al. 1993; summary by Buczynski and Chafetz 1993). Experiments indicate that bacteria trigger the precipitation of aragonite and calcite, exhibiting distinct morphologies that seem to be limited to bacterial contribution. The size of individual crystals, spheres and rods ranges between 0.1 and 0.4 µm; that of crystal aggregates between 5 and 100 µm. Comparable morphologies have been observed in laboratory cultures and in modern carbonate sediments. Because most bacteria except for cyanobacteria, are indifferent to light, bacterially-controlled carbonate precipitation is not restricted to shallow environments, but also occurs in deeper subtidal settings, various cryptic habitats and in deep restricted basins. A strong microbial contribution of microbes to the formation of ‘mud mounds’ is advocated by many authors (e.g. Pratt 1982; Lees and Miller 1985; see Sect. 16.2.2).
to basin traverses (Precambrian: Hoffman 1974; Late Devonian: Playford et al. 1976; Late Jurassic: Leinfelder 1993; Tertiary: Braga et al. 1995).
Biological versus environmental controls: The main processes of microbial carbonate formation are (1) trapping (agglutination) of sedimentary particles, (2) biomineralization (calcification) of organic tissues, and (3) mineralization (superficial precipitation of minerals on organisms and/or sediment). Fine-grained carbonate is trapped and produced within microbial mats, occurring under a wide range of environmental conditions in nonmarine and marine sites. The mats are dominated by various phototrophic, chemotrophic and heterotrophic microorganisms (cyanobacteria dominating in the top layer; colorless sulfur bacteria, purple sulfur bacteria, and sulfate-reducing bacteria harboring the underlying layers). Other numerically less important groups are nitrifying and denitrifying bacteria and methanogenic bacteria. The activity of aerobic heterotrophic organisms leads to oxygen depletion. Fermentative organisms provide growth substrates for sulfate-reducing bacteria. The vertically laminated structures develop as a result of microbial growth and activity sediment trapping and binding in the organic matrix, and sedimentation (Van Gemerden 1993). The shape and macrofabric of microbial carbonates are strongly influenced by variations in the depositional environment. Important controlling environmental parameters are the grain size of the substrate, the penetration of light, sedimentation and erosion rates, and grazing pressure (Walter 1976). Sedimentation and microbial mat composition are sensitive to water movement and light, respectively, and change with water depth. This is reflected in the morphology, texture and microfabrics. Examples were described from various shelf
Microbial carbonates are recorded by specific textures occurring within a wide range of scales. Common features are dense autochthonous micrites, micro- to megascaled laminated fabrics (e.g. stromatolites), millimeter- to centimeter-sized micritic crusts, non-laminated micritic and peloidal structures, and limestones exhibiting tiny tubelike microfossils referred to as remains of cyanobacteria.
9.1.2 How to Recognize Microbial Carbonates? Increasingly more authors support the idea of a strong microbial impact on the formation of Paleozoic and Mesozoic limestones, using external appearance, internal fabrics and geochemical signatures as evidence (see Facies, vol. 29, 1993). Box 9.1 lists some criteria which are commonly used as evidence for a bacterial contribution to the formation of limestones.
9.1.3 Describing and Classifying Benthic Microbial Carbonates
9.1.3.1 Terminology and Descriptive Criteria Microbial carbonates are carbonate deposits produced or localized by benthic microbial communities (Riding 1990) living in marine, marginal-marine, freshwater and terrestrial environments. The complex associations of bacteria, cyanobacteria (cyanophytes, bluegreens) and algae embrace photosynthetic prokaryotes, eukaryotic microalgae and chemoautotrophic as well as chemoheterotrophic microbes. In addition, encrusting invertebrates (e.g. foraminifers) may be of some importance (Sect. 9.2). Communities creating microbial carbonates are termed microbial mats (Gerdes and Krumbein 1987) or algal mats, reflecting the densely interlayered and intertwined orientations of the filamentous and coccoid cells involved and the resulting biolaminated sedimentary structures. Biolaminites characterized by organic-rich laminae and microbial mats are an essential criterion of microbially induced sedimentary structures (Noffke et al.
Benthic Microbial Carbonates
1996). The term microbialite (Burne and Moore 1987; changed to microbolite by Riding 1991) characterizes ‘organosedimentary deposits that have accreted as a result of benthic microbial community trapping and binding detrital sediment and/or forming the locus of mineral precipitation’. Microbialite is often used as an overall term, but also as a term restricted to non-laminated microbial carbonates only in contrast to stromatolite, which denotes laminated benthic microbial deposits (Riding 1999); see Box 9.1 using the classification developed by Logan et al. (1964; Fig. 9.3). The fabric describes lamination or non-lamination and the spatial composition (e.g. clotted). Microstructure refers to the type and fabric of the constituents (commonly micrite or microspar, various grains, and fenestral spar-filled voids, sometimes encrusting fossils and borings).
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9.1.3.2 Classification of Benthic Microbial Carbonates Box 9.2 summarizes the current classification of microbial carbonates (examples in Pl. 8 and Pl. 50). Comments on non-laminated microbialites Thrombolites (Pl. 8/1, 6, Pl. 50/5, Pl. 131/5). The term was proposed as a field term for ‘cryptalgal structures related to stromatolites but lacking lamination and characterized by a macroscopic clotted fabric’ (Aitken 1967). The name refers to micrite clots. The clots differ in color, form and/or texture from the intervening area and create a blotchy fabric. Kennard and James (1986) called the components of thrombolites mesoclots, which are typically dark colored and have a microcrystalline texture. They display a variety
Box 9.1. Which thin section and SEPM criteria provide evidence of a microbial contribution to the formation of limestones? Very fine-grained dense micritic matrix: Dead calcified bacterial cells can become calcified (Krumbein 1979) and form discrete micron-sized bodies. Small varieties found in micritic carbonates have been attributed to nannobacteria (size 0.05 to 0.2 µm) because the sphere- and bean-shaped objects seen in SEM exhibit similarities with regard to morphology and cluster-like distribution patterns (Folk 1983). Caution is needed because similar objects can be also formed abiotically (Kirkland et al. 1999, Abbott 1999). Laminated and undulated micritic structures: Comparing of laminated textures of fine-grained limestones with microtextures produced by bacteria in laboratory experiments shows that micritic finely laminated and undulated microstructures (Pl. 6/5, Pl. 50/4) can be formed by microbes that contribute to the precipitation of seafloor automicrite (see Sect. 4.1.1). Constructive micrite envelopes: Some of these envelopes (see Sect. 4.2.3) may represent biofilm calcification (Perry 1999). Bacteria adhere to surfaces for stability and create calcifying biofilm communities augmented by other microbes. Clotted fabrics: Clotted fabrics (Pl. 10/1) are widespread in stromatolites and thrombolites. The fabric appears to represent EPS calcification. It has been referred to as spongiostrome (Gürich 1906) and structure grumeleuse (Cayeux 1935), and can grade into dense micrite. Diffusely clotted micrite often forms clusters of rounded aggregates within microsparite and is associated with filamentous microfossils (Guo and Riding 1992). Calcimicrobes: Calcification of microbial, external polysaccharide-protected, sheets produces calcified fossils (calcimicrobes, e.g. Girvanella, Cayeuxia, Pl. 8/3, Pl. 53) which are represented by tiny tubes with micritic walls. Most of these fossils are cyanobacteria, and occur in association with finely peloidal and clotted fabrics. Peloids: Silt to sand-sized micritic grains and aggregates (20–60 µm) are common constituents of modern tropical carbonates. In-situ precipitation has been observed in Holocene reefs (Lighty 1985). These Mg-calcite components have been regarded as cement (Macintyre 1984, 1985), but also as calcified bacterial aggregates rimmed by euhedral calcite crystals (Chafetz 1986; Pl. 8/5; see Sect. 4.2.2). Potential microbial peloids are widespread in ancient reefs (Pl. 8/6), in stromatolites and thrombolites. Microbial ooids: The recognition of bacterially constructed carbonate crystals and grains (Fig. 4/24) in recent environments and the similarity of these grains with particles occurring in ancient carbonates (Gerdes et al. 1994; Reid 1987; see Sect. 4.2.5) offer a clue to the microbial character of some ooids enclosed in laminated textures. Microspar and spar: Fibrous, equant and dendritic spar precipitates (Hofmann and Jackson 1987) occur as external crusts on organic tissue and mineral surfaces of microbial carbonates, particularly those originating in fresh water (tufas and travertines, karst) and schizohaline settings. These spars also form rosettes and spherulites around cyanobacterial precipitates and on bacterial cells (Guo and Riding 1992; Defarge et al. 1996). Pores and allochthonous grains: Discrete voids ranging from tiny interstices to large growth cavities (Pratt 1995) are often associated with microbial carbonates. They include irregular fenestrae and tidal flat birdseyes. Trapped grains are important constituents of these carbonates (Pl. 20/4, Pl. 50/5, Pl. 124/1).
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Box 9.2. Classification of organosedimentary deposits controlled by benthic microbial communities (after Riding 2000). The classification is based on macro- and microfabric criteria. The presence or absence of lamination is used in distinguishing two major groups. Note that some authors restrict the term microbialite to nonlaminated microbial carbonates, whereas others use the term as a synonym for all microbial carbonates. Microbialites (Burne and Moore 1987): Non-laminated benthic microbial deposits: Thrombolites (Aitken 1967): Clotted mesofabric. Dendrolites (Riding 1988): Dendritic mesofabric, that may be distinct, crude or diffuse. Leiolites (Braga et al. 1995): Structureless, aphanitic mesofabric. Cryptic microbialites (Riding 1991): Lack distinct mesofabrics, but typically possess micritic, clotted or peloidal microfabrics. Stromatolites (Kalkowsky 1908): Laminated benthic microbial deposits: Agglutinated stromatolites (Riding 1991): Produced by trapping/binding of particulate sediment. Well-laminated and fine-grained, or crudely laminated and coarsegrained. Skeletal stromatolite (Riding 1977): Produced by in-place organisms that are commonly preserved as calcified fossils. Freshwater tufa stromatolites (Riding 1991): Produced by encrustations of external sheets of cyanobacteria and green algae. Terrestrial stromatolites (Riding 2000): The term refers to fabrics seen in laminar calcretes.
of shapes (subrounded, amoeboid, grape-like, arborescent, digitate, pendant) and occur isolated, interconnected, or amalgamated. The clots can make up in excess of about 40% of the volume of a thrombolite rock. Mesoclots may be lobate, cellular, microspherulitic, grumous, peloidal, vermiform, or mottled. Calcified microfossils such as Renalcis, Girvanella and Nuia are common in Paleozoic thrombolites. The clots are interpreted as a complex of irregular agglutination, insitu calcification of coccoid or coccoid-dominated microbial communities, skeletal encrustation, and erosional processes, but may result also from calcified filaments (Burne and Moore 1987). Walter and Heys (1985) consider some thrombolites as being disturbed stromatolites in which the original laminated structure was disrupted and modified by bioturbation or subsequent diagenesis. In thin sections thrombolites show a complicated microstructure. The individual micrite masses may consist of (1) dense micrite, (2) clotted micrite, (3) peloids of irregular shape and size, sometimes retaining porostromate structures, (4) calcite spherulites consisting of radially-oriented, non-ferroan calcite, typically 12– 20 µm in diameter with a cloudy micrite core, and
Organosedimentary Deposits
(5) various microfossils encrusting micrite and peloids. Many thrombolites contain micro-encrusters. Riding (2000) distinguished calcified microbial thrombolites typically displaying well-defined clots, coarse agglutinated thrombolites that incorporate sandand even gravel-sized sediment and commonly exhibit complex internal variations (e.g. Shark Bay), some of which are associated with sponge-tissue degradation (Reitner 1993; Reitner et al. 1995); arborescent thrombolites associated with decimetric dendritic fabrics (Schmitt and Monninger 1977, Armella 1994); tufa thrombolites occurring in freshwater lakes and streams (Moore and Burne 1994). Thrombolites are essentially subtidal, common in deeper-water settings, and form columns, domes, layers and thick crusts, typically dome-shaped meter-scale doughnut-like masses (kalyptrae: Luchinina 1975; Rowland and Gangloff 1988). They reflect an irregular and uneven supply of sediment on surfaces that were patchily colonized by microbes. Thrombolites range from the Neoproterozoic to the modern and were important as constituents of smallscaled reefs throughout much of the Cambrian and Early Ordovician (Shapiro and Rowland 2002; Webby 2002). Reports of post-Ordovician thrombolites are rare (Silurian: Kahle 2001; Late Jurassic: Leinfelder and Schmid 2000; Cretaceous: Neuweiler 1993; Miocene: Riding et al. 1991). Dendrolites (Pl. 8/4). These microbialites form large domes and columns that may be roughly layered, and exhibiting bush-like dendritic fabrics, either vertically erect or pendant. The fabric differs in color from adjacent areas and is distinct, crude or diffuse. Dendrolites formed by porostromate microfossils (e.g. Epiphyton; Pl. 82/4) are common features of Early Paleozoic reefs, growing at the surface, within dark cavities (endostromatolites: Monty 1982) or in fissures (Pl. 8/4). Rapid calcification and early cementation between the bushy fossils resulted in the formation of an organic framework that provided hard substrates for attached metazoans (e.g. archaeocyaths, stromatoporoids). Dendrolites were most conspicuous during the Early Cambrian and Early Ordovician, Late Devonian and Early Carboniferous. Bushy fabrics of dendrolites contribute significantly to the formation of nonmarine carbonates, e.g. travertines (Pl. 2/1). Leiolites form large domes in association with stromatolites and dendrolites. These microbial deposits with structureless micritic fabrics often appear ‘homogeneous’ or ‘massive’ and lack clear lamination, clots or dendritic fabrics. Leiolite formation is favored by a
Spongiostromata and Porostromata
steady uniform supply of well-sorted sediment on surfaces colonized by an even layer of microbes. Records of leiolites are rather rare (e.g. Late Jurassic: Dupraz and Strasser 1999, Leinfelder and Schmid 2000; Miocene: Braga et al. 1995). Cryptic microbial carbonates (Pl. 8/7). Many finegrained limestones, especially from reefal environments, may be microbial in origin but lack distinctive macrofabrics. They are characterized by inhomogeneous micritic, clotted or peloidal fabrics, showing locally some traces of filaments. Many of these fabrics are also present in stromatolites, dendrolites and thrombolites. Further classifications Microbialite features can be differentiated at various scales (Shapiro 2000) and according to the amount of compositional constituents (Schmid 1996; Fig. 9.1). Comments on Spongiostromata and Porostromata These groups were proposed by Pia (1927) for schizophycean blue-green algae of uncertain affinities. Spongiostromata refer to nodules or heads with well-
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defined growth forms but without (or very rare) with preserved organic microstructures (Pl. 50/4). Porostromata comprise forms exhibiting tubular microstructures (Pl. 50/1). Spongiostromate microstructures are very variable, including micritic, spongious, vermicular, fenestral, and peloidal textures (Gürich 1906). Comparable textures occur in recent and ancient stromatolites (Monty 1976). Porostromate microstructures are typified by loose or tangled, variously oriented, straight or sinuous tubes. Because these tubes record organisms of different systematic position Riding (1977) proposed the term ‘skeletal stromatolites’. The terms spongiostromate and porostromate are used in microfacies studies describing textural variations of biogenic carbonate crusts and oncoids (see Sect. 4.2.4.1). Spongiostromate refers to a laminated, poorly differentiated micritic and peloidal microfabric (Pl. 28/1) that may include various encrusting fossils. Porostromate microfabrics characterized by tiny calcified tubes occur in skeletal stromatolites (Pl. 50/1), but also in non-laminated microbialites (Pl. 124/1), and as constituents of biogenic nodules (Pl. 12/3, 4), in reef limestones as well as in carbonates formed in nonmarine and marine-nonmarine transitional environments (Pl. 50/6).
Fig. 9.1. Differentiation of microbialites. A: The four scales of microbialite investigation. Microbialites should be differentiated according to (1) the megastructure (referring to the large-scale features of microbial limestones, e.g. biostromal buildup), (2) macrostructure (designating the shape of the microbialite, with typical diameters of centimeters to meters, e.g. domal or columnal), (3) mesostructure (describing internal textures of macrostructural elements visible with the naked eye, e.g. laminated, clotted, or dendritic), and (4) the microstructure (referring to microscopic feature observed under the microscope or in SEM, e.g. peloids, filamentous microbes). The scale bars indicate the magnitude of the criteria described at each of the stages. Slightly modified from Shapiro (2000). B: Classification of Mesozoic microbialites (modified from Schmid 1996). This valuable classification relies on microstructures observed in thin sections. The system is apparently not applicable to the morphologically highly variable Precambrian and Early Paleozoic microbialites. See Pl. 50 for examples.
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9.1.4 Stromatolites are Laminated Microbialites In describing the siliciclastic Rogenstein deposit of the Early Triassic Buntsandstein of the German Harz Mountains, Kalkowsky (1908) introduced the term stromatolith for ‘limestone masses showing a fine, more or less planar, layered structure’ composed of a set of laminae (‘stromatoid’) and believed to be of vegetal origin (see Paul and Peryt 2000). The use and definition of the term stromatolite have undergone gradual revisions (Riding 1977; Krumbein 1983) and oscillate strongly between descriptive and genetic approaches (Monty 1977; Semikhatov et al. 1979; Awramik 1984, 1990).
Stromatolites
I recommend the definition proposed by Riding (1999): Stromatolites are laminated benthic microbial deposits. This definition emphases the two main elements included in the original definition by Kalkowsky: lamination and biogenicity. Recent stromatolites occur predominantly in freshwater (Freytet and Verrecchia 1999), marginal marine (Fig. 9.2) and shallow subtidal environments. Ancient stromatolites are known from the same settings, but from deep subtidal and basinal environments. Stromatolite fabrics can also form in geothermal systems as well (Jones et al. 2002) and result from anaerobic methane oxidation at cold seeps (Greinert et al. 2002).
Plate 50 Microbialite and Stromatolite Fabrics Biogenic laminated and non-laminated microbialite crusts forming various growth forms and consisting of micrite, grains and fenestral pores occur in different settings of shallow- and deep-marine as well as nonmarine carbonates. The fabric corresponds to that of stromatolites (skeletal stromatolite, –> 1; agglutinated stromatolite, –> 2) or non-laminated microbialites (e.g. thrombolite, –> 5). The microbialites are built by the combined effect of biota, sedimentation and diagenetic processes. Microbial contributions, rapid cementation as well as the assistance of encrusting organisms (foraminifera, serpulids) are essential in the formation of these microbialites. 1 Skeletal stromatolite crust growing on a colonial reef coral. The crust consists of slightly undulated spongiostromate micritic layers (ML) separated by layers of porostromate cyanobacteria filled with sparry calcite (SC). Late Triassic (Dachsteinriffkalk, Norian): Gosaukamm, Austria. 2 Laminated fine-grained agglutinated stromatolite produced by trapping/binding of sediment. Detail of microbial crusts consisting of thicker peloid layers (PL) and thinner micrite layers (ML). Peloids are very small and surrounded by calcite rims. They are isolated or occur in amalgamated masses. Micrite layers consist of micritic laminae with peloids and (white) bodies composed of radially arranged calcite crystals (Baccanella, see Pl. 99/8). Both, calcite-rimmed small peloids and Baccanella point to a microbial origin. These crusts are abundant in mm- to cm-sized framework voids of Triassic reefs. The peloidal texture is characteristic of ‘container organomicrites’ (see Sect. 4.1.1) formed in semi-restricted cavities. Late Triassic (Norian): Gosaukamm, Austria. 3 Microbial crusts around and between rugose corals (Peneckiella). Microbial circumcrustations around reef-building organisms contribute significantly to the stabilization and preservation of reef structures. The crusts exhibit peloidal and clotted structures. Note that the crust around adjacent corallites is interconnected, forming a framework. This sequence can be explained by the leeward position of the Peneckiella thickets on the forereef slope (Gischler 1995) that hampered a rapid early cementation. Interskeletal pores are filled with burial calcite cements. Black spots between the septa of the corals and between coral calices are ipsonite, an asphaltic probitumen derived from the metamorphism of hydrocarbons due to thermic effects. Atoll reef, formed by stromatoporoids and corals on top of a volcanic seamount. Late Devonian (Frasnian): Iberg, Harz, Germany. 4 ‘Spongiostromate’ stromatolite crust covering the wall of a cryptic reef cavity. Micritic stromatolite according to the classification by Schmid (1996). The term spongiostromate refers to the variable, undifferentiated microfabric. Note the variable thickness of the micritic laminae. These crusts which are several centimeters thick, are interpreted as microbially induced carbonate precipitation within biofilms coating the interior of semi-closed reef voids and forming autochthonous micrites (container organomicrite). The interior of the cavity is occupied by Baccanella (arrow), supposed to be of bacterial origin or as diagenetic products caused by the recrystallization of micritic High-Mg calcite and aragonite (Pratt 1997). Late Triassic (Norian): Gosaukamm, Austria. 5 Agglutinated microbialite consisting of amalgamated peloids (AP) leaving space for spar-filled cavities (C) forming a ‘laminoid fenestral fabric’ (see Sect. 5.1.5.2). The absence of a distinct lamination prohibits an assignment as stromatolite. Poorly structured thrombolite according to the classification by Schmid (1996). SMF 21. Late Triassic (Norian): Zelenica, Begunjscica Mountains, Slovenia. 6 Tufa stromatolite. Cement/algal bindstone, characterized by alternating thick layers of bladed elongate calcite cement crystals (CC) and layers consisting of radiating bundles of algal threads (arrow). The algal character of these threads is supported by distinct bifurcations. Schizohaline near-coastal environment. Tertiary (Miocene): Gulf of Suez, Egypt.
Plate 50: Microbialites and Stromatolites
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Recent Stromatolites
Fig. 9.2. Holocene stromatolites formed in an intertidal environment characterized by extreme salinity and limited water circulation. Stromatolites grow slowly, less than 1 mm and as little as 0.04 mm/year. Hamelin Pool, Shark Bay, Western Australia. Courtesy of R. Höfling, Erlangen.
Stromatolites are not necessarily carbonate, but may be siliceous (Walter et al. 1972), evaporitic (Friedman and Krumbein 1985; Gerdes et al. 1985; Renaut 1993; Gasiewicz and Peryt 1994; Rouchy and Monty 2000), or phosphatic (Krajewski et al. 2000). Widespread siliciclastic stromatolites occur on modern temperate tidal flats (Cameron et al. 1985; Gerdes and Krumbein 1986; Noffke 1998). Carbonate-siliciclastic stromatolites are common (Martin et al. 1993; Bertrand-Sarfati 1994). The laminated fabric of stromatolites is distinct, crude or diffuse (Pl. 8/2, Pl. 50/2), depending on the degree of continuous development of the fabric. Stromatolite formation is favored by a regular and even supply of sorted sediment over time and space onto surfaces colonized by an even layer of microbes. Subdivision of stromatolites (1) Many stromatolites result from trapping/binding of particulate sediment. These agglutinated stromatolites exhibit two subtypes: (a) Fine-grained, welllaminated stromatolites characterized by conspicuous lamination, due to episodic sedimentation and trapping of fine-grained sediment. Lamination may be smooth (subaqueous) or crinkled with fenestrae (due to exposure and desiccation on tidal flats or lake margins). (b) Coarse-grained, crudely laminated stromatolites characterized by irregular or inconspicuous lamination and sand-sized sediment. Recent examples are the columnar stromatolites of Shark Bay (Fig. 9.2) and the giant subtidal Bahamian stromatolites at Lee Stocking Island.
(2) Skeletal stromatolites, are produced by biochemical calcification of in-place organisms (cyanobacteria and algae) and characterized by the presence of tubular microfossils. This type seems to be restricted to marine and marginal-marine Paleozoic and Mesozoic carbonates. Some recent lacustrine stromatolites resemble skeletal stromatolites but mineralization is more important than biomineralization in these freshwater microbial carbonates. (3) Freshwater tufa stromatolites are formed by encrustations of sheets of cyanobacteria and green algae (Cladophora, Oocardium, Vaucheria; Arp 1995) associated with bacteria. (4) The term terrestrial stromatolites corresponds to laminar calcretes formed by microbial activity (Wright 1989; Riding 1991). Differentiating stromatolite structures Various approaches have been used in classifying and describing ancient stromatolites: (1) binary taxonomy, (2) geometric pattern (Logan et al. 1964; Fig. 9.3), and (3) morphometric analysis (Hofmann 1976). The use of spatial power spectrum image analysis is an effective means for the morphological study and comparison of stromatolites. The Logan et al. classification was developed to describe growth forms of intertidal and shallow subtidal stromatolites. These growth forms were believed to be predominantly controlled by the degree of water turbulence. Laminated mats are common in supratidal and intertidal quiet-water environments, domal growth
Microbial Mats and Stromatolites
Box 9.3. Selected references on modern marginal-marine and marine microbial mats and stromatolites. Most references deal with tidal and subtidal settings which certainly do not represent all the wide variations of the depositional environments of ancient microbialites. In modern reefs laminated microbial crusts occur as coatings on corals and other organisms and can contribute to the formation of high-energy coralalgal-stromatolite frameworks. Bahamas: Bathurst 1967; Black 1933; Cao and Xue 1985; Dill et al. 1986; Dravis 1983; Feldmann and MacKenzie 1998; Gebelein 1976; Ginsburg et al. 1977; Hardie 1977; Hardie and Ginsburg 1977; Macintyre et al. 1996, 2000; Mann and Nelson 1989; Monty 1965, 1967, 1972, 1976; Neumann et al. 1970; Park 1977; Paull et al. 1992; Pickney et al. 1995; Reid and Browne 1991; Reid et al. 1995; Riding 1994; Riding et al. 1991; Scoffin 1970; Shapiro et al.1995; Visscher et al. 1998 Florida: Frost 1974; Gebelein 1976; Ginsburg et al. 1954 Bermuda: Gebelein 1969, 1976; Sharp 1969 Mediterranean: Friedman and Krumbein 1985; Friedman et al. 1973; Gerdes and Krumbein 1984; Gerdes et al. 2000 Red Sea: Friedman 1972 Pacific: Defarge et al. 1994; Kempe and Kazmierczak 1990, 1993, 1997; Kempe et al. 1996 Shark Bay, western Australia: Bauld 1981, 1984; Chivas et al. 1990; Davies 1970; Golubic 1982, 1985, 1992; Hoffman 1976; Logan 1961; Logan and Brown 1986; Logan et al. 1974; Playford 1990; Playford and Cockbain 1976; Riding 1994; Skyring and Bauld 1990; Walter 1984 South Australia: Bauld et al. 1980; Skyring et al. 1983 Reef environments: Cabioch et al. 1999; Camoin and Montaggioni 1994; Camoin et al. 1999; Macintyre 1993; Macintyre et al. 1996; Marshall 1983; Montaggioni and Camoin 1993; Reid and Browne 1991; Reitner 1993; Reitner et al. 1995; Scholz 2000; Sorokin 1973; Steneck et al. 1998; Taylor 1977; Thiel et al. 1996; Wood 1962
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forms are common in subtidal settings. This interpretation may fit the rough shapes of large stromatolite bodies, but the shape of small stromatolites and their constructional composition depends strongly on the biota involved. Lamination, lamina shape, fabric and texture as well as gross morphology can be substantially controlled by the taxonomic composition of the mat-building community (Awramik and Semikhatov 1979; Hofmann 2000) and reflect the activity of microcommunities living under different microenvironmental conditions (Defarge 1996). Quantitative methods in the study of stromatolite fabrics facilitate the recognition of microenvironmental changes.
9.1.5 Occurrence and Significance of Microbialites and Stromatolites 9.1.5.1 Development through Time Stromatolites and other microbialites played an important role in the formation of Precambrian and Phanerozoic limestones originating in various settings including marine, transitional marginal-marine and nonmarine environments. Whereas the microbial contribution to marginal-marine carbonates was uninterrupted through time, the role of microbes varied in the formation of shallow-marine platform and reef carbonates (Fig. 9.5). The Golden Age of stromatolites was the period from 2 800 Ma to 1 000 Ma where stromatolites contributed significantly to the formation of tidal carbonates, platforms and reefs (Fig. 9.4), and during the Proterozoic reached a maximum in abundance, diversity and lateral environmental expansion (Grotzinger 1989; Grotz-
Fig. 9.3. Stromatolite classification after Logan et al. (1964). The classification is based on basic geometric forms expressed by the vertical and lateral arrangement of hemispheroids. Stromatolite growth forms as well as the shape of the lamination is described by symbols and formulas. These symbols can be used in the field and in the laboratory to describe thin sections and polished sections. LLH: Laterally Linked Hemispheroids with laminae whose domes are either Closely packed or Spaced somewhat apart (subtypes LLH-C and LLH-S). SH: Stacked Hemispheroids forming columns that are separated by sediment. The domes of the laminae have either a Constant diameter or Various widths (subtypes SH-C and SH-V). SS: Spheroidal Structures around a nucleus (corresponding to oncoids). Subtypes are SS-C (characterized by a Concentric structure; normal oncoid), SS-R (laminae Randomly overlapping), and SS-I (Inverted; laminae facing each other as concentric hemispheres), see Fig. 4/15. Mixed geometric forms can be indicated by a linear combination of symbols, e.g. LLH-SH. The relations of growth forms and microstructure is expressed by a fraction, whereby the numerator describes the macrostructure seen in the field and in hand specimens, and the denominator the microstructure seen on a smaller scale as in thin sections.
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Precambrian and Paläozoic Stromatolites
ter settings can be found in the papers included in the SEPM Special Volume 72 on Phanerozoic Reef Patterns (Kiessling et al. 2002).
Fig. 9.4. Late Cambrian (La Flecha Formation) stromatolite from the Eastern Precordillera, between San Juan and Mendoza (Argentina). Scale is 50 cm. Courtesy of W. Buggisch, Erlangen.
inger and Knoll 1999; Hoffman 2000; James et al. 2000; Semikhatov and Raaben 2000; Sumner 2000). By the end of the Precambrian, stromatolite development had declined, but stromatolites and other microbial carbonates were still important in the formation of reef structures in various parts of the Phanerozoic, as shown in Fig. 9.5 and summarized by Riding (2000). A wealth of new data on Paleozoic and Mesozoic microbial carbonates and their distribution in shallow- and deep-wa-
The decline of stromatolites near the Precambrian/ Cambrian boundary and the development of Cambrian and Ordovician tidal carbonates have been explained by competitive or noncompetitive restriction of stromatolites by grazing and burrowing organisms, e.g. gastropods (Garrett 1970; Mazzullo and Friedman 1977), but this interpretation has been strongly criticized (Pratt 1982; Browne et al. 2000; Rowland and Shapiro 2002). The microbial reef ecosystem flourished after the collapse of the archaeocyath reefs from the Middle Cambrian to the Early Ordovician, covering a period of about 40 Ma. Many Cambrian inter- and subtidal shelf carbonates and reefs exhibit dendrolitic, stromatolitic or thrombolitic textures (Aitken 1967; Ahr 1971; Soudry and Wissbrod 1993; Pl. 82/2, 4). Stromatolites occur throughout the Ordovician whereas thrombolites had a constant decline during the Early and Middle Ordovician, subsequent to their peak in the Late Cambrian (Webby 2002). The Silurian (Soja 1994) and Devonian (Browne and Demicco 1987) exhibit a marked decrease in the frequency of microbial carbonates, except during the upper Late Devonian where calcimicrobes and stromatolites contributed significantly to the formation of carbonate platforms during times of major environmental changes (Playford et al. 1976; Whalen et al. 2002). Microbial shelf and reef carbonates again be-
Fig. 9.5. Changes in the microbial control on limestones during time expressed by the percentage of microbial reefs (based on the evaluation of more than 2800 Phanerozoic reefs; see Kiessling 2002). The contribution of microbes to the formation of reef carbonates varied during the Phanerozoic. Stromatolites and calcimicrobes were abundant in Cambrian, Ordovician and Early Carboniferous carbonates, and common to abundant in periods following worldwide extinction events (Famennian; Early Triassic). The importance of microbial carbonates in reefs decreased after the Jurassic.
Importance of Stromatolites
came important in the Mississippian, as exemplified by the deep-water ‘Waulsortian’ mud mounds. Stromatolites gave way to thrombolites during the Viséan. Microbial contribution to Pennsylvanian and Early Permian reefs is indicated by the common association of micrite and carbonate cement crusts with encrusting organisms regarded to be microbial (e.g. Archaeolithoporella; Pl. 42/1). Similar associations occur in Middle and Late Permian shelf-margin reefs, e.g. the Capitan reef of Texas (Pl. 145/1). Shallow-water and deeperwater stromatolite reefs are widely distributed in the cratonic Zechstein basins of central Europe and the Kungurian basins of the Urals. Comparable to the abundance of microbial carbonates subsequent to the Frasnian/Famennian extinction event, microbial carbonates and small microbial reefs dominated after the Permian– Triassic extinction event during the Early Triassic. Microbes contributed to the construction of Middle and Late Triassic reef frameworks (Harris 1993; Brachert and Dullo 1994; Pl. 28/1, Pl. 116/2); some Late Triassic reefs were almost exclusively formed by stromatolites and microbial crusts. The importance of photic and aphotic microbes as reefbuilders increased again during the Late Jurassic, where low-energy microbe-dominated reefs and platform carbonates were widespread (Keupp and Arp 1990; Schmid 1996; Leinfelder et al. 2002), and continued being constituents of shallow-water and deeper-water environments during the Early Cretaceous (Strasser 1988; Neuweiler 1995). In the Late Cretaceous microbialites were gradually substituted by encrusting red algae. Some microbial-dominated reef carbonates are still known from the Tertiary.
9.1.5.2 Paleoenvironmental Significance of Microbial Carbonates Microbial carbonates, together with associated encrusting organisms, offer a high potential for interpreting ancient environments and their major controls by nutrients, light, water energy and water depth. Nutrification is influenced by relative sea-level changes, which result in shifts from oligotrophic to mesotrophic organisms, including microbes (Wood 1993), an increase in bioerosion, and platform drowning (Hallock and Schlager 1986; Hallock 1988; Hallock et al. 1988). Ecologic changes indicated by differences in microbial carbonates allow sequence stratigraphy frameworks of carbonate platforms to be refined and sea-level curves to be evaluated (Whalen et al. 2002). Differences in growth forms, morphotypes, microfabrics and the association of microbialites with fossils and specific sedimentary structures have been very suc-
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cessfully used in recognizing depositional environments of Precambrian and Phanerozoic limestones and discriminating shallow- and deeper-water settings (Hoffman 1974, George 1999, Riding and Awramik 2000). Fig. 9.6 illustrates generalized distribution patterns of Mesozoic microbialites based on about thirty case studies. Marine deep-water microbialites, formed on slopes and deeper ramps at water depths of 100 to 150 m, become important in the Late Jurassic (e.g. Gygi 1992), but were already present in the Precambrian and Devonian. Aphotic stromatolites are known from Late Triassic and Early Jurassic basinal settings (Böhm and Brachert 1993).
9.1.5.3 Economic Importance of Stromatolites Microbes contribute significantly to the precipitation of a wide variety of authigenic minerals, including oxides, carbonates, phosphates, sulfides and silicates (see overviews by Mendelsohn 1976; Ferris 2000). Microbial metabolic activity can alter Eh and pH conditions, influence oxidation and reduction processes, and trigger mineral precipitation and fixation of dissolved metal ions, thus contributing to the formation of economically important ore deposits (Dexter-Dyer et al. 1984; Westbroek and DeJong 1983; Schidlowski 1985; Fan et al. 1999). Many ore deposits are closely associated with stromatolites, particularly with stromatolite reefs. Examples are deposits related to the Precambrian Banded Iron Formation in America (James 1992), stratiform copper deposits in Africa, or some lead-zinc deposits of the Mississippi Valley type in North America (see papers included in Wolf 1976 and Parnell et al. 1990). Cyanobacteria and other bacteria are directly or indirectly implicated in the genesis of iron and manganese deposits (Nealson 1983; Yin 1990).
9.2 Biogenic Encrustations Encrusting organisms play a key role in carbonate sedimentation and the development of reefs. Skeletonized and non-skeletonized encrusting organisms grow on various hard substrates represented by hardgrounds (see Sect. 5.2.4.1), cobbles/boulders and skeletons of organisms. This subchapter addresses biogenic crusts developed on free exposed surfaces of benthic host organisms, e.g. the upper surface of colonial fossils, or in hidden and shadow places, e.g. the undersides of colonies, or within
380
Msozoic Microbialites
Fig. 9.6. Generalized distribution of microbialites in Mesozoic platforms and reefs. Microbialites comprise non-laminar and laminar types as well as microbial oncoids (see Sect. 4.2.4.1). Cement crusts refer to quantitatively frequent, microbially induced marine carbonate cements. The Middle Triassic (predominantly Ladinian) and Late Triassic (predominantly Norian) scenarios reflect case studies in the European western Tethys. The Late Jurassic scenario summarizes case studies in the Tethys and in epicontinental settings. The Middle and Late Cretaceous sketch describes the distribution of microbialites in the Tethys. Note similarities in the setting of microbial oncoids, but differences in the microbial contribution to reefs and mud mounds. Modified and expanded from Leinfelder and Schmid (2000).
cavities or in borings (Sect. 9.3). The composition and sequential order of encrustation patterns are important microfacies criteria allowing the study of original ecological relationships and the interactions and competition between encrusters to be studied, environmental parameters, e.g. nutrient constraints to be estimated. Encrusting organisms occur in all shallow- and deepmarine settings, both in warm-water and cold-water environments. Encrusting taxa act as binders and cementing organisms, but also as constructors, and are often initial and early colonizers.
Encrusters are commonly included in the binder guild, but for many encrusting organisms a considerable guild overlap exists (Sect. 16.2.3.2). The dominance of encrusters is used in limestone classifications (Sect. 8.2). Terminology: Organisms living on a substrate can be divided into epibenthic and endobenthic, depending on whether they live above or below the substrate surface. A simple nomenclature system was proposed by Taylor and Wilson (2002) for marine plants and animals that encrust or bore natural hard substrates. The
Encrusters
terms describe the identity of the colonizing organism, the nature of the substrate, and the location of the colonists (on the surface or within the surface). The classification was inspired by the carbonate classification system used by Folk (Sect. 8.3.2). Each term consists of two or three roots: The last root refers to the identity of the colonizing organism (animal, -zoan; plant, -phyte, or either, -biont). The preceeding root describes the type of the substrate (rock, -litho-; animal, -zoo-; plant, -phyto-; or any organic hardpart of unknown or uncertain status, living or dead, -skeleto). The location of the colonizing organisms is indicated by the prefix epi- or endo-. The collective term sclerobiont is used for any kind of animal or plant fouling any kind of hard substrate. Sclerobionts include encrusters pressed closely to the surface of the substrate (Fig. 9.7), sessile organisms that are cemented or organically anchored to the substrate surface and grow into the water column (Fig. 9.9), and borers that enter the substrate.
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9.2.1 Criteria and Constraints of Encrusters Biogenic crusts are formed by sessile benthic organisms including microbes, algae, foraminifera, sponges, bryozoans, corals as well as serpulids and mollusks. Encrusting organisms are solitary or colonial. The latter exhibit various growth forms (Jackson 1979, Taylor 1992) comprising runners (linear forms, e.g. some bryozoans), sheets, mounds, plates (with punctuated bases), vines (linear to branched, semierect forms) and trees (erect forms). Sheets and mounds are common growth forms of reefal encrustations, producing crusts oriented parallel to the substrate (Fig. 9.7) and indicating strong competition for space. Erect upright encrustations (Fig. 9.9) point to low sediment input. The success of encrustation depends on a number of factors including the size, structural complexity and stability of the substrate, the biochemical conditioning,
Fig. 9.7. Characteristic hard substrate encrusters. The bivalve shell (B) was first encrusted by an association consisting of sessile foraminifera (SF; Miniacina), serpulids (Ditrupa, D) and squamariacean red algae (SRA; today called peyssonnelaceans). Corallinacean red algae (CRA) form the later phase of the encrustation. Note intensive boring of the shell (black arrows) prior to the biogenic encrustation. Borings are infilled with fine bioclastic sediment. By contrast, the borings in the squamariacean algae are occluded by cement (white arrow). Applying the terminology proposed by Taylor and Wilson (2002) for organisms inhabiting hard substrates, the crust sequence consists of more or less synchronous episclerozoans and episclerophytes represented by squamariacean red algae, followed by epizoophytes represented by corallinacean red algae. Late Tertiary (Middle Miocene): St. Margarethen, Burgenland, Austria. Scale is 5 mm.
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Biogenic Encrustations
Plate 51 Biogenic Encrustations Biogenic encrustations originate from the growth of sessile organisms on various substrates. These epibionts grow on animal hosts (epizoans) or on plant hosts (epiphytes, e.g. sea grass or soft algae). Epibionts from ephemeral substrates are generally rare, in contrast to the considerable record of epibionts from hard substrates. Crust-forming calcareous organisms occur on living and dead surfaces of skeletons, particularly in reef environments, and on lithified rock surfaces (e.g. hardgrounds, see Sect. 5.2.4.1). Calcareous encrusters include microbes and algae (see Pl. 50), adherent foraminifera, sponges, corals, bryozoans, serpulids and some mollusks. The frequency, composition and diversity of crust-building communities have changed over time. Sessile foraminifera (–> 1, 2, 4, 7, 8) belonging to different suborders are common in Mesozoic and Cenozoic reef and shelf faunas (Adams 1962). Some taxa contributed significantly to the formation of Tertiary reefs (Plaziat and Perrin 1992). Attached foraminifera are valuable proxies for sedimentation rates, water energy and paleo-water depths (Ghose 1977; Martindale 1992). Biogenic encrustations take part in the stabilization of sediment, the rapid lithification of reef slopes (Keim and Schlager 1999), the formation of carbonate grains (–> 2, 4) and the construction of rigid reefs (–> 7). Major controls on the occurrence and distribution of encrusters are substrate, availability, water energy and (low) sedimentation rates. Succession analyses of encrusting associations provide clues to substrate types, sedimentation rates, paleo-water depths (Rasser and Piller 1997) and associated sea-level fluctuations (Reid and Macintyre 1988). Biogenic encrustations can be more easily described by using the simple classification system proposed by Taylor and Wilson (2002). Each term consists of two or three roots: The location of the colonizing organisms is indicated by the prefix epi- or endo-. The second root refers to the type of the substrate (rock, -litho-; animal, -zoo-; plant, -phyto-; or any organic hardpart of unknown or uncertain status, living or dead, -sclero-). The last root describes the identity of the colonizing organism (animal, -zoan; plant, -phyte, or either, -biont). Examples are given on this plate. 1 Attached agglutinated foraminifera (Ammovertella Cushman, Carboniferous-Permian). Note the flat lower surface of the tests representing adaption to an encrusting mode of life (e.g. on not preserved, poorly calcified algal blades). The foraminifera were probably epiphytozoans. Early Permian: Carnic Alps, Austria. 2 Foraminiferal encrustations (Tolypammina Rhumbler) attached to shells. Episkeletozoan crust. Long-lasting encrustations lead to the formation of foraminiferal oncoids. Late Triassic (Carnian): Central Carinthia, Austria. 3 Encrustations of sphinctozoid sponges (Uvanella; arrows) on and between corals facilitate the construction of rigid reef frameworks. Epizoozoan crust. Late Triassic (Hawasina Formation, Norian): Central Oman. 4 Formation of composite grains by encrustation of ooids with nubeculariid foraminifera and cyanobacteria. Nubeculariids are miliolid porcelaneous foraminifera, often attached to sedimentary grains or shells. They contribute to the formation of oncoids and Tertiary reef structures. The crust type can be classified as epilithozoan crust or more precisely as epizoozoan crust. Late Triassic (Late Carnian): Central Carinthia, Austria. 5 Encrustations of serpulids around a vanished recrystallized structure, probably a shell infilled with peloids. Episkeletobiont crust. The serpulid encrustation indicates the existence of a hard substrate. Early Cretaceous: Ostermünchen well, Southern Bavaria, Germany. 6 Repeated biogenic encrustations form characteristic sequences consisting of two and more taxonomically different organisms. Example: Cystoporid bryozoans (Dybowskiella, B) growing upon (now recrystallized) solenoporacean red algae (SO) are overgrown by sphinctozoid sponges (S). The relations between encrusting organisms include epiphytozoans followed by episkeletozoan crust types. Middle Permian reef limestone: Straza near Bled, Slovenia. 7 Modern reef limestone. Note the interaction of different encrusting organisms in the formation of an autochthonous boundstone structure: A meshwork resembling the microproblematicum Bacinella (B), interpreted as being of cyanobacteria origin is encrusted by corallinacean red algae (RA); followed by attached agglutinated foraminifera (F). The development of the encrusters was disturbed several times by the influx of ooid sand. San Salvador, Bahamas. 8 Encrustation of foraminifera (Palaeonubecularia) on a phylloid alga. Epiphytozoan crust. Note the difference in the preservation of the primary Mg-calcitic foraminifera and the primary aragonitic alga which has lost almost all its significant features due to the recrystallization of its aragonitic skeleton. Early Permian: Forni Avoltri, Carnia, Italy. 9 Alternations of calcimicrobes (Rothpletzella, R, Pl. 53/8) and fossils interpreted alternatively as foraminifera or cyanobacteria (Wetheredella Wood, W) are important in the formation of Paleozoic oncoids. Many Early Paleozoic biogenic crusts were built predominantly by microbes, sponges and stromatoporoids. Early Devonian (Emsian): Lake Wolaya, Carnic Alps, Austria.
Plate 51: Biogenic Encrustations
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384
Phanerozoic Encrusters
9.2.2 Phanerozoic Encrusters
Fig. 9.8. Rugose corals (Acanthophyllum sp.) encrusted by stromatoporoids (Clathrocoilona sp.) and thin layers composed of porostromate cyanobacteria (Girvanella sp.; not recognizable in the picture). Stromatoporoids and cyanobacteria form sheetlike crusts around the corals. The wide distribution of this crust pattern in Middle Devonian platform carbonates suggests that encrustation took place during the lifetime of the coral. Middle Devonian: Letmathe, Sauerland, Germany. Scale is 5 mm.
and the presence of biofilms. The available attachment surface area is critical for the survival of crust-building organisms. The biological potential of many encrusters is strongly microbially controlled. Colonizer-inherent factors are larval recruitment patterns and growth strategies. Ecosystem-inherent factors include sedimentation rates, water movement, predation/grazing pressure and disturbance frequency. In general, the dominance of encrusters with skeletons is reduced with depth, leading to complex succession patterns controlled by endogenic and exogenic factors. The former includes reproduction and larval strategies as well as competition for space and nutrients. Exogenic controls are mainly the size of the encrusting surface and predation pressure. Upper and lower surfaces of organisms acting as substrates for encrusters are likely to be subject to different light intensity, turbulence, grazing pressure and dissolved oxygen. These differences influence the distribution of generalists and specialists and cause faunal polarity between crusts developed upon upper and lower surfaces of host organisms. Lower surfaces often exhibit greater areal cover, higher density and a higher diversity of encrusters than upper surfaces.
Encrusting organisms range in size from less than a millimeter to several millimeters (cyanobacteria, foraminifera, many microproblematica) up to centimeters (red algae, stromatoporoids, corals). Micro-encrusters are encrusting microfossils, often occurring in close association with microbialites and within marine and nonmarine oncoids (Sect. 4.2.4.1). Cambrian and Ordovician encrusting communities lived predominantly in sheltered reef cavities (Kobluk 1988) and are dominated by microbialites, particularly cyanobacteria. Other common encrusters are bryozoans and sponges. The differentiation between encrusters occurring on exposed surfaces, cryptic habitats or in shadowed cavities seems to have taken place in the Ordovician. In Silurian and Devonian reef limestones, biogenic crusts on the surface of benthic organisms are composed of stromatoporoid sponges (Fig. 9.8), tabulate corals and porostromate cyanobacteria. Strong evidence for spatial competition among ancient marine hard substrate biotas comes from the study of overgrowth networks among Silurian bryozoans (Lidell and Brett 1982; Taylor 1984). Microbes, algae, smaller foraminifera, bryozoans and chaetetid sponges are important constituents of Carboniferous and Permian encrusting communities. Mesozoic reef and platform carbonates yield diverse encruster communities whose taxonomic diversity and ecologic complexity increase in the Late Jurassic and the Cretaceous. Tertiary succession patterns resemble those of modern crust communities, due to the dominance of red algae, foraminifera, bryozoans, barnacles and serpulids (Fig. 9.7 and Fig. 9.9).
9.2.3 Significance of Encrustation Patterns in Recognizing Depositional Settings and Environmental Controls The combined use of encrustation patterns, microbialite types and bioerosion criteria can serve as a guide to identifying inner and outer parts of carbonate platforms and ramps, as well as interpreting paleo-water depths and nutrient levels in ancient reefs. Prerequisites for the successful application of these criteria are (1) the knowledge of changes in the composition of biogenic crusts during time, and (2) the differentiation of micro-encrusters, if possible at low taxonomic levels. Changes in the composition of Triassic biotic crusts: Middle and Late Triassic crusts exhibit distinct differ-
Encrustation Patterns
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Jurassic micro-encruster associations subdividing carbonate ramp depositional environments: Schmid (1996) studying Late Jurassic reef and ramp carbonates in different parts of Europe recognized five microencruster associations characterized by composition and diversity (Fig. 9.10). Girvanella associations within oncoids are restricted to near coastal settings with fluctuating salinities.The widespread Bacinella-Lithocodium association occurs in shallow lagoonal and reefal settings affected by moderate environmental stress. Balanced environmental conditions in shallow lagoonal and reefal areas are indicated by high-diversity micro-encruster associations consisting of porostromate microbes (Cayeuxia, Girvanella), red algae (Marinella) and microproblematica (Bacinella, Thaumatoporella). The Tubiphytes-Koskinobullina association is abundant in mid-ramp positions, but also occurs in reef cavities. Low-ramp and middle- to outer-ramp positions are characterized by the low-diversity Terebella (a serpulid)Tubiphytes association that thrives in low-energy environments and appears to tolerate oxygen deficits at the sea bottom. Fig. 9.9. Encrustation of colonial scleractinian corals by corallinacean red algae (Archaeolithothamnium gosaviense Rothpletz). Both, the corals and the algae were attacked by macroborers. The coral was strongly eroded and bored prior to the overgrowth by the algae. The transition from moundlike to erect growth indicates quiet-water conditions and low and slow sediment input. Lagoonal back-reef environment. Late Cretaceous (Coniacian): Northern Alps, Bad Reichenhall, Bavaria, Germany. Scale is 2 mm. Courtesy of R. Höfling, Erlangen.
ences in the association and diversity of organisms that contributed to the formation of crusts in reef environments (Flügel 2002). Anisian crusts are rare but diverse; their encrusting organisms comprise sphinctozoid sponges, porostromate cyanobacteria, foraminifera, serpulids and some microproblematica. The latter are common and sometimes abundant in younger Triassic but also in Jurassic reef crusts. Crusts are frequent in Ladinian and Early Carnian reefs; they are characterized by various associations of sponges, microbes, and microbially induced carbonate cement. The composition of Late Carnian reefs differs significantly in that lowgrowing sponges and porostromate cyanobacteria dominate (Pl. 79/1). A complete change took place in the Norian (Late Triassic): Biogenic crusts become highly diverse and mainly built by microbialites (Pl. 116/2), sphinctozoan and chaetetid sponges (Pl. 81/1), and micro-encruster associations, which also dominate in Late Jurassic and Early Cretaceous reefs (e.g. Lithocodium/ Bacinella; Pl. 99/4, 5).
Nutritional models derived from microbialites, micro-encrusters, bioerosion, and faunal diversity: The paper by Dupraz and Strasser (2002) on coral-microbialite reefs from the Swiss Late Jurassic illustrates the high potential of microbial structures, encrusters and borings for evaluating changes in dominating nutrient regimes. Micro-encrusters associated with microbialites and reef-building corals include here red algae, foraminifera, bryozoans, serpulids and sponges. The semiquantitative micro- and macroscale distribution patterns indicate three trophic structures that parallel the different siliciclastic influx into the reef areas: Phototrophicdominated structures (indicated by Lithocodium-Bacinella and opportunistic serpulids and bryozoans) dominate in moderately diverse coral reefs growing in poor carbonate and nutrient-poor environments with only low bioerosion. Balanced phototrophic-heterotrophic structures developed in mixed siliciclastic platform environments and produced the most diversified coral reefs with common thrombolites. Heterotrophic structures dominated in the case of strong siliciclastic accumulation or a strong increase in nutrient availability; thrombolites, Terebella and other serpulids, and bryozoans are common. Bioerosion is moderate to high. Late Cretaceous encruster associations dependent on the reef type: Different encruster associations occur in reefs built by different reef organisms as exemplified by coral-stromatoporoid reefs, coral reefs, rudist reefs and algal reefs (Moussavian 1992). Encrusters are
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Encrusters in Jurassic Reefs
Fig. 9.10. Distribution patterns of micro-encruster associations in Late Jurassic reefs developed in different parts of carbonate ramps. Modified from Schmid (1996).
cyanobacteria, rhodophycean algae, various groups of sessile foraminifera (e.g. acervulinids), corals, hydrozoans, sponges including stromatoporoids and chaetetids, bryozoans, and microproblematica (e.g. Lithocodium and Bacinella).
9.3 Bioerosion, Boring and Grazing Organisms This chapter deals with the destruction of carbonates substrate by boring organisms and with the nature, intensity and significance of bioerosion. The term bioerosion (Neumann 1966) describes the processes by which biological activity destroys hard substrates. Bioerosion contributes significantly to sediment production, controls early diagenetic processes, and the taphonomic history of fossils. Many fine-grained carbonate particles, seen in thin sections, may be the result of biodegradation by boring. Especially particles in the silt-size range are candidates for this interpretation, as shown by SEM studies. In coral reefs, bioerosion depends on numerous environmental factors such as light availability, nutrient supply, and water depth. Boring occurs in shallowand deep-marine, freshwater and terrestrial environments (contributing to phytokarst formation and the degradation of limestones and marbles used as building stones and for art sculptures (e.g. Anagnostidis et
al. 1983), and affects carbonate rocks, life or dead skeletons as well as hardened sea floors. Endolithic borers (algae, polychaetes) and grazers (e.g. herbivorous fishes and sea urchins) modify and redistribute carbonate produced by reef-building organisms. The relative importance of endolithic borers and grazers varies between reef environments and with time of exposure (Kiene and Hutchins 1994). Macroborers are responsible for extensive substrate erosion, sediment production and the generation of secondary porosity in reef frameworks. Bioerosion occurs in marine warm-water and coldwater carbonate environments and is controlled by illumination, nutrients, sedimentation rate, siliciclastic influx and water depth, as drastically shown by the compositional difference of boring associations in shallow warm and deep cold-water environments. The latter, exemplified by deep-water Lophelia reefs in the Atlantic are characterized by the abundance of non-autotrophic bacteria, heterotrophic fungi as well as clionid sponges and bryozoans. Sediment production is high (Freiwald 1998), sometimes competing with figures known from shallow tropical environments (up to 0.5 to 2.5 cm/year). The interest in bioerosion and boring organisms has increased significantly during the last few decades. A bibliography of papers dealing with modern and fossil micro- and macroborers includes more than 800 references (Radtke et al. 1997).
Bioerosion, Boring, Grazing
Micro- and macroborers produce trace fossils The traces of micro- and macroborers are known throughout the Phanerozoic. Microborers at a scale < 1 µm to about 100 µm tunnel their way into the substrate using weak acids. They include bacteria, cyanobacteria, fungi as well as foraminifera and bryozoans. Macroborers are visible with the naked eye, and have diameters of 1 mm and greater. They use physical and physicochemical techniques for boring. Endolithic microborers attack hard substrates (skeletons, hardgrounds, carbonate grains, rocks) and produce borings that are filled with cement or sediment, or remain open. The products of boring activities are tunnels and caves whose shape may correspond to the morphology of the borers, thus allowing taxonomic assignments of the trace fossils. Most borings excavated by bacteria, algae or fungi are too small to be successfully studied in thin sections. Studies dealing with modern and ancient microborers are mainly based on SEM investigations of artificial casts produced by filling the boreholes with resin (Golubic et al. 1970, 1983; Pl. 52/ 1-4). Some of the macroborings can be morphologically differentiated in thin sections and attributed to systematic groups (Box 9.4). Grazers, scrapers and swallowers Bioeroding organisms include not only borers, but also grazers, scrapers and swallowers, which have a major impact on the bioerosion of modern shallowmarine carbonates (Hutchins 1986). Grazers (e.g. echinoids, gastropods, chitons, and fishes) are organisms feeding mainly on plants, e.g. algal and microbial mats. Most grazers search for algae in living or dead substrates. Grazing echinoids are major eroders of modern coral reefs. Chitons, already known from the Late Cambrian, are common grazers of endo- and epilithic algae along carbonate coastlines (Rasmussen and Frankenberg 1990). Carbonate removal by grazers may be responsible for a high percentage of the total bioerosion (Chazottes et al. 1995). Intensive microbial mat grazing started in the Late Precambrian. Grazing by conodontophorids and fishes may have become more intensive in the Middle and Late Cambrian. Grazing fishes are the dominant herbivores in most tropical environments. Herbivorous grazing fish species are important bioeroders on tropical shallow reefs (up to 7 kg/m2/year). Substrates already infested by boring algae and of low skeletal density are eroded fastest (Bruggemann et al. 1995). Scarid parrotfish and acanthurid surgeonfish, feeding by scraping off algae that grow on and in dead coral skeletons, have a major im-
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Fig. 9.11. Sponge borings dominate in modern and Neogene environments, but are already known in the Cambrian. Sponges are capable of destroying large amounts of reef frameworks. The Modiola mussel is intensively bored by a sponge (Cliona) characterized by a network structure. The borings consist of a series of galleries branching and anastomosing within the substrate and communicating with the surface via numerous small round pores that house the inhalant and exhalant papillae of the sponge. The sponge tissue has specialized etching edges responsible for boring, which results in freeing of characteristically shaped chips. These chips with a size of 40-60 µm are expelled by the oscula und carried away by currents. Clionid borings preserved as trace fossils are called Entobia. X-ray photograph. Scale is 2 cm.
pact on the destruction of modern reef surfaces. The first unequivocal fossil representative of the Scaridae is known from the Middle Miocene (Bellwood and Schultz 1991); scarids may represent ‘the reefs latest trophic innovation’ (Bruggemann 1994). Scrapers (e.g. regular echinoids, limpids, some parrot fishes) use beaks or claws to scratch skeletons in order to remove microbial nutrients or algal/microbial cover. Regular echinoids are known since the Carboniferous. Swallowers are detritus feeders that swallow sediment grains that are comminuted through their digestive tracks, usually producing pellets (e.g. holothurians, jawed fishes). Holothurians are known since the Middle Cambrian, jawed fishes since the Wenlockian.
9.3.1 Recent and Fossil Microborers Microboring is not only a product of light-dependent photosynthetic algae but also of heterotroph bacteria and fungi. Therefore, microborings have a wide bathymetric distribution ranging from supratidal and intertidal environments to subtidal and deeper-water environments. Many microborers are concentrated within shallow subtidal and intertidal depths and exhibit distinct zonation patterns. Endolithic biodegradation is common and highly diverse in near-coastal environments and warm
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Microborers
shallow seas, but is also important in cool-water skeletal shelf carbonates (Akpan and Farrow 1984; Young and Nelson 1988). Microborers are studied in order to understand • destructive processes resulting in the degradation of skeletal elements and the production and accumulation of fine-grained carbonate sediment, • processes responsible for the formation of micrite envelopes (cortoids; Pl. 52/6), • distributional patterns that can be used to differentiate modern and ancient marine environments (e.g. intertidal and supratidal sequences, Hoffman 1985), • the role of microborers as (paleo-)bathymetric and (paleo-)ecologic indicators, • the role of increased nutrient input to reefs that might result in reef demise (Hallock 1988; Wood 1993). Microboring groups Microborers include bacteria, algae and fungi that penetrate substrate by etching. Boring foraminifera and bryozoans are attributed to microborers, because the dimensions of the trace fossils are smaller than 100 µm. Bacteria: Non-photosynthetic bacteria and photosynthetic cyanobacteria occur in marine, limnic, fluvial and terrestrial environments. Both groups include the oldest microborers known since the Precambrian and recorded in carbonate rocks as old as 1700 Ma (Zhang and Golubic 1987). Bacteria have various shapes (spheres, bars, bent bars and spirals). The size of non-photosynthetic bacteria ranges between 1 and 5 µm with diameters between 0.2 and 1 µm. Modern endolithic cyanobacteria are represented by more than 20 genera (Budd and Perkins 1980). Green algae occur in marine and non-marine environments. Some chlorophycean algae live within carbonate substrates or have endolithic stages (Golubic et al. 1975); they are already recorded from Early Paleozoic skeletal grains and ooids (Podhalanska 1984) and possibly also from Precambrian rocks. Red algae are predominantly marine. Endolithic life stages are known from only a few marine genera that already occur in Early Paleozoic rocks. Fungi are heterotroph organisms without photosynthetic pigments. Boring fungi occur in shallow and deep marine environments down to 5000 m. Their trace fossils are differentiated by the shape, size, mode of branching and sporangia (Glaub 1994). The oldest records of boring fungi are from the Middle Ordovician. Foraminifera: Boring endolithic foraminifera appear to be of greater importance for the degradation of skeletal grains than previously assumed (Freiwald and
Fig. 9.12. Recent cold-water microborings produced by marine fungi attacking the deep-water coral Lophelia. Microborings are not restricted to shallow-marine environments (see Pl. 51/1-4), but also occur in the deep sea where bacteria, fungi and bryozoans act as borers (Freiwald and Wilson 1998). The abundance of fungal borings increases with water depth from deeper subtidal to bathyal environments (see Fig. 9.16). The SEM photograph of the resin cast shows linear fungal borings and globular sporocysts developed within 70 µm below the surface of the coral. Propeller Mound, Northern Atlantic, west of Ireland. Water depth about 700 m, water temperature 10 °C. Courtesy of L. Beuck, Erlangen.
Schönfeld 1996). They occur in groups characterized by agglutinated and calcareous perforate tests; they are known from cold-water and warm-water environments, and are reported from Jurassic to Quaternary carbonates (Venec-Pyre 1987; Baumfalk et al. 1983; Smyth 1988; Cherchi et al. 1990; Cherchi and Schroeder 1991). Bryozoans: Trace fossils of boring bryozoans can be easily recognized in SEM images, because the trace closely resembles the producer morphology and reflects the shape of the body fossil. The trace corresponds to a network that connects elongated chambers. These bryozoans are differentiated by the shape, size and position of the zooids, the relation between zooids and the stolonial system, and the size of surface openings (Pohowsky 1978). The diameter of the non-calcified slender structural elements is less than 10 µm. Most boring bryozoans known from Jurassic to Quaternary limestones are attributed to the Ctenostomata. The Paleozoic record of boring bryozoans is scanty.
9.3.2 Recent and Fossil Macroborers Macroborings occur in marine and non-marine environments and contribute considerably to bioerosion in marine shallow- and deeper-water, warm- and coldwater settings (Ekdale et al. 1989). The most effective
Macroborers
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Fig. 9.13. Macroborings produced by cirripedians of the group Acrothoracica. The borings are characterized by a amphoralike shape with a narrow neck at the anterior end. The borings occur within a reef limestone with abundant bryozoans (celllike structures). Note the reversed geopetal fabric in the borings indicating overturning of parts of the reef mound. These are the oldest Mesozoic cirripedian borings subsequent to the Permian-Triassic extinction event. The association of ascothoracid borings and trepostomate bryozoans in the Late Permian reefs and in the Anisian underlines the Paleozoic aspect of the early Middle Triassic. Middle Triassic (Anisian): Olang Dolomites, Southern Alps, Italy. Modified from Senowbari-Daryan et al. (1993). Scale is 2 mm.
groups in modern environments are sponges, bivalves and ‘worms’, followed by specialized cirripedians (Fig. 9.13). These groups were also the principal macroborers in the Mesozoic and Cenozoic. The identification of the producer of macroborings in thin sections is difficult, but should by tried using the morphological criteria summarized in Box 9.4. Macroborers are responsible for extensive substrate destruction, sediment production and the generation of secondary porosity in reef frameworks.
Bivalves: Boring into reef-building organisms by shells is extremely well developed within three families of bivalves (Warme 1975; Kleemann 1980, 1983, 1990; Bromley 1981). The Pholadidae are exclusively borers primarily using mechanical means of penetration. The Gastrochaeniidae and some Mytilidae penetrate predominantly calcareous substrates using mechanical and chemical means. Bivalve borings have a major impact on the destruction of littoral sediments,
Macroboring groups Sponges: Boring sponges have received the most attention of all groups of boring groups because of their role in bioerosion processes, sediment production and calcium carbonate dissolution (Neumann 1966; Futterer 1974; Rützler 1975; Torunski 1979; Reitner and Keupp 1991; Acker and Risk 1985). Boring sponges (Fig. 9.11) form large chambers, with smaller galleries branching off the main chambers. Bromley (1978) noted that the cavities can be identified even after fossilization, allowing sponge borings to be distinguished from other borings. Demospongid borings are differentiated according to the size of the external openings and the chamber system (Bromley and D’Alessandro 1984). Traces of sponge borings resembling present day Demospongiae have been found in Early Cambrian archaeocyathid reefs and are common in Permian to Cenozoic carbonates.
Box 9.4. Some morphological trace fossil criteria of macroborer groups (after Perry and Bertling 2000). Reviews assisting in the identification of borings were published by Warme (1975), Bromley (1978), Bromley and D’Alessandro (1984, 1987), Pleydell and Jones (1988), and Fürsich et al. (1994). Sponges: Complex branched multi-apertured networks of large chambers; multi-apertured irregular network without chambers; stellate structures with single aperture. Polychaete worms: Cylindrical borings. Round or dumbbell-shaped cross section. Mostly single-apertured. Sipunculids: Cylindrical, straight to gently curved, round cross sections, single aperture. Phoronids: Cylindrical and U-shaped, multi-apertures, or small, branched with cylindrical branches, multiapertured. Bivalves: Flask-shaped. Single apertured. Cirripedians: Small amphora-like borings with a slitlike apertura in the narrow exterior part.
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Bioerosion
the degradation of reefs (Scott and Risk 1988), and diagenetic processes affecting reef organisms (Jones and Pemberton 1988). Bivalve borings can be recognized in thin sections by vertical or oblique boreholes that often contain remains of the rock-boring shell (Pl. 52/8). Records of bivalve borings in hardgrounds, reefs and fossils are known from the Ordovician to the Quaternary, but do not become more abundant until the Triassic.
contribute to bioerosion. Millimeter-sized holes believed to have been made by gastropods have been reported from the Paleozoic, but these data are controversially discussed. Undisputable drill holes of carnivorous naticid gastropods occur in the Late Triassic. Gastropod drilling is very rare in the Jurassic to the Early Cretaceous interval. Starting with the Late Cretaceous gastropod drilling in mollusk shells become abundant (Kowalewski et al. 1998).
Gastropods: Carnivorous gastropods drilling into various fossils are predators rather than borers, but still
’Worms’: Trace fossils referred to boring activities of worms occur throughout the Phanerozoic and are
Plate 52 Bioerosion and Boring Organisms Bioerosion has a tremendous impact on carbonate sedimentation and carbonate diagenesis. The destruction of hard substrates by boring, rasping, grazing and browsing organisms is substantial in the breakdown of carbonate skeletons and for the production of fine-grained to sand-sized sediments both in warm-water and cool-water environments. A huge variety of microbes, plants and animals penetrate hard surfaces. Submarine bioerosion increases the porosity and surface area of skeletons and makes them more susceptible to dissolution. Microborers are highly effective in the biological degradation of subaerially exposed carbonate rocks (e.g. karst) and of limestones used as building stones. Microborers (–> 1-4) include autotroph (–> 1,2) and heterotroph bacteria, green algae (–> 3) and red algae, fungi (–> 4), as well as foraminifera and bryozoans. Important macroborers are sponges (e.g. clionids; Fig. 9.1.1), bivalves, ‘worms’ and cirripedian arthropods (Fig. 9.13). Microborers are known since the Proterozoic, the oldest macroborers have been reported from the Early Cambrian. Microborings are studied by stereoscan electron microscopy. SEM photos (–> 1-4) exhibit microborings of recent bivalve shells which were collected from an upper shelf slope covered by reefs and bioclastic sands (Günther 1990). The samples were first impregnated by plastic material, followed by dissolution of the carbonate so that casts of the boreholes can be studied. Macroborings are described and differentiated in thin sections and rock and fossil samples. Many groups produce morphologically distinct borings which are valuable in evaluating paleoenvironmental conditions. Distribution patterns of microborers provide a highly promising tool for the differentiation of paleodepth zones. 1 Modern shell-boring cyanobacteria (Mastigocoleus testarum Lagerheim). Long arched tunnels (Ø 6–10 µm) growing in all directions and forming a dense network. Short branches with terminal heterocysts (H). Generally, cyanobacterial microborings are concentrated in shallower waters. Cozumel shelf, Yucatan, Mexico. Water depth 42 m. 2 Modern shell-boring cyanobacteria (Hyella gigas Lucas and Golubic). Thick short filaments with rounded tips, arranged in clusters. Ø 10–40 µm. Upper and lower photic zone. Cozumel. Water depth 10 m. 3 Modern shell-boring green algae (Phaeophila engleri Reinke). Filaments are characteristically branched exhibiting distinct sporangia (arrows). Boring in the uppermost layer of a bivalve shell. Note the prismatic shell microstructure. Generally, green algal borings are common in shallower waters. Upper part of the photic zone. Cozumel. Water depth 2.5 m. 4 Modern shell-boring fungi. Spherical and stalked sporangia, oriented to the shell surface. Thin hyphae connecting the sporangia. Fungal borings occur in shallow-marine and deeper marine environments but increase in abundance in deeper waters. Lower photic zone. Cozumel. Water depth 20 m. 5 Macroborings in a brachiopod shell. The primary mineralogy of the brachiopod shell was Low-Mg calcite. Note differences in size, orientation and shape of the bore holes which may be attributed to the Trypanites group regarded as being excavated by polychaete worms. Early Devonian (Emsian): Anti-Atlas, southern Morocco. 6 Micrite envelope forming cortoids: Reworked rounded echinoderm clasts are surrounded by thin micritic linings, caused by multiple microboring (arrows), and subsequent infilling of the bore holes with microcrystalline calcite cement (see Sect. 4.1.2.3). The microstructure of the echinoderm fragments characterized by reticulate pattern and cleavage planes partly preserved and partly obliterated by recrystallization of High-Mg calcite to Low-Mg calcite. D: Dasyclad green algae. Early Permian: Carnia, Southern Alps, Italy. 7 Borings in crinoid bioclasts. The echinoderm microstructure is still preserved. Note the different infilling and the close setting of the boreholes. Middle Triassic (Muschelkalk): Southwestern Germany. 8 Lithophagid bivalve borings in colonial corals. Remains of the bivalves are still preserved within the boreholes (arrow). Late Jurassic (Kimmeridgian): Lower Saxony, northern Germany.
Plate 52: Boring Organisms
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abundant in Mesozoic and Cenozoic limestones. The oldest representatives occur in Early Cambrian reefs (James and Debrenne 1980; Rowland and Gangloff 1988). Many fossil records are attributed to marine annelid worms, specifically to polychaete worms, which include several rock- and wood-boring families. In thin sections borings attributed to polychaete annelids exhibit circular or thumb-shaped cross sections. The boring extends as a long cylindrical tube from a single entrance. A common trace fossil often (but not unequivocally) interpreted as borings of polychaete worms is Trypanites, first described from the German Triassic, and now known from the Early Cambrian to the recent. These borings may be observed in thin sections of reef-building organisms (e.g. stromatoporoids, rugose corals, bryozoans; Pemberton et al. 1988) and reef sediments as well as in hardgrounds (Palmer 1982) and at discontinuity surfaces (Pemberton et al. 1980). Trypanites is characterized by simple unbranched vertical to sinuous borings with a single opening to the surface, and with or without a flared entrance. The diameter (< 1 to about 2 mm) is more or less the same throughout the entire length of the boring which varies between a few millimeters to more than 20 millimeters. Geopetal sediment fillings within the boring may bear testimony to a reorientation of the bored fossils. Cirripedians: Some ascothoracican Cirripedia living in shallow seas create cavities in skeletons (e.g. corals, bivalves, echinoids) and various carbonate substrates including hardgrounds. Records of Paleozoic cirripedian borings (Webb 1987) are rare as compared with the Mesozoic (Fig. 9.13; Fürsich and Wendt 1977) and the Cenozoic. The millimeter-sized borings are shoeor sack-like with a narrow opening at a neck to the exterior. A slit-like extension at the anterior end gives the aperture the shape of a comma. Phoronids and sipunculids: Less common macroborer trace fossils are interpreted as being produced by phoronids and sipunculids. Phoronid borings have been described from the German Triassic Muschelkalk and from the Late Cretaceous chalk. The borings correspond to small irregularly winding tubes with diameters less than 0.5 mm. Modern sipunculid borings are common in coastal rocks and reef limestones. They appear as clubs in thin sections (Rice and Macintyre 1972). If all or some Trypanites borings were attributed to spirunculids instead of polychaete worms, the stratigraphic record would range from the Cambrian to the Quaternary.
Evolution of Borers
9.3.3 Micro- and Macroboring through Time Major boring organisms have very long ranges (Fig. 9.14), but the contribution of the groups varied strongly during the Phanerozoic. Older bioerosion processes that can be qualitatively and quantitatively compared with modern processes are relatively young and have been effective since the Late Tertiary. It has been suggested that bioerosion, predation and grazing features were only fully present in modern Mesozoic-Cenozoic reefs (Wood 1998, 1999). The biological destruction of recent coral reefs is an impressive example of the overwhelming role of organisms in degradation and recycling of carbonate sediments (Bromley 1978; Hutchings 1986). Although microborers have been known since the Precambrian and macroborers since the Cambrian, processes comparable in quality and quantity with the biological destruction in modern reefs seem to have developed relatively late in the Phanerozoic and did not reach their full intensity in reefs before the Late Tertiary (Fig. 9.15). Examples, indicating strong bioerosion are known however, from Early Paleozoic stromatoporoid reefs (Nield 1984), from Triassic coral reefs (Stanton and Flügel 1989) and from Jurassic sponge and coral reefs (Pisera 1987; Reitner and Keupp 1991; Flügel 1993).
9.3.3.1 Qualitative Changes in Micro- and Macroborer Groups Microborers were important in Paleozoic and postPaleozoic times, but the impact and the composition of macroborer communities seem to have increased during the Mesozoic and the Cenozoic, as exemplified by Fig. 9.15. The oldest macroborers in the Early Cambrian and Ordovician occurred in reef environments. In contrast, the most intensive boring in the Silurian took place not in reefs but in deeper-water muddy ramp settings. Copper (2002) discussed the possibility that boring developed in deeper shelf settings before invading shallow reef platforms. Macroboring occurred in Early and Middle Devonian reefs, but is only poorly documented in the Famennian and was not a dominant process in Carboniferous and Early Permian reefs (Webb 2002; Wahlman 2002). The rate of the destruction of Late Permian reefs by borers and grazers was low, because free substrates were rapidly occupied by encrusting organisms. Macroborers (predominantly sponges) were more important than microborers (Weidlich 2002). Interestingly, the composition of microborers did not change from the Permian to the Triassic, despite significant changes in the reefal host
Macroboring in Reefs MACROBORERS
Cyanobacteria Green algae Red algae Fungi Foraminifera Bryozoans Sponges Bivalves Gastropods ‘Worms‘ Cirripedia Phoronids
MICROBORERS
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Quaternary Tertiary Cretaceous Jurassic Triassic Permian Carboniferous Devonian Silurian Ordovician Cambrian
organisms, and remained similar during the Mesozoic and Cenozoic (Schmidt 1992; Balog 1996; Vogel et al. 1999). No macroborers are known from Early Triassic carbonate platforms. Anisian macroborers are bivalves and cirripedians. Macroborer diversity and abundance increased significantly in Norian and Rhaetian reefs,
Fig. 9.14. Fossil record of micro- and macroborers and substrate excavating grazers destroying carbonate skeletons and lithified substrates. The fossil record of boring organisms can be estimated through trace fossils. In many cases a reasonable delineation of the responsible group or borer is possible, but often the assignment of the trace fossils remains controversial. The range chart summarizes borings in different settings (reefs, hardgrounds). In addition, the presence of drill holes in marine invertebrate shells attacked by predators (mainly gastropods) is indicated. Black columns refer to times where the group contributed greatly to bioerosional processes; open columns refer to times with a less significant influence of borers. Note that the columns also include time intervals with no or questionable records. Based on Kobluk et al. (1978), Vogel (1993), Kowalewski et al. (1998), Perry and Bertling (2000) and other sources.
possibly indicating a parallel evolution of Triassic reefbuilding corals and macroborers (Perry and Bertling 2000). Norian and Rhaetian reefs were attacked by bivalves; sponge boring activity increased in the Rhaetian. The contribution of macroborers to bioerosion increased strongly during the Jurassic (Leinfelder et al. 2002). In the Liassic bivalves became the dominant group. Middle and Late Jurassic reefs suffered from boring sponges, bivalves and worms. Boring gastro-
Fig. 9.15. Impact of micro- and macroborings on reefs through time: The graph is based on the Erlangen PaleoReef Database (see Kiessling and Flügel 2002). Bioerosion was low to moderate in Paleozoic reefs, increased significantly during the Triassic, and maintained high but fluctuating levels from the Jurassic to the Recent. The Mesozoic increase may be explained by the coeval diversification of predators. Caution should be exercised when using the graph: The Early Triassic decline is caused by the single occurrence and rare records of microbial reefs, and the complete lack of metazoan reefs. These microbial reefs were not attacked by microborers; macroborers were not described. Grazers were abundant, as indicated by common herbivorous gastropods. The decline of microborers in the Middle Jurassic was probably caused by the low number of reefs in the database. Similarly, the decline in the Paleogene may be due to bias caused by the lack of sufficient data.
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Paleobathymetry of Microborers
Fig. 9.16. Bathymetry of different Phanerozoic reef types inferred from microborer associations. The figure summarizes investigations of modern shallow-water coral reefs of the Bahamas and deep-water coral reefs of the North Atlantic, and case studies of ancient reefs from the Silurian of Gotland Island, the Permian of Tunisia, the Middle Triassic Muschelkalk of Germany and Late Triassic reefs in the Northern Calcareous Alps and Turkey, and Late Jurassic reefs in southwestern Germany. The bathymetric zones are characterized by (1) the appearance of borings assigned to cyanobacteria (Hyella, Mastigocoleus) versus those affiliated with green algae (Ostreobium) or red algae (Conchocelis), (2) the frequency of algal-related borings perpendicular to the surface of the substrate versus those oriented parallel to the substrate surface, and (3) the occurrence of borings attributed to obligate photoautotrophs (cyanobacteria, algae) versus those produced by obligate or facultative chemoheterotrophs (fungi). Each ichnocoenosis is named after key taxa. Not shown in the figure is the shallow eutrophic zone I corresponding to the supratidal zone. The shallow eutrophic zone II covers the intertidal zone and probably the uppermost part of the subtidal zone. The shallow eutrophic zone III corresponds to the well illuminated part of the subtidal zone (light intensity 10% of the surface light). The deep eutrophic zone comprises the less to poorly illuminated subtidal zone (light intensity to about 1% of the surface light). The dysphotic zone (light intensity to 0.01 to 0.001% of the surface light) is not yet clearly defined by microborers. Similarly, the microborer associations of the aphotic zone characterized by a dominance of bacteria, fungi as well as sponge and bryozoan borings need further investigation. The asterisks mark the presence of key taxa of the respective photic zone. The use of diagrams showing light intensity versus water depths may lead to a rough estimation of absolute water depths. But note that light intensity depends on many factors, such as latitude, turbidity of the water, or overshadowing of the bored substrate by densely growing organisms. Modified from Vogel et al. (1999).
pods became common in Early Cretaceous reefs. Late Cretaceous rudist reefs exhibit macroborer associations consisting of bivalves, worms, sponges and cirripedians (Johnson et al. 2002). Paleogene macroborer associations are similar to those of the Late Cretaceous. A significant change in the abundance and composition of bioerosion by macroborers took place during the Oligocene and continued in the Miocene. Late Tertiary reefs exhibit boring patterns similar to recent ones (Pleydell and Jones 1988; Perrin 2002).
9.3.3.2 Quantitative Changes in the Intensity of Macroboring in Coral Reefs Boring intensity can be measured by the semi-quantitative assessment of ichnospecies abundance using field and laboratory samples, X-ray techniques, serial thin sections and point-counting (Modern coral reefs: Klein et al. 1991; Risk et al. 1995; Perry 1998). Preliminary data provided by Perry and Bertling (2000) suggest that ancient reefs were variously affected by macroborers
Describing Microbialites, Encrusters, and Borings
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Box 9.5. Describing microbialites, encrustations and borings. Microbial carbonates Micrite: Inhomogeneous or homogeneous texture? Peloidal micrite? Differences in color? With or without undulation? Peloids: Within the micritic matrix? In interskeletal or intraskeletal protected cavities? Silt- or fine sand-sized grains? Peloids with dark center surrounded by a clear rim of euhedral calcite crystals? Clotted fabric: Frequency of micritic and peloidal micritic areas? Relation of the clotted fabrics to non-clotted fabric areas? Remains of tubular and filamentous cyanobacteria? Calcimicrobes: Tiny tubes with micritic walls, often bifurcated. Morphological type (see Pl. 53), frequency, irregular occurrence or in layers? Fabric: Non-laminated or laminated fabric (–> microbialites or stromatolites; see Fig. 9.1 and Box 9.1)? Non-laminated fabric: Irregular, sometimes dome-shaped structures consisting of mesoclots (–> thrombolites)? Domeshaped, sometimes bushy structures exhibiting micritic, peloidal and coarsely zoned microstructures (–> dendrolites)? Dome-shaped structures consisting predominantly of undifferentiated micrite (–> leiolites)? Laminated fabric: Macroscopic growth form (dome shaped, columns, high or low forms; fine- or coarse-grained laminae; agglutinated laminae)? Internal structures (individual laminae composed of micrite peloids or mesoclots)? Disruptions of the lamination by sediment, borings or biogenic encrustations? Geometrical arrangement of basic growth elements (see Logan et al. classification, Fig. 9.3)? Biogenic encrustations Occurrence of the encrusters: On or within substrates (epibenthic or endobenthic)? Location of the crusts: Upward growing crusts on a surface or pendant crusts in cavities? Thickness of the crust: Uniform or laterally decreasing or increasing? Interruptions of crustbuilding: Layers separated by sediment or cement? Bored surfaces of layers? Substrate: Lithified surface, fossil or other sedimentary grains? Carbonate cement? Boundary encruster/substrate? Growth forms: Thin or thick crusts, sheets, mounds, plates, semi-erect or erect forms)? Sequential patterns: Composition of crusts consisting of different encrusters? Relation between parts formed by specific encrusters? Thickness of the individual crust layers? Relation between encrusters and microbialites: Crusts embedded within microbialites? Distinct or vague boundaries? Encrusting organisms: Identity (encrusting organisms should be differentiated to the lowest possible taxonomic level)? Species diversity and percent cover of upper and undersides of host organisms (e.g. corals, stromatoporoids, shells)? Proportion of autotrophic to heterotrophic encrusters? Bioerosion and borings Microborers: SEM characterization. Which groups (bacteria, cyanobacteria, algae, fungi, foraminifera, bryozoans)? Relative frequency? Growth sequences? Macroborers: Thin-section characterization. Simple tubes or branched networks? Shape of the bore holes (cylindrical, dumb-shaped, amphora-like)? Mode of the exterior opening (single or multi-apertured openings; circular, elliptical or slit-like openings)? Assignment of the trace fossils to major groups: Sponges, worms, bivalves or cirripedians? Dominating group? Dimension of the borings: Lengths and diameters of the bore holes? Orientation of the borings: Vertical, oblique, horizontal, irregular, or fluctuating? Substrate type: Rock surfaces? Lithified surface, e.g. hardgrounds and disconformity surfaces? Fossils? Sedimentary grains? Relation borings and substrate: Depth of the penetration below the bored substrate? Borings only in specific fossils? Cross-cutting patterns? Association of boring organisms and the organisms of the host substrate: Repeated co-occurrence of particular borers and specific host fossils? Borings only in specific biofacies types (e.g. only in coral reef limestones)? Quantitative data: Percentage of macroborer groups within a defined area? Density of macroboring measured by pointcounting or image analysis considering macroborers versus the non-bored substrate areas?
depending on reef type, environmental setting and dominant macroborers. The abundance of borings in coral reefs expressed by the ratio of the point counter percentage of borings to the percentage of non-attacked substrate areas per cm2 is low and moderate in the Late Triassic (Northern Alps: 0.01 – dominating macroborers bivalves and sponges, and 0.09 – mainly worms; Nayband Forma-
tion, Iran: 0.14 – worms dominate over bivalves and cirripedians) and in the late Early Jurassic (Morocco: 0.08 and 0.12 – dominated by cirripedians), moderate in the Middle Jurassic (Iran, India, Chile: 0.12 – bivalves and worms equally important), and increases in the Late Jurassic, Oxfordian, and Kimmeridgian microbial-coral reefs (central and western Europe: 0.330.59 – dominated by bivalves and worms followed by
396
sponges). In contrast, low-energy carbonates of the deep-water turbid facies dominated by bivalves exhibit a low boring intensity of 0.03-0.06. A moderate to high abundance of macroborers is found in Late Cretaceous coral reefs (Coniacian, Bavaria: 0.18 - bivalves and sponges dominate over worms; Turonian, France: 0.44 - worms dominating over bivalves). Factors influencing these patterns are apparently the existence of clear or siliciclastic-influenced water and lagoonal or shallow/deep fore-reef settings. Microfacies analyses of reef limestones but also of platform carbonates should provide more quantitative data that can be used as measures of the changes in macroboring and bioerosional intensity during earth history. 9.3.4 Microborer Associations are Proxies for Paleo-Water Depths The bathymetric distribution of modern endolithic microborers depends predominantly on the intensity and spectral composition of light. The association of microborers is different in the eu-, dys- and aphotic zones. Microborers exhibit strong water depth controls; the association patterns therefore represent one of the best paleobathymetric criterium (Vogel et al. 1999). This approach is based on the • SEM study of the composition and distribution of modern shallow- to deep-marine microborers, • investigation of trace fossils of microborers from ancient environments that can be assigned to specific paleo-water depths according to paleontological and facies criteria, • morphological comparison of ancient and modern microborings, which surprisingly exhibit a high degree of correspondence, and • establishment of characteristic ichnocoenoses defined by the relative dominance of variously lightadapted groups (cyanobacteria, green algae, red algae, fungi) and the occurrence of the most common endoliths. These studies resulted in bathymetric subdivisions of ancient subtidal, slope and deeper basinal sequences. The subdivisions correspond to bathymetric zones recognized in modern environments (see summary in Glaub 1994) and identified in the Silurian (Glaub and Bundschuh 1997), Devonian (Vogel et al. 1987), Permian (Balog 1996) and Triassic (Schmidt 1990, 1992, 1993; Glaub and Schmidt 1994; Balog 1996), Jurassic and Cretaceous (Glaub 1994; Hofmann 1996), and Tertiary (Radtke 1991) The distribution patterns have a high potential for estimating the water depths of ancient reefs (Fig. 9.16).
Practical Advices
9.4 Practical Advice: How to Describe Microbialites and Stromatolites, Biogenic Encrustations and Borings? Box 9.5 summarizes the questions which should be kept in mind when studying the biological aspects of the formation and destruction of carbonate sediments. Basics: Limestones are biological sediments Microbial carbonates and stromatolites Aitken, J.D. (1977): Classification and environmental significance of cryptalgal limestones and dolomites. With illustrations from the Cambrian and Ordovician of southwestern Alberta. – J. Sed. Petrol., 14, 405-441 Atlas, R.M., Bartha, T. (1981): Microbial ecology: Fundamentals and application. – 569 pp., Reading (Adisson-Wessley) Burne, R.V., Moore, I.S. (1987): Microbialites: organosedimentary deposits of benthic microbial communities. – Palaios, 2, 241-254 Castanier, S., Le Metayer-Kevrel, G., Perthuisot, J.-P. (1999): Ca-carbonates precipitation and limestone genesis - the microbiologists point of view. – Sedimentary Geology, 126, 9-23 Chavetz, H.S. and Buczynski, C. (1992): Bacterially induced lithification of microbial mats. – Palaios, 7, 277-293 Characklis, W.G., Marshall, K.C. (eds., 1990): Biofilms. – 796 pp., New York (Wiley) Cohen, Y., Castenholz, R.W., Halvorson, H.D. (eds., 1984): Microbial mats: stromatolites. – MBL Lectures in Biology, 3, 498 pp., New York (Alan R. Cis) Gerdes, G., Krumbein, W.E. (1987): Biolaminated deposits. – Lecture Notes in Earth Sciences, 9, 1-183 Grey, K. (1989): Handbook for the study of stromatolites and associated structures. – Stromatolite Newsletter, 14, 82-171 Grotzinger, J.P., Knoll, A.H. (1999): Stromatolites in Precambrian carbonates: Evolutionary mileposts or environmental dipsticks. – Annual Review of Earth and Planetary Sciences, 27, 313-358 Kalkowsky, E. (1908): Oolith und Stromatolith im norddeutschen Buntsandstein. – Zeitschrift der deutschen geologischen Gesellschaft, 60, 68-121 Kennard, M.M., James, N.P. (1986): Thrombolites and stromatolites: two distinct types of microbial structures. – Palaios, 1, 492-503 Logan, B.W., Rezak, R., Ginsburg, R.N. (1964): Classification and environmental significance of algal stromatolites. – Journal of Geology, 72, 68-83 Monty, C. (ed., 1981): Phanerozoic stromatolites. Case histories. – 249 pp., Berlin (Springer) Paul, J., Peryt, T.M. (2000): Kalkowsky’s stromatolites revisited (Lower Triassic Buntsandstein, Harz Mountains, Germany). – Palaeogeography, Palaeoclimatology, Palaeoecology, 161, 435-458 Pratt, B.R. (1982): Stromatolitic framework of carbonate mudmounds. – J. Sed.Petrol., 52, 1203-1227 Reid, R.P., Visscher, P.T., Decho, A.W., Stolz, J.F., Bebout, B.M., Dupraz, C., Macintyre, O.G., Paerl, H.W., Pinkney, J.L., Prufert-Bebout, J.L., Steppe, T.F., DesMarals, D.J. (2000): The role of microbes in accretion, lamination and early lithification of modern marine stromatolites. – Na-
Basics: Microbes, Encrusters, Borers
ture, 406, 989-992 Riding, R.E. (ed., 1991): Calcareous algae and stromatolites. – 571 pp., Berlin (Springer) Riding, R.E. (1999): The term stromatolites: towards an essential defintion. – Lethaia, 32, 321-330 Riding, R.E. (2000): Microbial carbonates: the geological record of calcified algal mats and biofilms. – Sedimentology, 47, 179-214 Riding, R.E., Awramik, S.M. (eds., 2000): Microbial sediments. – 331 pp., Berlin (Springer) Schneider, J., Campion-Alsumard, T. (1999): Construction and destruction of carbonates by marine and freshwater bacteria. – European Journal of Phycology, 34, 417-426 Shapiro, R.S. (2000): A comment on the systematic confusion of thrombolites. – Palaios, 15, 166-169 Walter, M.R. (ed., 1976): Stromatolites. – Developments in Sedimentology, 20, 790 pp. Further reading: K038, K137 Biogenic encrustations Brett, C.A. (1988): Paleoecology and evolution of marine hard substrate communities: an overview. – Palaios, 3,374-378 Dupraz, C., Strasser, A. (2002): Nutritional models in coralmicrobialite reefs (Jurassic, Oxfordian, Switzerland): Evolution of trophic structure as a response to environmental change. – Palaios, 17, 449-471 Jackson, J.B.C. (1979): Morphological strategies of sessile animals. – In: Larwood, G., Rosen, B.R. (eds.): Biology and systematics of colonial organisms. – 499-556, London (Academic Press) Moussavian, E. (1992): On Cretaceous bioconstructions: composition and evolutionary trends of crust-building associations. – Facies, 26, 117-144 Rasser, M., Piller, W. (1997): Depth distribution of calcareous encrusting associations in northern Red Sea (Safaga, Egypt) and their geological significance. – Proceedings of the 8th International Coral Reef Symposium, 1, 743748 Schmid, D.U. (1996): Marine Mikrobolithe und Mikroinkrustierer aus dem Oberjura. – Profil, 9, 101-251 Taylor, P.D., Wilson, M.A. (2002): A new terminology for marine organisms inhabiting hard substrates. – Palaios, 17, 522-525 Walker, S.E., Miller, W. (1992): Organism-substrate relations. Toward a logical terminology. – Palaios, 7, 236-238 Further reading: K213 Bioerosion, micro- and macroborers Anagnostidis, K., Golubic, S., Komarek, J., Lhotsky, O. (eds., 1988): Cyanophyta (Cyanobacteria). Morphology, Taxonomy, Ecology. – Algological Studies, 50-53, 584 pp.
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Bertling, M. (1999): Late Jurassic reef bioerosion - the dawn of a new era. – Bulletin of the Geological Society of Denmark, 45, 173-176 Bromley, R.G. (1970): Borings of trace fossils and Entobia cretacea Portlock, as an example. – In: Crimes, T.P., Harper, J.C.C. (eds.): Trace fossils, 49-90 Bromley, R.G. (1990): Trace fossils: Biology and taphonomy. – Special Topics in Palaeontology, 3, 280 pp. Bromley, R.G. (1994): The palaeoecology of bioerosion. – In: Donovan, S.K. (ed.): The paleobiology of trace fossils. – 133-154, London (Belhaven) Glaub, I, (1994): Mikrobohrspuren in ausgewählten Ablagerungsräumen des europäischen Jura und der Unterkreide (Klassifikation und Palökologie). – Courier Forschungsinstitut Senckenberg, 174, 1-324 Golubic, S., Friedman, I., Schneider, J. (1981): The lithobiontic ecological niche, with special reference to microorganisms. – J. Sed. Petrol., 51, 475-478 Hallock, P. (1988): The role of nutrient availability in bioerosion. Consequences to carbonate buildups. – Palaeogeography, Palaeoclimatology, Palaeoecology, 63, 275-291 Hassan, M. (1997): Modification of carbonate substrate by bioerosion and bioaccretion on coral reefs of the Red Sea. – Ph.D. Thesis Kiel University, 126 pp. Hutchins, P.A. (1983): Biological destruction of coral reefs: a review. – Coral Reefs, 4, 239-252 Kobluk, D.R., James, N.P., Pemberton, S.G. (1978): Initial diversification of macroboring ichnofossils and exploitation of the macroboring niche in the lower Paleozoic. – Paleobiology, 4, 163-170 Kowalewski M., Dulai, A., Fürsich, F.T. (1998): A fossil record of holes: the Phanerozoic history of drilling predatory. – Geology, 26, 1001-1004 Perry, C.T., Bertling, M. (2000): Spatial and temporal patterns of macroboring within Mesozoic and Cenozoic coral reef systems. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform components and interactions. – Geol. Soc. London, Spec. Publ., 178, 33-50 Lynch, J.M., Hobbie, J.E. (1988): Micro-organisms in action: concepts and applications in microbial ecology. 2nd edition. – 363 pp., Oxford (Blackwell) Radtke, G., Hifmann, K., Golubic, S. (1997): A bibliographic overview of microscopic and macroscopic bioerosion. – Courier Forschungsinstitut Senckenberg, 201, 307-340 Vogel, K.(1993): Bioeroders in fossil reefs. – Facies, 28, 109-114 Vogel, K., Balog, S.-J., Bundschuh, M., Gektidis, M., Glaub, I., Kruschina, J., Radtke, G. (1999): Bathymetric studies in fossil reefs, with microendoliths as paleoecological indicators. – Profil, 16, 181-191 Further reading: K038
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Chronostratigraphic Time Scale
Fig. 9.17. Chronostratigraphic time scale for the Phanerozoic (Golonka and Kiessling 2002) used throughout this book.
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10 Fossils in Thin Section: It is Not That Difficult
Many students of carbonate rocks are bewildered and sometimes frustrated by the morphological and microstructural diversity of skeletal grains and are satisfied by distinguishing major fossil groups. The present chapter will hopefully demonstrate that more detailed identifications of thin-section fossils are not so difficult and that identifications at lower systematic levels can provide fresh insights into the environmental controls on depositional processes. However, keep in mind, that thin-section fossils present only a part of the biota present in the rock and that thin-section diversity is not equivalent to biotic diversity (see Sect. 6.2.1.4). The chapter consists of four parts. The first part discusses specific problems related to the study of fossils in thin sections. The second part documents the principal criteria used to distinguish fossil groups commonly or abundantly occurring in thin sections of limestones, their distribution in time and their significance with regard to carbonate production, the recognition of paleoenvironments and depositional settings. The fourth part deals with the potential of thin-sections in recognizing assemblage zones Because a good picture may be worth more than a thousand words, the last chapter includes a list of monographs and papers showing useful thin-section microphotographs. Many fossils are restricted to specific time intervals. Knowledge of distinct time ranges is necessary for estimating their value as guide fossils or in recognizing time-dependent facies types. Age and time ranges of the fossils discussed in this chapter and throughout the book are noted within the frame of the chronostratigraphic time scale shown in Fig. 9.17, see left page!
10.1 Specifics of Thin-Section Fossils 10.1.1 How to Determine Fossils in Thin Sections? Several well-illustrated books assist in a first identification of skeletal grains: Horowitz and Potter (1971), Majewske (1974), Lucas et al. (1976), Scholle (1978), and Adams and MacKenzie (1998).
Determining fossils in thin sections used in microfacies analysis may be difficult, because: • randomly cut sections of microfossils may exhibit very different shapes and outlines, resulting in an often confusing taxonomic situation of ‘thin-section fossils’ (Racki and Sobon-Podgorska 1992), • diagenetic alterations may produce ‘diagenetic taxa’ as, e.g. exemplified by calcimicrobes (e.g. Renalcis: Pratt 1984), or some foraminifera whose tubular structures are difficult to distinguish from micritized shells or agglutinated worm burrows, • the lack of modern counterparts of many extinct fossils may cause very different opinions varying from author to author regarding the systematic position and classification of fossils known only from thin sections. Despite their uncertain position, these ‘microproblematica’ are often excellent facies indicators and can be used as guide fossils, • very small-sized skeletal grains may present problems in attributing the grains to specific systematic groups as demonstrated by the investigation of silt-sized fragmented material from known organisms (Feray et al. 1962; Matthews 1966; Force 1969). In general, skeletal identity is lost when the particle size lies below 62 µm. Additional SEM studies significantly increase systematic assignment, particularly of mollusc shells. Some fossils (e.g. larger foraminifera) common in thin sections require oriented sections to follow defined planes. These sections may be rare in randomly cut microfacies samples, contrary to the opinion of many geologists who believe that thin sections yielding abundant microfossils must be useful in generic or specific determinations. The problem is demonstrated by an Early Permian fusulinid limestone from Oman (Fig. 10.1). The sample was studied in six over-sized thin sections cut in various directions. These thin sections exhibit more than 2000 sections through the spindle-shaped foraminifers. Only a very small number of the sections can be used in systematic assignments as shown by a thin section with 457 sections of fusulinids. 25.5% of the fossils represent broken shells and 10.5% of the samples are
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_10, © Springer-Verlag Berlin Heidelberg 2010
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Cutting Effects
Fig. 10.1. Highly fossiliferous fusulinid foraminiferal limestone. The sample yields very abundant shells, but only a few sections exhibit morphological criteria allowing a taxonomic assignment, which is imperative if the age determination of the sample has priority. Early Permian: Baid area, eastern Oman. Scale is 1 cm.
cut extremely obliquely and are therefore not suitable for taxonomic determinations. 64% of the sections can be attributed to specific cutting planes, but most of these sections run only parallel or oblique to the coiling axis of the shells. These sections exhibit morphological criteria which are of taxonomic value only in combination with axial and sagittal sections showing the arrangement of the whorls around the initial part of the shell (Pl. 68). Sections allowing a generic determination (Parafusulina) are extremely rare: 4.5% of all sections present in the thin section. What can we learn from this example? Be aware of which problems should be solved with your sample. If you are aiming at indications of paleoenvironment, a sample exhibiting the abundance and orientation of the foraminiferal tests will be sufficient. But if you want to know the age of the sample, you should send a sufficiently large sample that allows oriented thin sections to be prepared by your paleontology colleague, specialized in the study of larger foraminifera (look for lists in the net), and ask for friendly cooperation.
Several attempts have been made to reduce the geometric diversity of organisms with hardparts and the wealth of figures seen in thin sections.
Cutting effects Obviously, the reconstruction of the three-dimensional geometry of fossils from two-dimensional cross sections can be hazardous if only a few sections are examined (Fig. 10.2). Microstructural features are of value in discriminating fragments of mollusks and echinoderms, but other groups (e.g. foraminifera, calcareous algae) require oriented sections along defined planes.
10.1.2 Which Fossils in which Time Interval?
The approach used by Horowitz and Potter (1971) concentrates on the basic geometry and distinguishes five major groups: (1) Straight or coiled single cones, tubes or cylinders; (2) colonial tubes; (3) valves; (4) multi-element or multi-plated shells, and (5) reticulate or rectangular and cystose networks. The identification key developed by Flügel (1982) considers the two-dimensional thin-section figures and distinguishes six groups: (1) Circular and elliptical figures, (2) U and V shaped figures, (3) curved figures, (4) net-like figures, (5) layered structures, and (6) chains of segments. Although both authors underline the pitfalls of these groupings, beginners and people eager to become acquainted with the mysteries of biotic identifications in thin sections, will find these ‘cookery books’ useful.
Most thin-section fossils have wide stratigraphic ranges, but different stages are characterized by the association and abundance of common and dominating fossil groups found in limestones. Fig. 10.3 and Box 10.1 give a rough picture of the larger groups which can be expected in limestones of different age and different depositional settings.
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A
B
Fig. 10.2. Cutting effects. Random cutting of microfossils results in very different figures. Relatively few of these figures show all the morphological criteria necessary for a determination at low taxonomic levels. Generally a combination of several figures and oriented sections are needed for generic or specific determinations of thin-section fossils. In contrast, random cuts can often be assigned to major fossil groups, if additional criteria (e.g. size, microstructure) are considered. A: Different sections through the cylindrical calcareous skeleton of the Late Paleozoic dasyclad green alga Velebitella (see Pl. 59/1) produce various figures. The diagnostic taxonomic criteria must be derived from longitudinal, oblique, tangential and cross sections. Modified from Kochansky-Devidé (1964). B: The Late Paleozoic smaller foraminiferal genus Climacammina is characterized by a calcareous test whose chambers are biserially arranged in an early stage and uniserially in the last growth stage. The aperture is cribrate. The figure displays oriented sections which in combination exhibit these criteria and allow the generic determination of the microfossil. Modified from Cummings (1958).
10.1.3 Practical Advice The following text summarizes some points which should be considered looking at fossils in thin sections. Browsing among thin-section photographs in books and microfacies monographs (the penultimate point) is certainly very tempting, but often misleading. • Use various techniques (polarized light, CL staining etc.). • Check the orientation of the thin section with re-
spect to bedding planes and geopetal structures (vertical, parallel, oblique). • Consider the size of the grain. It is imperative to know the field of view of the microscope at each magnification. • Evaluate the presence of major fossil groups (e.g. bivalves, echinoderms, calcareous algae) and decide which group should be investigated in detail (e.g. calcareous algae), because they give indications of paleoenvironmental constraints and the age of the sediment.
Fig. 10.3. Generalized distribution of the major fossil groups occurring in thin sections of Phanerozoic limestones.
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Preservation • Test the preservation of the fossils. • Consider the possibility of recrystallization, diagenetic solution or replacement of the original microstructure. • Are the fossils complete or fragmented? • Consider the degree of fragmentation and the resulting sizes. • Are the fossils in place or transported? • In place fossils may need several sections in order to recognize the variability of internal structures. • In-place fossils, particularly colonial reef organisms, are often large and their study requires oversized thin sections or peels (up to 10 and 20 cm in length). • In-place fossils often have characteristic growth forms which will influence the figures seen in thin sections. Quality of the thin sections • Compare the thin sections with the rock sample, and if necessary, make additional and better oriented thin sections. • Differentiate between random sections and oriented sections. • Some fossil groups, particularly sessile benthos (e.g. corals, coralline and stromatoporoid sponges) need at least two sections (parallel and perpendicular to the growth direction). • Indicate the orientation of the section using defined terms (e.g. vertical, tangential, oblique). • Look especially for sections that contain defined morphological criteria necessary for taxonomic determinations (e.g. axial structures in larger foraminifera). • Decide which of the taxonomically important criteria are shown by the sections and try to combine these
Occurrence ot Thin-Section Fossils
criteria in reconstructing the shape and internal structure of the fossils. Skeletal mineralogy and microstructure • Look at the mineralogical composition of the fossils (calcareous, siliceous, phosphatic, organic). • Decide whether the original mineralogy of the skeletal grain was aragonitic or calcitic. Preservation of the bioclast as a mold or cast indicates a primary aragonitic mineralogy that has been subsequently altered. Well-preserved microstructures point to a primary calcitic mineralogy. • For shells try to discriminate the original microstructures and structures produced by diagenesis. • Describe the microstructure by looking at the shape, size and orientation of the component crystals. Advancing systematic determinations • Bear in mind the age of the rock examined, and the likely biota will be encountered. • Browse among the thin-section microphotographs in the microfacies monographs (see Sect. 10.3). • If you think you have found a reliable assignment to a specific group or on taxa, look for further information in the relevant parts of Sect. 10.2.
10.2 Diagnostic Criteria of Fossils in Thin Sections Thin-section fossils comprise six informal groups: • cyanobacteria (calcimicrobes) and calcareous algae, • foraminifera and other protozoa (radiolarians, calpionellids),
Box 10.1. Fossil groups common in thin sections of marine limestones formed in different settings. Ranking indicates the approximate frequency. Note that this list is highly subjective and does not consider specific associations, e.g. in reef limestones and slope sediments, as well as microfossils of uncertain systematic position present predominantly in inner and outer platform environments and reefs. Late Tertiary Inner shallow platform: Smaller benthic foraminifera, corallinacean red algae, mollusks, bryozoans, barnacles, serpulids, ostracods Edge: Reefs with corallinacean red algae, corals, bryozoans Slope: Benthic and planktonic foraminifera Basin: Planktonic foraminifera Early Tertiary Inner shallow platform: Smaller and larger benthic foraminifera, calcareous dasyclad algae and udoteacean green algae, corallinacean and squamariacean red algae
Edge: Larger foraminifera Slope: Planktonic and benthic foraminifera Basin: Planktonic foraminifera Upper Late Cretaceous Inner shallow platform: Benthic foraminifera, mollusks, bryozoans, echinoderms Edge: Rudists Slope: Planktonic and benthic foraminifera, ‘calcispheres’, siliceous sponges Basin: Planktonic foraminifera, ‘calcispheres’ Continued next page!
Occurrence ot Thin-Section Fossils
Box 10.1. Continued Lower Late Cretaceous Inner shallow platform: Benthic foraminifera, gastropods, rudists and other bivalves, Edge: Rudists Outer platform: Benthic foraminifera, echinoderms, cephalopods Basin: Planktonic foraminifera, ‘calcispheres’, cephalopods Early Cretaceous Inner shallow platform: Benthic foraminifera, dasyclad and udoteacean green algae, gastropods Edge: Corals, calcimicrobes, benthic foraminifera Outer platform: Benthic foraminifera, sponges, echinoderms, bivalves, cephalopods Basin: Planktonic foraminifera, cephalopods Late Jurassic Inner shallow platform: Benthic foraminifera and dasyclad green algae, solenoporacean red algae, calcimicrobes, gastropods, serpulids Edge: Corals and ‘hydrozoans’, siliceous sponges, brachiopods, bivalves, echinoderms Slope: Pelagic bivalves, echinoderms, benthic foraminifera, ‘calcispheres’, cephalopods Basin: Calpionellid tintinnids, radiolarians, pelagic crinoids, pelagic bivalves, ‘calcispheres’; siliceous sponges, bivalves, cephalopods Early Jurassic Inner shallow platform: Benthic foraminifera, dasyclad green algae, calcimicrobes, gastropods Edge: Benthic foraminifera Outer platform: Benthic foraminifera, bivalves Basin: Radiolarians, sponge spicula, cephalopods Late Triassic Inner shallow platform: Benthic foraminifera, dasyclad green algae, calcimicrobes, bivalves Edge: Reefs with scleractinian corals and coralline sponges, foraminifera, solenoporacean red algae Slope: Echinoderms, bivalves, brachiopods Basin: Thin-shelled pelagic bivalves, radiolarians, cephalopods Middle Triassic Inner shallow platform: Dasyclad green algae, calcimicrobes, gastropods, benthic foraminifera Edge: Reefs with coralline calcareous sponges, scleractinian corals, smaller foraminifera Slope: Reefs with encrusting microfossils (e.g. Tubiphytes), coralline calcareous sponges, brachiopods, echinoderms Basin: Radiolarians, thin-shelled pelagic bivalves Early Triassic Inner shallow platform: Gastropods, bivalves, small benthic foraminifera Edge: Mollusks, echinoderms Basin: Bivalves, cephalopods
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Late Permian Inner shallow platform: Dasyclad and gymnocodiacean green algae, phylloid algae, fusulinid foraminifera, gastropods Edge: Reefs with calcareous coralline sponges, bryozoans, calcimicrobes, rugose corals Slope: Reefs with encrusting microfossils (e.g.Tubiphytes), brachiopods, echinoderms Basin: Radiolarians Early Permian Inner shallow platform: Ancestral red algae, dasyclad green algae, phylloid algae, fusulinid foraminifera, smaller forminifera, bryozoans, gastropods Edge: Reefs with microbes, phylloid and dasyclad green algae Outer platform: Brachiopods, echinoderms Basin: Radiolarians Late Carboniferous Inner shallow platform: Ancestral red algae, dasyclad green algae, smaller foraminifera, fusulinid foraminifera, gastropods Edge: Bryozoans, coralline sponges, echinoderms, brachiopods Slope: Echinoderms, bivalves Basin: Radiolarians, echinoderms, bivalves Early Carboniferous Inner shallow platform: Smaller benthic foraminifera, dasyclad green algae, various calcareous red algae, ‘calcispheres’ Edge: Reefs with calcimicrobes, bryozoans Outer platform: Smaller benthic foraminifera, various red algae, mollusks, echinoderms, rugose corals Basin: Radiolarians Devonian Inner shallow platform: Stromatoporoid sponges, tabulate and rugose corals, ‘calcispheres’, dasyclad and udoteacean green algae, benthic foraminifera Edge: Reefs with stromatoporoids and tabulate corals, calcimicrobes, rugose corals Slope: Echinoderms, brachiopods Basin: Tentaculitids, cephalopods Silurian Inner shallow platform: Dayscladacean green algae, calcimicrobes, mollusks Edge: Reefs with stromatoporoid sponges, tabulate and rugose corals, bryozoans, echinoderms Outer platform: Brachiopods, trilobites, echinoderms Basin: Tentaculitids, cephalopods Ordovician Inner shallow platform: Green and red algae, mollusks Edge: Reefs with calcimicrobes, sponges and bryozoans Outer platform: Trilobites, bryozoans, brachiopods, echinoderms, sponges Basin: Cephalopods Cambrian Inner shallow platform: Calcimicrobes, mollusks Edge: Reefs with archaeocyath sponges and calcimicrobes Outer platform: Trilobites, sponges, echinoderms Basin: Echinoderms.
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• benthic sessile organisms (sponges, corals, bryozoans), • shell-bearing organisms with valves (brachiopods, mollusks, serpulids, ostracods) and multi-element/ multi-plate shells (trilobites, crustaceans, echinoderms), • rare thin-section fossils (including vertebrates, conodonts, graptolites, and wood), • fossils of uncertain systematic position (microproblematica). These groups and their major systematic units are presented by instructive plates and with accompanying text. The text starts with a definition of the group, followed by the diagnostically important morphological criteria, takes a short glance at the classification and discusses the environment and the biofacies, and finally gives an overview of the temporal distribution. The significance of the group as well as hints for avoiding confusions with other fossil groups are summarized in the last part.
10.2.1 Cyanobacteria and Calcareous Algae Definition: Algae are aquatic photosynthetic benthic and planktonic plants, ranging from micron-sized unicellular forms to giant kelps several meters long. These eucaryotic organisms display extremely varied lifecycles, characterized by different complexes of photosynthetic pigments, which together with other characters are used to distinguish major groups, i.e. red, green and brown algae (Rhodophyta, Chlorophyta, Phaeophyta). In contrast to these groups, the so-called ‘bluegreen algae’ (Cyanophyta) are procaryotes, lacking a discrete nucleus and membrane-bound organelles within the cell. They are physiologically more closely related to bacteria and are included within a separate division of the Procaryota (‘Cyanobacteria’). In comparison with eucaryotic algae, the cyanobacteria are adapted to a wider range of environmental conditions (which may include the loss of photosynthetic activity) and occur both in freshwater and marine environments. Because of these differences, the paleoenvironmental interpretation of ‘algae’ has to consider the systematic membership of the fossils as well as the differences in the calcification potential. Calcification: Calcareous algae constitute a highly artificial group comprising various benthic and planktonic algae belonging to different systematic groups, which are capable of removing carbon dioxide from the water in which they live and secreting or depositing carbonate around the algal body (thallus). Calcifi-
Cyanobacteria and Calcareous Algae
cation has been interpreted as a light-shading function, protection against predators and epiphytes, or as means of receiving higher amounts of light and nutrients. The mode and objective of calcification are topics of ongoing research (Borowitzka 1989; Pentecost 1991). The precipitation of calcium carbonate may occur within the cell (planktonic coccolithophorids), in association with cell walls (coralline red algae), outside the cell (many green algae) or as a surface deposit on the thallus (many blue-greens). The skeleton of calcified green and red algae displays distinct ultrastructures (Flajs 1977). Cyanobacterial calcification results from complex, in places, bacterially-initiated processes within and upon the mucilaginous sheath surrounding the cell wall. Binding and trapping of carbonate at the sheat surface enhance the formation of structures capable of fossilization (calcimicrobes). In all groups, calcification is associated with organic matter, primarily polysaccharides, which may stimulate nucleation and control subsequent growth of calcium carbonate crystals (Merz 1992). Less than 10% of the total of modern benthic algal species and cyanobacteria exhibit calcification, but the percentage of calcareous algae may increase to about 20% in warm shallow-marine environments. In the cyanobacteria, considerably more calcified species are known from freshwater than from marine environments, in contrast to green algae, with a reversed situation. Calcification of red algae is almost completely restricted to the oceans. Disregarding the diatoms (a group of the Chromophyta) whose cell walls are impregnated with silica, modern skeletal algae exhibit skeletons consisting of either aragonite or calcite (see Box 4.9). Aragonite is the exclusive mineral of marine green algae, but also occurs in some red algae (e.g. squamariaceans). The skeletons of corallinacean red algae consist of HighMg calcite containing magnesium amounts ranging up to 30 mole percent. Low-Mg calcite is the common mineral of nonmarine cyanobacteria as well as of another group of the Chromophyta, the marine planktonic coccolithophorids. The calcified reproductive organs of freshwater charophyceans consist of Low-Mg or HighMg calcite. The first well-calcified cyanobacteria are known from the Late Precambrian, the first calcified algae from the lowermost Cambrian. Since the calcification of cyanobacterians depends strongly on environmental conditions favoring precipitation of calcium carbonate, the abrupt appearance of calcified cyanobacteria near the Precambrian-Cambrian boundary has been interpreted as possibly reflecting a change in seawater chemistry (decrease of the Mg2+/Ca2+ ratio: Riding 1982). Environmental controls on algae: Controls on the
Distribution of Cyanobacteria and Algae
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distribution of recent benthic algae comprise (a) physical factors: intensity and quality of light (necessary for photosynthesis), water temperature, water energy and substrate; (b) chemical factors: water chemistry, salinity, dissolved gases, pH values; and (c) biological factors: nutrients, competing and herbivorous organisms, and ecological successions. Generally, the intensity of light decreases with increasing water depth. About 50 % of the light penetrating clear water is absorbed in the uppermost ten meters. Despite this, photosynthetic algae can occur at greater water depths if they have special pigments that can utilize minimal light intensities and specific wave lengths. Modern red algae, adapted to blue and green light sectors, also occur at greater water depths, whereas calcified green algae assimilate most strongly in the red sector and are, therefore, preferentially limited to shallow seas. Preservation of calcareous algae: Because of the widely differing solubilities of calcium carbonate phases, the primary mineralogy of calcareous algae is a critical factor in their diagenesis and in the preservation of algal fossils. Aragonitic skeletons of udoteacean and dasyclad green algae are susceptible to solution and replacement by sparry calcite. Coralline red algae display differing degrees of diagenetic alterations, de-
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pending on varying magnesium contents. Low-Mg calcite algae and calcimicrobes are commonly well-preserved. Classification: Calcified cyanobacteria have been classified into formal taxonomic units (Chuvashov 1987; Dragastan et al. 1996) or into informal morphological groups (Riding 1991). The latter classification is more practicable (Fig. 10.4). Living algae are classified according to pigments, nature of the food reserve, kind of flagellation, and chloroplast ultrastructure. The presence of skeletal carbonate is not a classification criterium. Fossil calcareous algae are classified according to morphological characteristics used in classification schemes of analogous living forms. These characters include external morphology, preserved tissue, and reproductive bodies. Many fossil algae, however, can not be related to major groups of living algae. These algae are included in non-systematic collective groups (e.g. ancestral red algae, phylloid algae) or in taxonomic units which can only be tentatively assigned to major groups of green or red algae (e.g. gymnocodiacean algae). Still other ancient extant algae exhibit morphological characters that are similar to green or red algae, but do not allow an attribution to known systematic subgroups (e.g. ung-
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Cyanobacteria Corallinacean red algae Squamariacean red algae Ancestral red algae Solenoporacean red algae Udoteacean green algae Gymnocodiacean green algae Phylloid algae Dasyclad green algae
Fig. 10.4. Distribution of Phanerozoic calcimicrobes and major calcareous algal groups. Note that the thickness of the bars does not reflect taxonomic diversity, but instead indicates the relative frequency in thin sections.
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darellacean algae). Fig. 10.4 exhibits the temporal distribution of major groups of calcareous algae. Further smaller groups are dealt within the discussion of red and green algae. Significance of algae in microfacies analysis Fossil calcified algae are most important in microfacies analysis, because of their high potential in reconstructing paleoenvironments and ecological conditions, their biostratigraphic potential, and their large contribution to the formation of carbonate sediments. Paleoenvironment: Fossil calcified algae are used in reconstructing environmental factors such as water temperature (see Pl. 105), salinity and relative water depths (see Pl. 108). This approach has to consider several points: (a) Modern calcareous green and red algae exhibit different ecological controls with regard to water depths. Some algal groups occupy wide bathymetric ranges: Coralline red algae and udoteacean green algae are not restricted to ‘shallow’ depths but also thrive at water depths of several tens of meters, corallinaceans even in depths of more than two hundred meters. Paleobathymetric interpretations should be based on specific algal groups rather than on the generalized statement of the occurrence of ‘algae’. (b) A few algal groups exhibit differences in preferred modern and ancient biotopes. Ancient calcimicrobes are common fossils in marine limestones, whereas modern calcified cyanobacteria flourish in waters with nonmarine salinities. Modern dasyclads prefer lagoonal and shelf habitats in contrast to ancient forms common in lagoonal and shelf carbonates, but also occur in reef limestones. The major factors controlling the distribution of algae may be modified by local factors, e.g. substrate conditions (Flügel 1983). (c) Not all modern calcified algae are highly lightdependent. Green algae and red algae exist in shadowy habitats. Ancient calcimicrobes are common constituents of cryptic biotopes. (d) Mean annual water temperature has a strong control on the distribution of calcareous algae, but for modern corallinacean red algae this factor is distinctly taxadependent. (e) Algal bioclasts are easily transported and redeposited. They may not represent autochthonous constituents of the sediment, but rather grains transported into other carbonate settings. Downslope and turbidity current transport of green algal debris from modern carbonate platforms to deep water environments is a common feature. Transport of shallow-marine algae to
Algae and Sediment
deeper subtidal, slope and basinal environments as well as the growth of various red and some green algae at water depths of several tens and even a few hundred meters is also responsible for the possible occurrence of algal material in deep-marine settings. Storms can transport algal debris from open-marine platforms to near-coastal environments. Even a transport of marine algae to nonmarine settings by storms is likely (Pl. 131/4). To avoid some of these pitfalls, the use of distribution patterns is recommended. This approach relates the composition of algal associations to other microfacies data (limestone textures and structures, additional fossils). The algal associations are characterized with regard to their taxonomic composition, frequency and dominance of specific taxa or groups, and diversities. The distribution patterns of algal associations can not be directly compared with the ecological zonation of living algae, because many ancient taxa are extinct and their relationship to living forms may be unknown. Nevertheless, established empirical distribution patterns have a high potential for reconstructing paleoecological factors, e.g. paleoclimate and paleoceanography and paleoenvironments. Stratigraphic zonation: Calcareous algae are valuable in subdividing rock units and establishing zonations. These zonations are based on the occurrence of characteristic associations of species, or the range of single species, or the abundance and climax of species. Commonly used marine algal groups are dasyclad green algae, some red algae, and charophycean algae in nonmarine sediments. Sedimentation: Calcareous algae are major producers of carbonate sediments in shallow marine environments as well as in deep-water settings. Modern sediment-producing algal groups in shallow platform and inner ramp environments comprise predominantly udoteacean green algae (e.g. Halimeda; see Sect. 4.1.2) and corallinacean and squamariacean red algae. The contribution of dasyclad algal skeletal grains to recent sediment production is minor as compared with ancient carbonate platforms, particularly of Late Paleozoic and Mesozoic age. Udoteacean green algae were important producers of carbonate mud and sand at least since the Late Triassic, but probably already in the Paleozoic. Common Late Paleozoic sediment producers were ancestral red algae forming rhodoids, and phylloid algae contributing to the accumulation of sediment in reef mounds. The deposition of calcitic planktonic algae (coccolithophorid chrysophyceans, calciodinoflagellate dinophyceans, and other related groups) in pelagic deep-
Basics: Calcimicrobes and Algae
shelf environments and open-marine basinal settings started during the Triassic, increased in the Jurassic and became a dominant process in Cretaceous and Cenozoic times. The influx of shallow-water algal material derived from platforms into slope and basinal settings is known already from the Cambrian, and is a common feature in Mesozoic periplatform carbonates. Cyanobacteria, charophycean algae as well as a few green algae are important constituents of nonmarine Mesozoic and Cenozoic lacustrine carbonates. Basics: Calcimicrobes and algae Far more general texts and overviews deal with calcareous algae than most invertebrate groups. Taxonomically based overviews are listed at the end of the relevant paragraphs. General texts Barattolo, F., de Castro, P., Parente, M. (1993): Studies on fossil benthic algae. – Bolletino della Società Paleontologica Italiana, Spec. Vol., 1, 420 pp. Bassoullet, J.P., Bernier, P., Conrad, M.A., Deloffre, R., Jaffrezo, M. (1978): Les algues dasycladales du Jurassique et du Crétacé. – Geobios, Mém. spec., 2, 330 pp. Chuvashov, B.I., Shuysky, V.P., Schaikin, I.M.(1987): Iskopaemye izvestkovye vodorosli. – Akademya Nauk SSSR, Sibirskoe otdelenie, Trudy instituta geologii i geofiziki Novosibirsk, Vypusk, 674, 200 pp. Flajs, G. (1977): Die Ultrastrukturen des Kalkalgenskelettes. – Palaeontographica, B, 160, 69-128, Stuttgart Flügel, E. (ed., 1977): Fossil algae. Recent results and developments. – 375 pp., Berlin (Springer) Ginsburg, R., Rezak, R., Wray, J.L. (1972): Geology of calcareous algae (Notes for a short course). – Sedimenta, I, 70 pp., Miami (Comparative Sedimentology Laboratory, Rosenstiel School of Marine and Atmospheric Science) Höfling, R., Moussavian, E., Piller, W.E. (eds., 1993): Facial development of algae-bearing carbonate sequences in the Eastern Alps. – International Symposium and Field-Meeting ‘Alpine Algae ’93’, Guidebook, 206 pp., MünchenWien Johnson, J.H. (1961): Limestone-building algae and algal limestones. – 295 pp., 139 Pls., Golden (Colorado School of Mines) Maslov, V.N. (1956): Iskopaemye izvestkovye vodorosli SSSR. – Akademya nauk SSSR, Trudy instituta geologicheskikh Nauk, Vypusk, 297 pp. Monty, Cl. (ed., 1981): Phanerozoic stromatolites. Case histories. – 249 pp., Berlin (Springer) Pentecost, A. (1991): Calcification processes in algae and Cyanobacteria. – Riding, R. (ed.): Calcareous algae and stromatolites. – 3-21, Berlin (Springer) Pentecost, A., Riding, R. (eds., 1978): Biomineralization in lower plants and animals. – The Systematic Association, Spec. Vol., 30, 400 pp., London Pia, J. (1926): Pflanzen als Gesteinsbildner. – 355 pp., Berlin (Borntraeger) Piller, W. (ed., 1994): Proceedings of the International Symposium and Field-Meeting ‘Alpine Algae ’93’. – Beiträge zur Paläontologie, 19, 264 pp., Wien Poignant, A.F., Deloffre, R. (eds., 1979): 2nd International
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Symposium on Fossil Algae. – Bull. Centres Rech. Explor.Prod. Elf-Aquitaine, 3, 385-885, Pau Riding, R. (ed., 1991): Calcareous algae and stromatolites. – 571 pp., Berlin (Springer) Round, F.E. (1981): The ecology of algae. – 653 pp., Cambridge (Univ. Press) Toomey, D.F., Nitecki, M.H.(eds., 1985): Paleoalgology. Contemporary research and applications. – 375 pp., Berlin (Springer) Walter, M.R. (ed., 1976): Stromatolites. – Developments in Sedimentology, 20, 790 pp., Amsterdam (Elsevier) Wray, J.L. (1977): Calcareous algae. – Developments in Palaeontology and Stratigraphy, 4, 184 pp., Amsterdam (Elsevier) Reviews of calcareous algae from different time intervals Awramik,S.M. (1991): Archaean and Proterozoic stromatolites. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 289-304, Berlin (Springer) Chuvashov, B., Riding, R. (1984): Principal floras of Palaeozoic marine calcareous algae. – Palaeontology, 27, 487500 Emberger, J. (1976): Les algues (Chlorophyceae, Prasinophyceae, Rhodophyceae) du Carbonifère et du Permien. Essai d’un inventaire bibliographique, géographique, stratigraphique. – Bulletin de l’Institut de Géologie du Bassin d’Aquitaine, Numero Spécial, 168 pp. Emberger, J. (1976): Les algues (Chlorophyceae, Prasinophyceae, Rhodophyceae) du Dévonien. Essai d’un inventaire bibliographique, géographique, stratigraphique. – Bulletin de l’Institut de Géologie du Bassin d’Aquitaine, Numero Spécial, 94 pp. Emberger, J. (1979): Les algues (Euchlorophyceae, Prasinophyceae, Rhodophyceae)du Trias. Essai d’un inventaire bibliographique, géographique, stratigraphique. – Bulletin de l’Institut de Géologie du Bassin d’Aquitaine, Numero Spécial, 157 pp. Flügel, E. (1991): Triassic and Jurassic marine calcareous algae: a critical review. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 481-503, Berlin (Springer) Johnson, J.H. (1961): A review of Ordovician algae. – Quarterly, Colorado School of Mines, 56/2, 1-101 Johnson, J.H. (1961):The Jurassic algae. – Quarterly, Colorado School Mines, 59, 129 pp. Johnson, J.H. (1963): Pennsylvanian and Permian algae – Quarterly, Colorado School of Mines, 58, 1-218 Johnson, R. (1966): A review of the Cambrian algae. – Quarterly, Colorado School of Mines, 61, 1-162 Johnson, J.H. (1969): A review of the Lower Cretaceous algae. – Professional Contributions Colorado School of Mines, 6, 180 pp. Johnson, J.H., Konishi, K. (1956): A review of Mississippian algae. – Quarterly, Colorado of School Mines, 51/4, 1-84 Johnson. J.H., Konishi, K. (1958): A review of Devonian algae. – Quarterly, Colorado School of Mines, 53/2, 1-84 Johnson, J.H., Konishi, K. (1959): A review of Silurian (Gotlandian) calcareous algae. – Quarterly, Colorado School of Mines, 54/1, 1-114 Mamet, B. (1991): Carboniferous calcareous algae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 370451, Berlin (Springer) Riding, R. (1991): Cambrian calcareous Cyanobacteria and algae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 305-334, Berlin (Springer)
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Riding, R., Guo, L. (1991): Permian marine calcareous algae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 452-480, Berlin (Springer) Roux, A. (1985): Introduction a l’étude des algues fossiles Paléozoiques (de la bacteria a la tectonique des plaques). – Bull. Centre Rech. Explora.-Prod. Elf-Aquitaine, 9, 465-699 Roux, A. (1991): Ordovician algae and global tectonics. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 335348, Berlin (Springer) Roux, A. (1991): Ordovician to Devonian marine calcareous algae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 349-368, Berlin (Springer) Further reading: K129 to K136
10.2.1.1 Cyanobacteria and Calcimicrobes Definition, calcification and morphology: Cyanobacteria are phototroph Procaryota. The cells are arranged in colonies (coccoid forms; Croococcales) or arranged within threads forming trichomes (filamentous forms; Hormogonophyceae). The threads can be branched and enclosed in a protecting mucilaginous sheath. The development of organic sheaths is strongly environmentally controlled. Some modern filamentous and coccoid cyanobacteria calcify through impregnation within or encrustation upon the extracellular sheaths with carbonate. Encrusted filaments are potential fossils, whereas impregnated filaments commonly disintegrate to carbonate mud. The mineralogy reflects water chemistry (freshwater: Low-Mg calcite; marine and restricted environments: Aragonite and High-Mg calcite). Calcification is species-specific. Calcification in recent cyanobacteria corresponds to a nucleation of calcium carbonate upon or within the mucilage sheath (Pentecost and Riding 1986; Pentecost 1990). Considerably more cyanobacteria have been observed in modern freshwater lakes and streams compared with the sea. However, calcification occurred widely in marine cyanobacteria during the Paleozoic and Mesozoic, perhaps reflecting temporal changes in oceanic carbonate chemistry. Merz-Preiss and Riding (1999) comparing calcification patterns of cyanobacteria from the Everglades in Florida and from spring tuffs stress the point that the preservation potential of fossilized filamentous cyanobacteria probably depends on the degree of carbonate oversaturation of the ocean water. Calcification of filaments produces tubelike microfossils, varying in external diameter between ten and several tens of microns. The tubes are bordered by microcrystalline calcitic ‘walls’. Calcified cyanobacteria occur as millimeterto centimeter-sized discrete microfossils (calcimicrobes) consisting of tubes or irregular bodies, or as amalgamated masses exhibiting laminated or massive growth forms. Laminated growth forms include skel-
Cyanobacteria and Calcimicrobes
etal stromatolites and oncoids, sometimes exhibiting discrete tubular microstructures. Other growth forms are microbushes, thrombolites and dendrolites. Classification: Cyanobacteria are extremely variable with regard to their morphology. Some fossil calcimicrobes exhibit high morphological similarities with groups of modern cyanobacteria, indicating the existence of very conservative long-range characters. The tubular microstructure was used to define the artificial group ‘Porostromata’. Some authors still prefer the term ‘porostromate algae’ but one should be aware that this group includes cyanobacterian fossils with close morphological similarity to recent calcified blue-greens, but also various green algae. A handy as well as morphological classification of calcimicrobes proposed by Riding (1991) is shown in Fig. 10.5. Another approach, based on comparative studies of porostromate taxa and recent cyanobacteria, results in a classification that relates ancient taxa to recent counterparts (Dragastan 1985, 1988, 1989, 1990; Dragastan et al. 1996). Environmental constraints: Modern cyanobacteria have a wide range of habitats. As pioneer settlers in marine freshwater and on terrestrial sedimentary deposits, benthic cyanobacteria prepare the ground for the successive invasion and expansion of eukaryotic flora and fauna (Golubic et al. 2000). Benthic cyanobacteria are the major constituents of microbial mats, specifically of the topmost oxic layers. The mats form in inter- and supratidal, sometimes hypersaline coastal environments as well as in soda and freshwater lakes or hot springs. Mat-building cyanobacteria exhibit significant tolerance to changes in light intensity, anoxic conditions or sporadic drying. Effective environmental factors are light, water temperature and nutrients (C, N, P, Fe). Planktonic cyanobacteria are common in freshwater but also occur in open oceans. Ancient benthic cyanobacteria are known both from marine and nonmarine environments. Calcimicrobes are common constituents of Mesozoic restricted, lagoonal as well as on open-marine platform carbonates and contributes to the formation of reefs and mud mounds. Distribution: Nonskeletal cyanobacteria were abundant in the Precambrian and have been recorded from rocks as old as 3.5 billion years. Heavily calcified calcimicrobes appeared in the Vendian (1.0–0.85 billion years) coincident with the decline in stromatolites. Vendian reefs are composed of abundant thrombolites and calcimicrobes. The absence of Proterozoic calcified cyanobacteria Text continued on p. 412
Classification: Calcified Cyanobacteria
409 Angulocellularia Group: Precambrian to JurassIc Criteria: Micritic, solid, bush-like form, erect or pendant growth; similar to recent Oscillatoriaceae Examples: Angulocellularia: Cambrian, Early Ordovician; Frutexites: Paleozoic, Jurassic Epiphyton Group: Cambrian and Early Ordovician; Late Devonian Criteria: Micritic narrow dendritic filaments, which may be rod-like, apparently chambered or tubular; erect or pendant growth Examples: Epiphyton: Cambrian to Devonian; Tharama: Silurian to Devonian Girvanella Group: Abundant in the Paleozoic Criteria: Thin-walled tubes, tangled, coiled or aligned; prostrate growth; similar to recent Oscillatoriaceae Examples: Girvanella: Cambrian to Cretaceous Hedstroemia Group: Cambrian to Cretaceous Criteria: Branching clusters of tubes; clusters with fan-like longitudinal sections, radial erect growth. Adjacent tubes may share common walls or may be separate; taxa separated according to patterns of branching. Similar to recent rivulariacean cyanobacteria Examples: Hedstroemia: Silurian; Ortonella: Carboniferous; Cayeuxia: Triassic to Cretaceous Renalcis Group: Abundant in Cambrian and Early Ordovician; Late Devonian; common in reefs Criteria: Clusters of hollow reniform bodies, usually with thick walls which may be micritic, fibrous or clotted; irregular or erect growth Examples: Renalcis: Cambrian to Early Carboniferous; Shuguria: Cambrian to Carboniferous Garwoodia/Mitcheldeania Group: Late Paleozoic and Mesozoic Criteria: Coarse, thin-walled, radiating clusters of tubes which are commonly juxtaposed (systematic position uncertain, may be green algae) Examples: Garwoodia: Devonian to Cretaceous; Mitcheldeania: Carboniferous to Cretaceous Rothpletzella Group: Silurian and Devonian Criteria: Flat, curved and encrusting sheets of juxtaposed tubes which branch dichotomously in one plane, systematic position uncertain Examples: Rothpletzella: Devonian
Fig. 10.5. Practical classification of calcified cyanobacteria based on growth forms and filament patterns appearing as variously oriented tubiform structures or irregular bodies (modified from Riding 1991). The tubes occur in clusters within commonly millimeter-sized nodules, bushes or sheets. Preferred growth direction is erect, pendant or prostrate (horizontal, parallel to the substrate). The Hedstroemia group and the Garwoodia/Mitcheldeania group comprise microfossils, often described as porostromate algae or microbes. The systematic position of these fossils (cyanobacteria or green algae) as well as the generic differentiation is still a matter of discussion. Thin-section pictures: Frutexites occurring in a deep-water stromatolite. Note the bush-like growth form and the microstructure consisting of radially arranged fibres of Fe-Mn oxides and calcite. Frutexites was probably non-phototrophic and seems to have preferred oxygen deficient, low-energy environments (Böhm and Brachert 1993). Early Jurassic: Near Lunz, Austria. - Epiphyton: The genus appears close to the Precambrian-Cambrian boundary and is widespread in Cambrian reef limestones. Early Cambrian: Southern Sardinia. - Girvanella forms small nodules, oncoids or sheets composed of loosely or densely tangled thin-walled tubes. Early Ordovician: Western Newfoundland. - Cayeuxia and morphologically similar taxa are abundant fossils in Mesozoic platform carbonates. The genus was placed in the synonymy of Rivularia, a modern cyanobacterium occurring in freshwater and marine environments. Late Jurassic (Tithonian): Sulzfluh, Switzerland. - Renalcis often occurs together with Epiphyton (bottom of the picture). Both genera are principal constituents of Cambrian archaeocyath reefs. Early Cambrian: Southern Sardinia. - Garwoodia and other taxa of this group are characterized by coarser tubes than of Hedstroemia and similar genera. The size of the tubes has been used as an argument for assigning these microfossils to udoteacean green algae. Late Jurassic (Tithonian): Sulzfluh, Switzerland. - Rothpletzella occurs in platform and reef carbonates. Early Devonian (Currajong limestone): Canberra, Australia.
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Calcified Cyanobacteria
Plate 53 Calcified Cyanobacteria: Porostromate Calcimicrobes Cyanobacteria (also known as cyanophytes or blue-green algae) are procaryotes. They contribute to sediment production (microbial mats), sediment construction (e.g. travertines and tufa; reefs), and sediment destruction (boring and micritization). Cyanobacterial microboring is the oldest known form of bioerosion. The first calcimicrobes appeared near the Precambrian-Cambrian boundary. Calcified cyanobacteria were abundant in Paleozoic and Mesozoic shallow-marine carbonates. Late Mesozoic and Early Tertiary calcimicrobes are commonly found in lacustrine or marginal-marine limestones. Calcified cyanobacteria (calcimicrobes) form millimeter- to centimeter-sized bushy masses, radial fans or tangled masses, often composed of very small ‘tubes’ (designed by calcification within organic sheaths), or irregularly shaped bodies. Calcimicrobes comprise microfossils consisting of tubes organized in erect radial clusters of branched tubes (e.g. Cayeuxia: –> 1, 2), prostrate felts composed of non-oriented intermingled tubes (e.g. Girvanella: –> 7), dendritic micritic filaments, apparently chambered (e.g. Epiphyton; Fig. 10.5)), or aggregates of hollow reniform bodies (e.g. Renalcis; Fig. 10.5). Most recent Cyanobacteria use Low-Mg calcite, but a few also High-Mg calcite or aragonite. Fossil calcimicrobes consist of Low-Mg calcite. At present, intense cyanobacteria calcification appears to be essentially a freshwater phenomenon and is rare in modern subtidal environments, in contrast to ancient cyanobacteria which occupied tidal and subtidal environments and were common in reef and platform settings. Common criteria used in the differentiation of groups or taxa of tubiform porostromate forms are the arrangement of tubes (e.g. erect, pendant, prostrate; –> 1, 7, 8), branching patterns (e.g. simply dichotomous, polydichotomous, V-shaped or right-angle branching; –> 1, 3, 6), and the shape and dimensions of the tubes (inner diameter, wall thickness and wall structure). Growth forms (hemispherical, –> 1; tufted, –> 4; nodular, – > 6; laterally expanded crusts and sheets, –> 7, 8) as well as growth zonation patterns (–> 2, 4) reflect environmental and biological constraints. Do not confuse tubular porostromate calcimicrobes with other thin-section fossils! Solenoporacean filaments shown in longitudinal sections have smaller dimensions and often exhibit horizontal partitions (Pl. 55). Bryozoan tubes are conspicuously larger and morphologically differentiated (Pl. 85). 1 Cayeuxia Frollo. Longitudinal section of a nodular thallus. The small nodules exhibiting a hemispherical outline are composed of outward radiating tubes that originate from a point located peripherally on the thallus. The tubes diverge upward at a narrow angle. Cayeuxia is very similar to the recent alga Rivularia; some authors consider the genera as being identical (Dragastan 1985). Late Jurassic (Tithonian): Kapfelberg, Southern Franconian Alb, Germany. 2 Cayeuxia Frollo. SEM sample. Note growth zonation (GZ) and branching pattern (arrows). Same locality and age as –> 1. 3 Cayeuxia. Note nodular growth form, long branched tubes (arrow) and the difference in tube diameter as compared with –> 1. Species of Cayeuxia (and Rivularia) are differentiated according to divergence angle, tube diameter and branching pattern. Late Jurassic (Tithonian): Sulzfluh Mountains, Graubünden, Switzerland. 4 Porostromate cyanobacteria (Pycnoporidium lobatum Yabe and Toyama) are characterized by erect growth and differentiation into marginal tufts (arrow). Late Jurassic: Lower Saxonia, Germany. 5 Ortonella Garwood. Cross-section (CS) and longitudinal section (LS) of a globular growth form. High-energy grainstones of an inner carbonate ramp. Early Carboniferous (Kohlenkalk, Viséan): Aachen, Germany. 6 Garwoodia Wood. Note difference in thickness of the tubes (as compared with –> 1 and 5) and right-angle branching pattern (arrow). Some authors place this genus into the green algae. Garwoodia and morphologically similar taxa are known from Paleozoic and Mesozoic platform carbonates. Jurassic and Cretaceous species are abundant in back-reef and inner shelf environments. Early Cretaceous: Subsurface, Hohenlinden, Bavaria, Germany. 7 Girvanella ducii Wethered, characterized by thin-walled tangled, coiled or aligned tubes and prostrate growth. The genus is common in Paleozoic carbonates. Filaments of micrite-walled calcimicrobes as Girvanella might have been a major source of lime mud in the Early Paleozoic when the tubes fell apart (Pratt 2001). Middle Devonian (Givetian reef): Sauerland, Germany. 8 Rothpletzella Wood, characterized by flat, curved or encrusting sheets of tubes branching dichotomously in one plane and appearing as ‘strings of pearls’ in oblique sections. Rothpletzella is a common Devonian calcimicrobe occurring both in shallow and deep-water settings, and characteristically found at reef-margins and on fore-reef slopes. The genus acted as constructors, but also as crypts within cavities. Early Devonian: Near Canberra, Australia.
Plate 53: Calcified Cyanobacteria
411
Coralline Algae
412
and their appearance in the Late Precambrian has been explained by a decrease in the Mg2+/Ca2+ ratio near the Precambrian-Cambrian boundary (Riding 1982), extreme calcium carbonate oversaturation of Precambrian ocean waters (Knoll et al. 1993), and biological evolutionary controls (Defarge et al. 1994). Calcified filamentous cyanobacteria are common in marine sediments from the Early Cambrian to the Late Cretaceous. Calcimicrobes flourished in Cambrian and Ordovician reefs and open-marine shelf carbonates; they were less common in Silurian to Permian carbonates, prominent again in lagoonal and reef environments during Middle Triassic to Mid-Cretaceous time, but scarce in the Late Cretaceous. In reefs calcimicrobes acted as frame-builders (e.g. Cambrian and Devonian Renalcis; Pl. 82/4. 5) as well as crypts within cavities (e.g. Frutexites; Fig. 10.5). Porostromate calcimicrobes are common constituents of Early Carboniferous ramp deposits (Pl. 53/ 5) and Mesozoic inner platform and lagoonal carbonates (Pl. 53/1-4). Depth distribution was not necessarily limited to very shallow environments, as shown by the occurrence of calcimicrobes in Devonian deeper forereef positions. Starting with the Tertiary, the group became primarily limited to nonmarine and transitional environments (Monty 1973; Pentecost and Riding 1986; Riding 1992). Records of calcified cyanobacteria from Early Cenozoic and modern shallow-marine carbonate environments are rare, probably due to the major change in the preferred biotopes of blue-greens during the Mesozoic. Starting with the Late Jurassic and Early Cretaceous, calcified cyanobacteria adapted to marginal-marine and freshwater conditions, where they become conspicuous carbonate producers in lacustrine and terrestrial environments. Basics: Calcified cyanobacteria Dragastan, O., Golubic, S., Richter, D.K. (1996): Rivularia haematites: a case of recent versus fossil morphology. Taxonomical considerations. – Revista Española de Micropaleontologia, 28, 43-73 Dragastan, O., Richter, D.K., Gielisch, H., Kube, B. (1998): Environmental siginificance of some Mesozoic ‘Porostromata’ calcareous algae. – Revista Española de Micropaleontologia, 30, 59-101 Elicki, O. (1999): Palaeoecological significance of calcimicrobial communities during ramp evolution: an example from the Lower Cambrian of Germany. – Facies, 41, 27-40 Fournie, D. (1967): Les Porostromata du Paléozoique. Étude bibliographique. – Bull. Centre Rech. Pau-SNPA, 1, 21-45 Merz-Preiss, M. (2000): Calcification in Cyanobacteria. – In: Riding, R., Awramik, S.M. (eds.): Microbial sediments. – 50-56, Berlin (Springer) Pia, J. (1927): Thallophyta. – In: Hirmer, M. (ed.): Handbuch der Paläobotanik. – 1, 31-136, München (Oldenbourg) Pratt, B.R. (1984): Epiphyton and Renalcis – diagenetic micro-
fossils from calcification of coccoid blue-green algae. – J. Sed. Petrol., 54, 948-971 Riding, R. (1991): Calcified cyanobacteria. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 55-87, Berlin (Springer) Further reading: K033, K129-K136
10.2.1.2 Corallinacean and Peyssoneliacean Red Algae Sections 10.2.1.2 to 10.2.1.4 deal with red algae. Calcified red algae are common fossils in thin sections of Phanerozoic limestones (Fig. 10.6). They comprise some well-defined systematic units (Corallinaceae, Peyssoneliaceae, Solenoporaceae), Paleozoic fossils interpreted as ‘ancestral corallines’ (Archaeolithophyllaceae), and microfossils attributed to red algae but of unknown systematic position (Ungdarellaceae, Stacheinaceae and other Paleozoic groups). Coralline algae Definition and calcification: Corallinacean algae (also known as coralline algae) are a distinct, exclusively marine group within the Rhodophyta. They represent the most heavily calcified red algae. Calcification occurs intracellularly by High-Mg calcite with varying amounts of MgCO3 (up to > 30 vol%). The stabilization of the High-Mg calcite connected with a magnesium loss is interpreted as a congruent dissolution of High-Mg calcite and precipitation of Low-Mg calcite with a definite textural change (Oti and Müller 1985). Classification: Traditionally coralline algae have been divided according to growth morphology into two major groups – the non-geniculate (non-articulated) crustose corallines and the geniculate (articulated) corallines (Fig. 10.7). Non-geniculate encrusting corallines comprise crusts, a few microns to several centimeter thick, on rocks, coral skeletons, shells, other algae or sea grass. Encrusting corallines vary in size from thin crusts to massive, nodular or lamellate thalli several tens of centimeters in diameter. Some non-geniculate corallines are free-living and occur as rhodoids. Geniculate corallines are erect branching, tree-like plants attached to the substrate by crustose or calcified, root-like holdfasts. The plants are flexible owing to non-calcified sections (genicula) separating longer calcified segments. Individual segments vary in shape from cylindrical to flattened. The segments consist of a central part (medulla) extending upwards to the next
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Calcareous Red Algae Thin-section criteria Corallinaceans Fine net-work structure; cell size < 5 to 15 µm; differentiation according to cell size and arrangement; case-shaped cavities in lines or isolated within the tissue; crusts, nodules, branches; in transmitted light dark. . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pl. 54
Stratigraphic range Early Cretaceous to recent, possibly already Early Paleozoic
Peyssoneliaceans Net-work structure; cell size 20 - 30 µm; basal calcification with botryoids; distinctly set sheets; blades, nodules; in transmitted light golden, yellowish, gray . . . . . . . . . .Pl. 64/10
Early Cretaceous to recent possibly also Late Paleozoic
Solenoporaceans Closely packed vertical filaments, sometimes with horizontal partitions; cell size 10 - 100 µm; no cell differentiation, no case-shaped cavities; nodular and branched . .Pl. 55
Cambrian to Tertiary (Miocene)
Ancestral red algae (Archaeolithophyllaceans) Net-work structure; cell size approx. 10 - 30 µm; cell differentiation; blades and nodules; in transmitted light yellowish, gray . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pl. 56/2, 8
Carboniferous and Permian (Viséan to Late Permian)
Problematic algae Ungdarellaceans Net-work structure; cell size about 20 µm; cell differentiation;very small encrusting or isolated structures; in transmitted light often golden, yellowish . . Pl. 56/5-7, Pl. 108/7-8
Carboniferous and Permian Common: Viséan–Moscovian
Stacheinaceans Indistinct network; very small encrusting structures; in transmitted light often golden, yellowish . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . Pl. 108/10
Carboniferous (Viséan to Moscovian)
Fig. 10.6. General thin-section criteria of some common calcareous red algae.
segment and composed of partitioned filaments, and peripheral cellular filaments (cortex) surrounding the medulla (Fig. 10.7G; Pl. 64/11). At present, the morphology and taxonomy of coralline genera are under revision. The identification of species should be reserved for specialists. Both, traditional morphological criteria and neontological features are used in determining coralline taxa. Identification keys have been published by Braga et al. (1993) and Rasser and Piller (1999). Commonly used features for identi-
fying fossil genera are growth form, development of cells within different parts of the thallus, and reproductive organs. Morphology: The thallus consists of adjacent calcified filaments. Consecutive cells within core filaments are connected by primary pit connections or by cell fusions. The calcified cellular tissue is characterized by definite arrangements, shapes and sizes of cells. Cells are very small (5–20 µm).
Fig. 10.7. Internal morphology of calcified coralline algae. Both encrusting and erect articulated corallines exhibit differentiation of the plant tissue into a basal (central) and a peripheral (upper) part differing in the geometry, arrangement and size of the filaments. A: Lithothamnion crust overgrowing an older algal nodule. B to E: Basic types of hypothallus (H) and perithallus (P) organization in encrusting corallines. F: Typical growth form of erect, articulated corallines (Corallina). G: Internal organization of the calcified segments. A: After Fott (1971), B-G: After Wray (1977).
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Most encrusting corallines (Pl. 54) are dorsoventrally organized. They have a basal-most (dorsal) part (formerly called the hypothallus) near the substrate, with cells oriented more or less parallel to the substrate, responsible for the lateral growth, and an outer (ventral) part (formerly called the perithallus) with commonly very small cells, developed above and upon the hypothallus and oriented perpendicular to it, responsible for the plant’s thickness. The basal part may con-
Coralline Red Algae
sist of multilayered or unilayered filaments (Fig. 10.7B– E) The cells of the multilayered basal part are described with regard to size, shape, arrangement of cells into single, multilayered and arcuate layers (coaxial; Pl. 54/ 8; Fig. 10.7 C) or they are non-coaxial (plumose). A thin surface layer (epithallus) known from modern corallines is very rarely preserved in fossil algae. Additional criteria revealed by SEM and ultrathin section studies (e.g. connections between contiguous cells)
Plate 54 Corallinacean Red Algae Tertiary limestones composed predominantly of red algae are important hydrocarbon reservoir rocks, e.g. in Libya and Indonesia. Corallinacean algae (or coralline algae) are a red algal group characterized by a strongly calcified tissue exhibiting a fine net-like structure in thin sections. The group is traditionally subdivided into encrusting (or crustose) corallines and segmented articulate corallines. The former develop millimeter- to centimeter-thick crusts, the latter erect, flexible branches consisting of calcified segments and non-calcified parts. Corallinacean algae are abundant constituents of Cretaceous and Cenozoic shelf and reef carbonates. The plate shows Tertiary and Cretaceous encrusting corallines forming rigid skeletons. The algal body (thallus) consists of adjacent calcified filaments. Corallinacean algae are most valuable paleoecological markers. They are used in assessing paleoclimate, paleolatitudes and paleodepths. But note that the occurrence of corallinacean red algae is not a priori indicative of a shallow tropical warm-water environment; coralline algae grow in temperate and cold-water environments as well. In contrast to calcareous green algae, coralline algae flourish down to several decades of meters even down to about 250 m. The study of corallines in thin sections needs high magnifications, otherwise algal fragments appear nonstructured and may be confused with micrite clasts. Similar confusion may exist for the red alga Marinella. This alga (Pl. 136/3, 5), abundant in Late Jurassic and Cretaceous (Barattolo and DelRe (1985) and known until the Oligocene, is characterized by millimeter-sized thalli composed of extremely fine threads (about < 5 to 15 µm). Corallines appear dark in transmitted light and white in hand specimens and polished sections. 1 Red algal nodule (rhodoid) consisting of Lithothamnion, a member of the melobesioid coralline algae. Large spar-filled structures occurring within the thallus are reproductive organs (conceptacles, C). These structures should not be confused with borings (B). The genus is characterized by encrusting growth, a multi-layered internal structure and multiporate conceptacles (black arrows). Interspaces between algal specimens are filled with sediment (peloids, foraminifera). Angular algal fragments (white arrows) are caused by vadose dissolution and breakage of the algal limestone. The growth form of the nodules and peloidal infilling indicate restricted water energy within a reef environment. Late Tertiary (Leithakalk, Miocene): St. Margarethen, Burgenland, Austria. 2 Bioclastic algal debris facies with broken branches of Lithophyllum. Note the different cell size of the hypo- and perithallus. The genus is an important framebuilding coralline alga in Cenozoic reefs. The sample comes from sediments deposited adjacent to small patch reefs. Late Tertiary (Leithakalk, Miocene): Oslip near Rust, Burgenland, Austria. 3 Sporolithon (formerly called Archaeolithothamnium). The genus is characterized by a multilayered hypothallus and a thick perithallial tissue with individual sporangia arranged in rows (‘sori’; arrow). Late Tertiary (Miocene): Großhöflein west of Eisenstadt, Burgenland, Austria. 4 Mesophyllum. Arrows point to conceptacles with a single-pored roof (–> 6). Tertiary (Miocene): Southwestern Turkey. 5 Corallina: Articulated coralline alga characterized by erect, jointed thalli consisting of segments. The segments consist of straight central (medullary, M) equant cells and a cortex (C) composed of tiny cells. Quartz-rich algal rudstone. Note the detrital subangular quartz grains (white). Transgressive near-shore shelf facies (Reijers and Petters 1997). Mid-Cretaceous (Mfamosing Limestone, Middle Albian): Calabar Flank, Niger Delta, Southern Nigeria. 6 Lithophyllum. Conspicuous single-pored conceptacle (arrow). Temperate near-coastal facies. Tertiary (Oligocene): Pohlkotte, Lower Saxony, Germany. 7 Sporolithon. Conceptacles are borne in ‘sori’ (S; compare Pl. 64/12). Same locality as Fig. 6. 8 Rhodoid composed of crustose branches of Lithophyllum. Note differences in cell sizes in the basal part (hypothallus, HT) and the perithallus (PT). Late Tertiary (Miocene): Turkey. 9 Articulated coralline alga. Subterranophyllum. Tertiary (Late Oligocene/Early Miocene): Drowned reef complex in the Sulu Sea south of Palawan, Philippines. –> 1, 2: Dullo 1983; 5: Oti and Koch 1990; 6, 7: Flügel 1970; 9: Wiedicke 1987
Plate 54: Corallinacean Red Algae
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have become increasingly more important (Johansen 1981; Woelkerling 1988; Bosence 1991). Reproductive organs are borne in calcified case-like chambers (conceptacles). Conceptacles may be uniporate (Pl. 54/4), multiporate (Pl. 54/1), or arranged in rows (sori; Pl. 54/7). The type, the kind of conceptacle roof formation, and the location of conceptacles are diagnostic criteria. Environmental constraints: Modern corallines range from tropical to polar oceans and extend from intertidal areas to depths of about 250 m. Tropical taxa occur in depths down to about 80 m, cold-water taxa between about 20 and 250 m (Pl. 4/1). The distribution of corallines is controlled by environmental factors, predation (grazing) and competition. Competitive interaction among crustose corallines is influenced by the thickness of the crusts and the height to which the growing margin is raised (Steneck et al. 1991). Most corallines prefer hard rough substrates, dim light conditions and agitated water. They occur in warm and cold waters from low to high latitudes and are commonly adapted to normal marine salinities. Growth rates depend mainly on light conditions. Many corallines are shade-adapted. Distribution: The definite record of abundant corallines begins near the start of the Cretaceous, perhaps already in the latest Jurassic. The recognition of Ordovician and Silurian fossils with coralline-like features extends the earliest record to the Early Paleozoic (Brooke and Riding 1998; Riding et al. 1998). Sediment production: Coralline algae are one of the most abundant marine carbonate producers. Unattached corallines live on hard and soft sea bottoms and can form rhodoids (Sect. 4.2.4.2; Pl. 12/6, Pl. 149/6) that contribute to rhodolith pavements and maerl (composed of branched rhodoids and coralline branches; Pl. 4/2; Sect. 2.4.4.3). Attached corallines form algal ridges (late successional stage of coral reef development forming a constructional cap over corals as the reef approaches sea level) and algal cup reefs (intertidal cupshaped reefs formed by intergrowing coralline algae and invertebrates), trottoirs (intertidal buildups, usually growing on steep rock shores, e.g. Mediterranean and the Northern Atlantic) and the coralligène de plateau (occurring in deeper shelf waters of about 20 to 160 m, and passing laterally into carbonate sands and gravels or terrigenous sands). Crustose corallines are important Cenozoic reef builders in tropical and subtropical regions, as framebuilders and binders, and as cementing agents for bio- and lithoclasts (Pl. 149/2).
Significance of Corallinacean Algae
Coralline bioclasts contribute significantly to the formation of marine warm-water, temperate-water, coldwater and even polar sand- and gravel-sized carbonates (Pl. 33/2, Pl. 35/1). Reef frameworks are constructed from crustose or branching coralline algae both in high- and low-energy environments, but differ in the density of the crusts and branches. Frameworks cemented by fibrous aragonite and microcrystalline High-Mg calcite occur in highenergy coralline algal ridges. Low-energy frameworks are characterized by intraskeletal cements precipitated only during the life of the algae. Bioerosion by microand macroborers is strong in high-energy reefs and less common in low-energy reefs. Sediments derived from high-energy reefs comprise coarse particles from breakage by waves and fine particles caused by borers. Quietwater frameworks are broken by wave action and bioturbation; wackestones and packstones are deposited around the reefs (Bosence 1985). Significance of corallinacean algae: Most Tertiary genera occur even today. Because of these long-term ranges, the study of present-day coralline algae offers an excellent potential for paleoenvironmental interpretations. Of particular value are thallus shape and generic associations. The morphology of coralline crusts and branches reflect different water energies (Bosence 1983). Rhodoids are sensitive to turbulence in shallowmarine environments (Bosence and Pedley 1982; Sect. 4.2.4.2). Because modern crustose coralline genera have existed from the Miocene onward, the paleobathymetry can be successfully evaluated at least for the Neogene (Basso 1998). Basics: Corallinacean red algae Adey, W.H., MacIntyre, I.G. (1973): Crustose coralline algae: a re-evaluation in the geological sciences. – Bull. Geol. Soc. Amer., 84, 883-904 Basso, D. (1998): Deep rhodolith distribution in the Pontian Islands, Italy: a model for the paleoecology of a temperate sea. – Palaeogeogr., Palaeoclimat., Palaeoecol., 137, 173187 Bosence, D.W. (1983): Coralline algal reef frameworks. – J. Geol. Soc. London, 140, 365-376 Bosence, D.W. (1991): Coralline algae: mineralization, taxonomy, and palaeoecology. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 98-113, Berlin (Springer) Bosence, D.W., Pedley, H.M. (1982): Sedimentology and palaeoecology of a Miocene coralline algal biostrome from the Maltese Island. – Palaeogeogr., Palaeoclimat., Palaeoecol., 39, 9-43 Braga, J.C., Bosence, D.W., Steneck, T.S. (1993): New anatomical characters in fossil coralline algae and their taxonomic implications. – Palaeontology, 36, 535-547 Johansen, H.W. (1981): Coralline algae, a first synthesis. – 239 pp., Boca Raton, Florida (CRC)
Peyssoneliacean Red Algae
Nebelsick, J.H., Bassi, D. (2000): Diversity, growth forms and taphonomy: key factors controlling the fabric of coralline algae dominated shelf carbonates. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems: compounds and interactions. – Geol. Soc. London, Spec. Publ., 178, 89-107 Woelkerling, W.J. (1988): The coralline red algae: an analysis of the genera and subfamilies of non-geniculate Corallinaceae. – 268 pp., British Museum Natural History, London (Oxford Univ. Press) Look at http://www.botany.uwc.ac.za/clines/bibliog/Intro.htm for the bibliography of Rhodophyta (prepared by W.J. Woelkerling with over 4500 references), Coralline News, or Rasser: http://www.paleoweb.net/rasser/welcome-main.htm.
Peyssoneliacean red algae Definition, calcification and morphology: Peyssoneliaceans (formerly called squamariaceans) are a small group of marine calcareous red algae, some of which secrete an aragonitic skeleton. Calcification is both intracellular and extracellular. The latter occurs as a layer of small botryoids below the thallus projecting downwards from basal hypothallus cells. The botryoids are similar to those forming aragonite cements (Sect. 7.4.2.1). The internal structure resembles some crustose corallinaceans. It consists of rectangular hypothallus cells (about 30 µm high and 20 µm wide) from which vertical or arcuate rows of smaller perithallus cells arise. The algae grow as arched millimeter- to centimetersized sheets on soft mud substrates and hard rock surfaces, form a bridge-like network between corals in reefs, and occur as concentric layers within nodules (Fig. 9.7, Fig. 10.8). Growth forms are layered and foliose. The former consist of superimposed layers with little or no space between, similar to the growth of encrusting corallines. The latter is an open structure composed of irregularly attached arcuate crusts forming overlapping sheets with open cavities in-between. In thin sections the crusts of fossil peyssoneliaceans differ from corallinaceans by a transparent golden or yellow-brownish tint in contrast to the uniform dark appearance typical of coralline algae in general. Environmental constraints: Calcified peyssoneliacean algae occur in temperate, and tropical and subtropical waters in depths between a few meters to more than 100 meters. They are common down to 50 meters. Distribution is controlled by light intensity, current regime and sediment input. The algae are often associated with coralline red algae and encrusting foraminifera.
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Fig. 10.8. Peyssoneliacean red algae. A: Peyssonelia. The thallus is built of thin blades. B: Internal structure of Peyssonelia composed of a thin hypothallium with basal aragonite calcification between downward projecting tubules, and a thick perithallium. A: After Pia (1926). B: After Wray (1964).
Distribution: The confirmed fossil record starts in the Early Cretaceous. Peyssoneliaceans are widely distributed in Cretaceous and Tertiary warm-water and temperate shelf carbonates. The algae are common in Paleocene and Eocene reef limestones associated with corals, or limited to rhodoid accumulations near reefs. Important genera are: Ethelia (also called Polystrata), Peyssonelia and Pseudolithothamnium. Significance: Striking similarities between calcified peyssoneliacean algae and some Late Carboniferous and Permian phylloid algae suggest that the latter may be closely related, thus allowing a less speculative interpretation of economically important phylloid algal buildups (James et al. 1988). Basics: Peyssoneliacean red algae Buchbinder, B., Halley, R.B. (1983): Occurrence and preservation of Eocene squamariacean and coralline rhodoliths: Eua. Tonga. – In: Toomey. D.F., Nitecki. M.H. (eds.): Paleoalgology, contemporary research and application. – 248-256, Berlin (Springer) Denizot, M. (1968): Les algues Floridées envroutantes (à l’exclusion des Corallinacées). – Thése, Laboratoire de Cryptogamie, Museum Nationalè d’Histoire naturelle Paris, 280 pp. James, P., Wray, J.L., Ginsburg, R.N. (1988): Calcification of encrusting aragonitic algae (Peyssoneliaceae): implications for the origin of late Palaeozoic reefs and cements. – J. Sed. Petrol., 58, 291-303 Massieux, M., Denizot, M. (1964): Rapprochement de genre Pseudolithothamnium Pfender avec le genre actuel Ethelia Weber van Bosse (Algues Florideae, Squamariaceae). – Revue de Micropaléontologie, 7, 31-42 Moussavian, E. (1988): Die Peyssoneliaceen (auct.: Squamariaceae; Rhodophyceae) der Kreide und des Paläogen der Ostalpen. – Mitteilungen der Bayerischen Staatssammlung für Paläontologie und historische Geologie, 28, 89124
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10.2.1.3
Solenoporacean Red Algae
Solenoporacean Red Algae
Definition, calcification and morphology: Solenoporaceans are extinct encrusting calcareous algae that developed nodular and branched growth habits ranging in size from several millimeters to a few centimeters. The thalli consist of closely packed threads of vertically divergent calcified cells, which appear circular or polygonal in cross-sections. Solenoporaceans are commonly preserved as Low-Mg calcite, but rare records and the strong and patchy recrystallization indicate a primary aragonitic composition. In contrast to coralline algae no distinct differentiation of the cellular tissue exists. Undisputable reproductive structures are not known. Cell diameter and length is commonly larger than those of corallinaceans (about 10 to 100 µm). Taxonomic differentiation is based on zonation patterns, the arrangement, shape and size of cells, partitions within the filaments and walls and cell connections. The presence/absence and regular or irregular spacing of horizontal partitions within the filaments have been
used as generic characteristics. The traditional species differentiation only on the basis of cell dimensions is highly artificial. The group has to be restudied. Vachard et al. (1989) drew attention to the possibility that many solenoporacean ‘species’ correspond to diagenetic stages rather than to well-founded taxonomic units. Systematic position: The Solenoporaceae are now regarded as a subgroup of the red algal group Florideae because of the existence of pit-like connections between the cells (Moussavian 1989). The group has been considered to represent possible ancestral forms of the Corallinaceae, but still distinct evidence for this interpretation is missing. Distribution: Cambrian to Miocene. Solenoporaceans are constituents of Paleozoic and Mesozoic subtidal open-marine shelf carbonates and reef limestones. In some Silurian, Permian and Late Cretaceous reefs the algae are concentrated near the open-marine reef margin. They are common in Middle and Late Trias-
Plate 55 Solenoporacean Red Algae Solenoporacean algae, known from the Cambrian to the Late Tertiary, are common thin-section fossils in Paleozoic and Mesozoic platform and reef carbonates. The encrusting calcareous algae had nodular and branched growth habits and range in size from several millimeters to a few centimeters. Solenoporaceans have been regarded as possible ancestors of corallinacean red algae. Longitudinal sections show radially or vertically divergent calcified filaments appearing as long ‘tubes’ (–> 3) or superposed cells within the filaments (–> 5). Cell diameter (about 10-100 µm) is larger than that of most corallinacean red algae. Taxonomic differentiation is based on zonation patterns, the arrangement, shape and size of cells (–> 3, 5, 8), partitions (–> 4, 5) within the filaments and walls and cell connections. The group has to be restudied regarding additional morphological criteria (–> 8). 1 Solenporacean floatstone. The nodular algal thalli are bored and marginally micritized. Solenoporacean algae, often associated with porostromate algae, are common constituents of small-scaled Frasnian and Famennian buildups. Latest Devonian (Strunian): Dubic-Debnik near Krakow, southern Poland. 2 Solenopora. Nodular thalli in life positon with distinct growth zonation. They occur together with dendroid corals (not shown) and form a reef framework with sizes larger than 1 square meter. Late Triassic: Jabal Kawr, Oman. 3 Solenopora cassiana Flügel. Longitudinal (LS) and cross-section (CS). Note distinct filaments exhibiting pearl-shaped outlines (arrow) that correspond to lateral connections. Reef facies. Late Triassic (Cassian Formation, Carnian): Seeland Alm near Cortina d’Ampezzo, Southern Alps, Italy. 4 Parachaetetes johnsoni Maslov. Note different preservation obscuring the algal structure in the lower part of the thallus. Same locality as –> 1 5 Parachaetetes johnsoni Maslov. Horizontal partitions are situated at the same level between tubes, and separate the long cells. The well-preserved alga is encrusted by foraminifera. Same locality as –> 1. 6 Transported solenoporacean red algae exhibiting the common preservation feature. Recrystallization of the originally aragonitic thalli has obliterated the algal structure, which is only recognizable in traces. In contrast, micro- and macroborings (black and white arrows) are well preserved because of early infilling with calcite. Early Permian: Teufelsschlucht, Karawanken Mountains, Slovenia. 7 Solenopora. Longitudinal section. Note growth zonation and microborings (arrows). Early Cretaceous: Subsurface, United Arabian Emirates. 8 Parachaetetes asvapatii Pia. The filaments are differentiated into ‘long’ and ‘short’ cells (occurring in the dark zones) similar to cell differentiation in modern floridean red algae. The cosmopolitan species is known from the Albian to the Paleocene. Early Tertiary: Italy. –> 2: Weidlich et al. 1993
Plate 55: Solenoporacean Red Algae
419
Ancestral Red Algae
420
sic, and locally abundant in Jurassic and Cretaceous patch reefs. Common genera are: Solenopora, Parachaetetes, Cordilites (Late Cretaceous), Elianella. Basics: Solenoporacean red algae Johnson, J.H. (1960): Palaeozoic Solenoporaceae and related red algae. – Quarterly, Colorado of School Mines, 55, 177
Mamet, B., Roux, A. (1977): Algues rouges dévoniennes et carbonifères de la Téthys occidentale, part 4. – Revue de Micropaléontologie, 19, 215-266 Moussavian, E. (1989): Über die systematische Stellung und Bestimmungskriterien der Solenoporaceen (Rhodophyceae). – Courier Forschungsinst. Senckenberg, 109, 51-91 Wright, V.P. (1985): Seasonal banding in the alga Solenopora jurassica from the Middle Jurassic of Gloucestershire, England. – Journal of Paleontology, 59, 721-732
Plate 56 Late Paleozoic Ancestral Red Algae and Problematic Red Algae The terms ‘ancestral red algae’ or ‘ancestral coralline algae’ refer to Late Paleozoic algae characterized by differentiated blade-like thalli with hypo- and perithallus structures, thus reminding one of corallinacean algae. Other taxa exhibit branched (–> 6) or irregular, encrusting forms (–> 4). Many taxa exhibit yellowish hyaline structures pointing to a primary aragonite skeletal mineralogy that is also indicated by high strontium concentrations in Archaeolithophyllum. Some genera (Archaeolithophyllum; –> 1, 2, 8) contribute to the formation of mounds, others are common in open-marine shoals but also in restricted lagoonal carbonates. Archaeolithophyllum appears to be related closely to modern aragonitic Peyssoneliaceae. A few taxa, sometimes called ‘pseudoalgae’, have been attributed to sponges, but are probably algae (Ungdarella, –> 7; Fasciella, –> 3, 4; Komia, –> 6). Some of these taxa are index fossils, particularly useful for biostratigraphic differentiation based on foraminiferal assemblage zones of Early Carboniferous shelf carbonates. Fig. 10.9. Late Paleozoic ungdarellacean algae (Ungdarella, Komia) and stacheinacean algae (Aoujgalia, Efluegelia, Stacheoides). The interpretation of these mm-sized encrusting microfossils is controversial, but many workers assign them to red algae because of the structural similarity with genuine rhodophytes. After Chuvashov et al. (1987). 1 Archaeolithophyllum Johnson. Ancestral coralline alga. Criteria: Crustose multifoliate sheets of differentiated cells. Prominent hypothallus with large polygonal cells (arrows), poorly developed perithallus. Single high-arched conceptacles in the perithallus tissue. Thalli often recrystallized, yellowish in transmitted light. Originally aragonitic skeleton. Top right: Cross-section with distinct cell structure. Time range: Early Carboniferous to Early Permian, abundant in Middle and Late Carboniferous. The genus is common in shallow-marine mounds and shoals. Late Pennsylvanian (Auernig Formation): Gugga, Carnic Alps, Austria. 2 Archaeolithophyllum Johnson. Note central hypothallus (HT) cells. The perithallus (arrow) is largely destroyed. Same locality as –>1 3 Fasciella Ivanova. Formerly named Shartymophycus. Regarded as red or green alga. Criteria: Encrusting nodular thallus, composed of irregular concentric layers of elongated cells. Walls yellowish, hyaline. Time range: Carboniferous (Viséan to Bashkirian). Common in open-marine shelf grainstones. Early Carboniferous (Viséan): Betic Cordillera, Spain. 4 Fasciella Ivanova. Note nodular growth and irregular layers (arrow). Middle Carboniferous: Nassfeld, Carnic Alps, Austria. 5 Efluegelia Vachard, enclosed in an oncoid. Probably an ancestral red alga. Criteria: Encrusting thallus consisting of a few thin rows of minute, subparallel laminae with quadratic cells. Walls yellowish, hyaline. Commonly enclosed in biogenic crusts. Time range: Carboniferous and Early Permian. Early Permian (Asselian): Carnic Alps, Austria. 6 Komia Korde. Criteria: Ramified robust thallus composed of rectangular and quadrangular cells. Walls yellowish, hyaline. Time range: Middle Carboniferous to Early Permian. Common in shelf carbonates. Late Pennsylvanian (Gzhelian): Hüttenkofel, Carnic Alps, Austria. 7 Ungdarella Korde. Criteria: Ramified thallus, perithallus with long thin rows of subquadratic cells. Time range: Carboniferous and Early Permian. Common in Middle Carboniferous shelf carbonates. Carboniferous (Viséan/Namurian): Rif Mountains, northern Morocco. 8 Rhodoid with thin crusts of Archaeolithophyllum lamellosum Wray. The alga has coated phylloid algae (Eugonophyllum, E; Pl. 58/3) as well as a rugose Dibunophyllum-type coral (C). Note the loose organization of the cm-sized oncoid and the inclusion of mud chips (MC), indicating that coating was a nonselective process. Late Pennsylvanian (Leavenworth limestone, Virgilian; transgressive part of a megacyclothem): Osage County, northern Oklahoma, U.S.A. –> 8: Toomey 1983
Plate 56: Ancestral Red Algae
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Box 10.2. Late Paleozoic ancestral red algae common in reef limestones, and possible red algae, abundant in platforms and ramps. Although the systematics of the latter are controversial and need substantiation, the fossils have a high potential in microfacies analyses. Specific associations of these algae and of foraminifera allow a detailed biozonation and correlation of Carboniferous shelf carbonates (Armstrong and Mamet 1977, Mamet et al. 1987). Archaeolithophyllacean red algae Central hypothallus coaxial, consisting of large polygonal and rectangular cells. Perithallus poorly developed, cells much smaller. Large conceptacles at the periphery, commonly exhibiting a yellowish color in transmitted light. Often strongly recrystallized. Common taxa: Archaeolithophyllum (Pl. 56/1, 2, 8). Viséan to Late Permian. Cosmopolitan. Abundant in Pennsylvanian and Early Permian reefs (reef mounds and mud mounds). Ungdarellacean red algae Millimeter-sized elongated and often ramified encrusting nodules, exhibiting rectangular and subquadratic cells that are sometimes differentiated according to size and arrangement. Skeletal elements hyaline, often with a golden or yellowish tint in transmitted light. Common taxa: Ungdarella (Pl. 56/7, Pl. 108/7, 8), Komia (Pl. 56/6). Early Carboniferous to Permian. Cosmopolitan. Abundant in Late Viséan to Moscovian shelf carbonates. Stacheinacean and related red algae Millimeter-sized microfossils encrusting a central support or a substrate. Characterized by a succession of arched or subquadratic cells arranged irregularly or in rows. Skeletal elements hyaline. Commonly yellowish in transmitted light. Common taxa: Cuneiphycus, Efluegelia (Pl. 56/5), Epistacheoides (Pl. 108/9), Stacheia (Pl. 108/10), Stacheoides. Early Carboniferous to Permian. Common in Viséan to Moscovian platform and ramp carbonates. Rare in Early Permian shelf carbonates.
10.2.1.4 Ancestral Red Algae and Problematic Red Algae Several thin-section fossils, predominantly of Late Paleozoic age, are believed to be ancestors of younger red algal groups, or are assigned to red algae, more or less persuasively. Some exhibit thalli composed of a cellular tissue known from true red algae taxa (e.g. Archaeolithophyllum). Other taxa can not be related to existing red algal groups and are therefore assigned to separate algal groups (e.g. Ungdarella) or have been excluded from the algae and compared with foramin-
Udoteacean Algae
ifera, hydrozoans or sponges. Two groups of these microfossils are common in Late Paleozoic shallow-marine carbonates and used in the stratigraphic and environmental subdivision of platforms and ramps (Fig. 10.9; Box 10.2). Basics: Ancestral corallines and problematic red algae Armstrong, A.K., Mamet, B. (1977): Carboniferous microfacies, microfossils and corals, Lisburne Group, Arctic Alaska. – Geological Survery Professional Paper, 849, 144 pp. Johnson, J.H. (1956): Archaeolithophyllum, a new genus of Paleozoic coralline algae. – Journal of Paleontology, 30, 53-55 Mamet, B., Roux, A., Nassichuk, W. (1987): Algues carbonifères et permiennes de l’Arctique canadien. – Geological Survey of Canada, Bulletin, 342, 1-143 Wray, J.L. (1964): Archaeolithophyllum, an abundant calcareous algae in limestones of the Landsing Group (Pennsylvanian), southeastern Kansas. – Kansas Geological Survey Bulletin, 170, 13 pp. Further reading: K129
10.2.1.5 Udoteacean Green Algae and Gymnocodiacean Algae Most modern benthic green algae live in freshwater environments, about 10% of the species also occur in marine settings. Ancient green algae are almost exclusively marine fossils. Benthic calcified green algae are represented by the Udoteaceae and the Dasycladaceae. The systematic position of the extinct Gymnocodiaceae (green or red algae) is still under discussion, but better arguments are in favor of green algae (Mu 1991; Bucur 1994). Other algal groups related to or interpreted as benthic green algae are the Early Paleozoic receptaculitids (common in Ordovician and Silurian, ranging to the Permian) and cyclocrinitids (Ordovician and Silurian). Both groups differ in the size of the globular shape (several centimeters) of typical calcareous algae, but resemble dasyclad green algae in their internal organization. They have a central main axis and radially arranged tightly packed lateral branches forming hexagonal or rhombohedral facets on the outer surface (Rietschel 1977; Nitecki et al. 1987; Beadle 1988, 1991). Some of the fossils included within the informal group of ‘phylloid algae’ are most probably udoteacean green algae (Sect. 10.2.1.6). Fossil records of marine planktonic green algae are known from the Volvocales and Chlorococcales (Sect. 10.2.1.9). Calcifying planktonic green algae also contribute to the formation of modern and Tertiary freshwater carbonates (Müller and Oti 1981).
Udoteacean Algae
Udoteacean green algae This algal group is well-known to sedimentologists because of the fundamental importance of Halimeda in the production of carbonate muds and sands in modern shallow seas. Which name is the right one? Taxonomists like to change systematic categories. Both, erect and filamentous forms were formerly called Codiaceae, a group of the Bryopsidales (formerly Siphonales). Since modern Codiaceae include biologically different types, the group has recently been split up into Udoteaceae, morphologically represented by the genus Halimeda, and into the non-calcified Codiaceae, represented by the genus Codium. Most erect fossil forms are now included within the Udoteaceae. Some authors use the name Halimedaceae which has priority over Udoteaceae (Hillis 1991). Filamentous fossil taxa are attributed to the porostromate Cyanobacteria or are sometimes regarded as members of separate groups of siphonate green algae. Definition and morphology: Udoteacean algae are plants consisting of cm-sized segmented thalli attached to the substrate by rhizoids. The algae develop erect,
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filamentous (nodular), delicately branched, and phylloid (bladed) growth forms. The basic pattern of the internal organization is differentiated into a central region (medulla) and a peripheral zone (cortex). The medullary filaments are arranged parallel to the axis of the thallus, interwoven or branched, and give rise to lateral cortical filaments. The medulla is poorly calcified. Hence, in thin sections the central region often exhibits no structures and is strongly recrystallized. The thinner cortical filaments differ greatly in shape and are straight, swollen or constricted. Calcification of living udoteaceans varies in intensity and takes place by precipitation of needle-shaped aragonite crystals on the surface of filament membranes within the thallus. In fossil udoteaceans, calcification is mainly recorded in the cortex. Many udoteacean fossils appear as calcareous ‘tubes’ or blades, exhibiting a peripheral zone perforated by pores and an often strongly recrystallized or spar-filled central region with faint or no filaments. Fossil reproductive structures are commonly absent, reported occurrences are questionable. Bladed udoteaceans are described in the context of the discussion of phylloid algae (Sect. 10.2.1.6). Ecological constraints: Modern codiacean green algae occur in normal saline warm waters of tropical shelf,
Fig. 10.10. Udoteacean algae. Species richness and sedimentological role. A: The figure shows the number of species of three closely related genera characterized by similar patterns in the differentiation of cortical filaments (Pl. 57). In recent papers some more new species have been described, but the general trend in species richness is the same as shown in the figure. Halimeda was also reported from Permian deposits. B: The contribution of erect udoteaceans to the sand fractions of platform and reef carbonates increased notably during the Tertiary and reached in Miocene lagoons values of more than 75% similar to those of today. PL: Pleistocene. Modified from Flügel (1991). Pictures of the morphotypes after Elliott (1982).
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lagoonal and reefal environments. They are commonly concentrated in depths of a few meters, but extend to about 100 m (Halimeda thickets). The plants grow on sand and mud bottoms below intense water agitation. Distribution: Calcified udoteaceans appear in the Ordovician. The algae are common in the Ordovician, abundant in the Early and Middle Devonian, and relatively rare in the Late Paleozoic and Triassic. Calcified udoteacean algae are locally common in Late Triassic and Jurassic limestones and abundant in Early Creta-
Gymnocodiacean Algae
ceous carbonate platforms. They became abundant during the Cenozoic. Many taxa appear to be restricted to stage or substage intervals. Udoteaceans related to or identical to Halimeda are common in shallow-marine carbonates of Cretaceous and Early Cenozoic age, where they occur in similar environments as today (Fig. 10.10). For a survey of the Phanerozoic species see Bassoullet et al. (1983). Some rock-building fossils of problematic systematic position that have been tentatively placed in the udoteacean or codiacean algae will be discussed in
Plate 57 Calcareous Udoteacean Green Algae and Gymnocodiacean Algae Modern udoteacean green algae are a major source of calcareous mud from postmortem disintegration of their skeletons. Halimeda contributes to the formation of carbonate muds and sands in shallow environments but also to Halimeda ‘reefs’ occurring today on deeper slopes. Udoteaceae are calcareous green algae with erect (–> 1-6), nodular or bladed growth forms. The internal organization of erect segmented forms is exemplified by Halimeda (–> 1, 2). The thallus consists of a central zone (medulla composed of filaments arranged parallel to the axis) and the peripheral cortex zone exhibiting smaller branched filaments. Calcification (aragonite) takes place between and around the filaments. Ancient udoteaceans, known since the Ordovician, are predominantly classified according to growth form (branched, unbranched, segmented, non-segmented, bladed), and the structure of the medulla and cortex (size, shape, number, arrangement and branching mode of the filaments). Gymnocodiacean algae (–> 7, 8) are extinct (Devonian to Tertiary) cylindrical or sack-like algae commonly occurring as mm-sized fragments. They developed erect, branched or unbranched, segmented or non-segmented thalli composed of central filamentous medulla and a peripheral thin cortex zone with pores. The primary mineralogy is considered to have been aragonite. Mass accumulations in Late Permian and Early Cretaceous Tethyan carbonate platforms are indicative of warm, shallow seas with low to moderate water energy. 1 Halimeda. Longitudinal section of a segment. The internal structure is characterized by loosely packed filaments forming an outer cortical (C) zone of thin lateral branches that develops from a broad central zone (medulla, M) with thicker filaments. Dark areas between the filaments are carbonate. Recent reef lagoon: Florida Bay, U.S.A. 2 Halimeda. Note the differences in the size and arrangement of filaments. The medullary filaments are tangled and form utricles (swellings; arrow) passing into the cortical zone. Recent reef lagoon: French Reef, Florida Bay, U.S.A. 3 Boueina marondei Flügel. Longitudinal (LS) and tangential section (TS). Note the gross similarity of the internal structure with Halimeda. M: Central medulla, C: Peripheral cortex. Late Triassic (Norian): Western Thailand. 4 Pseudopalaeoporella lummatonensis (Elliott). The cortex (C) consists of densely spaced, branched and subparallel filaments. Middle Devonian (Eifelian): Graz, Austria. 5 Litanaia graecensis Hubmann. Cross section of a cylindrical thallus, differentiated into central medulla (M) with thick filaments, and peripheral cortex (C). Note bifurcation and thickening of cortical filaments (arrow). Litanaia and similar taxa (e.g. Lancicula) are common in Early and Middle Devonian shelf carbonates. Inner ramp (see Pl. 135). Middle Devonian (Eifelian): Graz, Austria. 6 Cross section of Boueina hochstetteri Toula, a cosmopolitan Late Jurassic and Cretaceous species occurring predominantly in Early Cretaceous platform carbonates. Cretaceous (Early Aptian): Southern Carpathians, Romania. 7 Gymnocodiacean grainstone, consisting of broken Permocalculus thalli. Arrows point to thin-walled fragments with fine densely spaced branched cortical filaments. Permocalculus is characterized by tightly bound segments with a wide, poorly calcified medulla and a narrow cortex intersected by numerous branched pores (see Fig. 10.11). Parallel alignment of algal fragments and some coated shells indicates current deposition. Back-reef facies. SMF 18-GYMNO. Middle Permian (Wordian): Djebel Tebaga, southern Tunisia. 8 Compacted gymnocodiacean floatstone, consisting of densely packed thalli of Gymnocodium. Arrows point to longitudinal sections exhibiting the characteristic irregular outline of the thalli caused by thick peripheral filaments. SMF 18GYMNO. Late Permian: Taurus Mountain Range, Turkey. –> 4: Courtesy of B. Hubmann (Graz); 5: Hubmann 1990; 6: Courtesy of I. Bucur (Cluj-Napoca)
Plate 57: Udoteaceans and Gymnocodiaceans
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Gymnocodiacean Alagae
Sect. 10.2.1.5 (Nuia, Pl. 98/1) and Sect. 15.1.2 (Microcodium, Pl. 128/4). Sedimentological significance: In modern shallowmarine, low- to moderate energetic environments udoteaceans are a major source of calcium carbonate mud from postmortem disintegrations of their skeletons. Halimeda is abundant (1) in tropical lagoonal environments (Hillis 1991) similar to fossil counterparts, (2) in reef and forereef environments, and (3) on deep slopes, forming banks and reefs (Drew 1983; Marshall and Davies 1988). Ancient Halimeda-form algae contributed to sediment production, locally in the Late Triassic (Pl. 57/3) and more commonly in the Cretaceous (Pl. 57/6) and significantly since the Early Tertiary. The sedimentological role of Paleozoic udoteaceans seems to be restricted to some Ordovician and Devonian shelf carbonates (Pl. 57/4, 5).
Gymnocodiacean algae Systematic position: These algae were originally regarded as codiaceans, but later attributed to the red algae because of some similarities of the reproductive structures to those in a genus of recent chaetangiacean red algae (Elliott 1955). The affinity of the group is not yet settled. A red algal affinity has been accepted by many authors, but was recently questioned again in favor of a possible assignment to green algae (Mu 1991; Bucur 1994). The gymnocodiaceans resemble fossil erect udoteaceans in growth form, internal structure, and palecological and paleogeographical distribution, but seem to differ in containing internal reproductive organs (Mu and Riding 1983).
Fig. 10.11. Longitudinal section of Permocalculus. This gymnocodiacean genus is of rock-building importance in Permian and Cretaceous carbonates. Late Permian: Bergama, northwestern Anatolia, Turkey. Scale is 1 mm.
Fig. 10.12. Gymnocodiacean limestone. Various sections through segments of Gymnocodium bellerophontis Rothpletz. These algae were poorly calcified in the central part (which is commonly filled with sediment) and strongly calcified in the peripheral cortex, where the branched filaments may terminate in the form of funnels or bowls, appearing as polygonal structures in tangential sections. Faint relics of medullary filaments occur within the grayish interior of two sections. Middle Permian: Balya, Anatolia, Turkey. Scale is 1 mm.
Morphology: The extinct algae have erect, branched or unbranched, segmented or unsegmented cylindrical thalli, consisting of cylindrical, oval, or pyriform segments. The internal structure of these segments is characterized by a medullar zone and peripheral cortical zones. Calcification (primary aragonite) was significantly stronger in the cortical than in the medullar zone; the center of the thalli, therefore is commonly filled with sediment or cement. Medullar filaments run parallel along the axis, cortical filaments are branched several times and form a distinct pattern. Reproductive structures are represented by oval or spherical cavities in the cortex or in the area between the cortex and medulla. Distribution: The records of this group range from the Middle Devonian (Mamet et al. 1994) to the Late Tertiary, with most abundant occurrences at the Late Permian (Mu 1991) and common occurrences in some Cretaceous carbonates (Elliott 1958; Bucur 1994). Limestones composed more or less completely of gymnocodiacean algae are common in the Middle and Late Permian of the Tethys (e.g. Dolomite Mountains, South Tyrol) and most abundant in the Capitanian and the Early Changsiangian stages. Sedimentological significance: During the Late Permian, gymnocodiaceans were important sediment pro-
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Phylloid Algae
ducers in shallow, low to moderate energetic shelf seas (Fig. 10.12, Pl. 122/3). They acted as bafflers and contributed to the formation of bioclastic wackestones and packstones in ramp and outer shelf environments through fragmentation and transport of their thalli. Common geographically widespread taxa are Gymnocodium (restricted to the Permian) and Permocalculus (Permian to Aptian; Fig. 10.11). Basics: Udoteacean green algae and gymnocodiacean algae Bassoulet, J.P., Bernier, P., Deloffre, R., Genot, P., Poncet, J,, Roux, A., Jaffrezo, M. (1983): Les algues udoteacées du Paléozoique au Cénozoique. – Bull. Centre Recherches Explor.-Prod. Elf-Aquitaine, 7, 449-621 Bucur, I.I. (1994): Lower Cretaceous Halimedaceae and Gymnocodiaceae from the Southern Carpathians and Apuseni Mountains (Romania) and the systematic position of the Gymnocodiaceae. – Beiträge zur Paläontologie, 19, 13-37 Deloffre, R. (1992): Révision des Gymnocodiaceae (Algues rouges, Permien-Miocène). Taxonomie, biostratigraphie, paléogéographie. 3ème partie. Inventaire taxonomique critiques des espèces de Gymnocodiacées du Permien et du Trias. – Revue de Micropaléontologie, 35, 23-37 Elliott, G.F. (1955): The Permian alga Gymnocodium. – Micropaleontology, 1, 83-90 Elliott, G.F. (1958): Algal-debris facies in the Cretaceous of the Middle East. – Palaeontology, 1, 254-259 Flügel, E. (1988): Halimeda: paleontological record and paleoenvironmental significance. – Coral Reefs, 6, 123-130, Berlin Hillis, L. (1991): Recent calcified Halimedaceae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 167-188, Berlin (Springer) Hubmann, B.(1993): Zur heutigen Situation der Systematik und Taxonomie devonischer Grünalgen am Beispiel ‘Lancicula’. – Sitzungsberichte der Österreichischen Akademie der Wissenschaften, mathematisch-naturwissenschaftliche Klasse, 200, 151-161 Mamet, B., Roux, A., Shalaby, H. (1984): Rôle des Algues calcaires dans la sédimentation Ordovicienne de la platforme du Saint-Laurent. – Geobios, Mémoir Spécial, 8, 261-269 Mankiewicz, C. (1988): Occurrence and paleoecological significance of Halimeda in late Miocene reefs, southeastern Spain. – Coral Reefs, 6, 271-279, Berlin Marshall, J.F., Davies, P.J. (1988): Halimeda bioherms of the Great Barrier Reef. – Coral Reefs, 6, 139-148, Berlin Mu, Xinan (1991): Fossil Udoteaceae and Gymnocodiaceae. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 146-166, Berlin (Springer) Mu, Xinan, Riding, R. (1983): Silicified gymnocodiacean algae from the Permian of Nanjing, China. – Palaeontology, 26, 261-276 Roux, A. (1991): Révision des Gymnocodiaceae (Algues rouges, Permien-Crétacé). Taxonomie, biostratigraphie, paléogéographie. 2ème partie. Inventaire taxonomique critiques des espèces de Gymnocodiacées du Permien et du Trias. – Revue de Micropaléontologie, 34, 136-173
Fig. 10.13. Reconstruction of a phylloid algal community. The algae dominate the sea bottom. The algal leaves are home to various encrusting foraminifera and polychaete worms. Brachiopods live in sheltered areas below algal blades. Gastropods and echinids graze the algal meadow. Some bivalves occur on and within the sediment. Slightly modified from Toomey (1981). Roux, A., Deloffre, R. (1990): Révision des Gymnocodiaceae (algues rouges, Permien-Crétacé). Taxonomie, biostratigraphie, paléobiogéographie. 1re partie: Généralités sur la famille. – Revue de Micropaléontologie, 32, 123-137 Further reading: K129
10.2.1.6 Phylloid Algae Definition and morphology: The informal name refers to the leaflike form of the thalli, which consist of variously shaped calcified blades comparable in shape to the modern green alga Udotea and to non-calcified brown algae (Pray and Wray 1963). The fossil record of these algae consists of fragmented blades exhibiting various internal structures. Some phylloid algae are differentiated into central medullary and cortical zones (Anchicodium, Ivanovia, Eugonophyllum) or consist of only one layer of pores (Calcifolium). The medullary zone is commonly replaced by coarse sparite. Primary calcification consists of aragonite. Taxonomic differentiation is based on the structure of the cortex. Some phylloid algae have internal cellular structures that relate them to red algae (Archaeolithophyllum, see Sect. 10.2.1.4). Systematic position: Most authors place genera of ‘phylloid algae’ into various groups of green or red algae (Udoteaceae; Peyssoneliaceae). The green algal interpretation is, at least for Eugonophyllum, supported by strong arguments (Kirkland et al. 1993). Some authors include Eugonophyllum (Pl. 58/5), Ivanovia (Pl. 58/4) and Anchicodium in the separate green algal group Anchicodiaceae.
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Distribution: Predominantly Middle Carboniferous to Late Permian. A few records are known from the Middle Devonian and the Late Triassic (Reid 1986; Mamet and Preat 1983). The algae flourished during the Late Pennsylvanian and Early Permian (Asselian and Sakmarian). Sedimentological significance: Phylloid algae are dominant rock constituents in bedded shelf carbonates and contribute significantly to the formation of Pennsylvanian and Early Permian mud mounds, reef mounds, and banks. Mounds and banks dominated in shelf settings and at shelf margins (Wahlman 2002). Phylloid mounds are characterized by flat bases, undulatory surfaces and high micrite content. Rapid growth and recovery rates of phylloid algae were ideal for life in settings strongly controlled by short-term sea-level fluctuations. Stacked mounds reach a thickness of sev-
Phylloid Algae
eral decades of meters. Banks are more extended and may cover areas of more than 50 km2. Location and development of phylloid algal mounds seems to be controlled by paleocurrent directions and nutrient levels. Phylloid algae acted mainly as important sediment bafflers, but some of them have also built reefs by constructing organic frameworks (Samankassou and West 2002). Some phylloid algae (e.g. Eugonophyllum) exhibit a felt of aragonite needles interspersed between small parallel medullar tubes, similar to Halimeda. These needles might have been a source for carbonate muds of mounds and banks. Laterally extended phylloid algal buildups are major hydrocarbon reservoirs in the southwestern United States (Montgomery et al. 1999) because of high intraskeletal and interskeletal porosities and common subaerial exposure. The depositional texture of the phyl-
Plate 58 Phylloid Algae ‘Phylloid algae’ is a non-taxonomic term referring to predominantly Late Paleozoic calcareous algae characterized by leaf-like growth forms (Fig. 10.13). The group includes udoteacean green algae (Pl. 57), ancestral red algae (Pl. 56) as well as squamariacean red algae with erect or inclined growth forms. Viewed in cross-sections in limestones, phylloid algae appear as thin sinuous lines, sometimes looking like ‘potato-chips’. Many phylloid algae are recrystallized and lack sufficient internal features for taxonomic identification. Epifluorescence petrography assists in overcoming this difficulty (Dawson 1992). Generally, the central (medullary) part of the blades is filled with carbonate cement (–> 5). The very poor preservation as well as rare records indicate a primary aragonite mineralogy. Taxonomically important criteria are the shape and arrangement of cortical branches (–> 2, 4 and 5). Blades can be several centimeters long and up to 1 mm thick. 1 Phylloid algal floatstone. All bent, curved and straight elements are recrystallized algal blades. Mollusk shells differ in the shape and lack of peripheral pores. Many algal blades are bored (B). Other biotic elements are scarce (gastropods - G, and rare echinoderm plates - E). The chaotic distribution of the blades resembles the edgewise fabrics of conglomerates, and points to reworking and deposition caused by storm currents. Edgewise fabrics have been described mainly from peritidal environments, but also occur in subtidal settings, e.g. as storm or swell lags deposited in shoreface environments, lateral accretion lags deposited in tidal channels, or as transgressive lags formed by ripping up of the underlying strata as the sea level rose during platform drowning (Demicco and Hardie 1994). SMF 5. Late Pennsylvanian (Lower Pseudoschwagerina beds, Gzhelian): Carnic Alps, Austria. 2 Neoanchicodium Endo. Blades with poorly calcified tubes and one or two rows of ellipsoidal cells in the peripheral zone appearing as a ‘ring of pearls’ in transverse sections (arrow). The genus is common in Early Permian open-marine platform carbonates. Early Permian shelf platform: Forni Avoltri, Carnia, Italy. 3 Eugonophyllum Konishi and Wray, an udoteacean green alga. The thallus consists of large blades standing erect on the substrate. The alga is most commonly preserved as isolated and often heavily recrystallized fragments. Filaments in the central medulla are often obliterated by recrystallization. The blades are strongly encrusted by foraminifera (arrows). E: Epimastopora (see Pl. 60/4). Early Permian: Carnic Alps, Austria. 4 Ivanovia Khvorova. The central (medullary) zone is recrystallized. The well-preserved cortical zone exhibits thin branches oriented perpendicular to the medullary zone. The sample comes from phylloid algal packstones overlying a Archaeolithoporella-Tubiphytes reef. Middle Permian (Wordian): Djebel Tebaga, southern Tunisia. 5 Eugonophyllum mulderi Racz. Well-preserved cortical zone (C) with crescent-shaped branches. The medulla (M) is obliterated by dissolution and cement filling. E: Echinoderm plate. Late Carboniferous (Gzhelian): Hüttenkofel, Carnic Alps, Austria. –> 4: Toomey 1991
Plate 58: Phylloid Algae
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loid algal sediment is a principal factor in the localization of effective porosity and permeable zones. Basics: Phylloid algae Cys, J.M. (1985): Lower Permian phylloid algal mounds, southern Tatum basin, southwestern New Mexico, U.S.A. – In: Toomey, D.F., Nitecki, M.H. (eds.): Paleoalgology. Contemporaneous research and applications. – 179-187, Berlin (Springer) Kirkland, B.L., Moore, C.H., Jr., Dickson, J.A.D. (1993): Well preserved, aragonitic phylloid algae (Eugonophyllum, Udoteaceae) from the Pennsylvanian Holder Formation, Sacramento Mountains, New Mexico. – Palaios, 8, 111-120 Montgomery, S.L., Chidsey, T.C., Eby, D.J., Lorenz, D.M., Culham, W.E. (1999): Pennsylvanian carbonate buildups, Paradox Basin: increasing reserves in heterogeneous, shallow shelf reservoirs. – American Ass. Petrol. Geol., Bull., 83, 193-210 Toomey, D.F. (1980): History of a Late Carboniferous phylloid algal bank complex in northeastern New Mexico. – Lethaia, 13, 249-267 Toomey, D.F., Winland, H.D. (1973): Rock and biotic facies associated with a middle Pennsylvanian phylloid algal buildup, Nena Lucia field, Nolan County, Texas. – American Ass. Petrol. Geol., Bull., 57, 1054-1074 Wahlman, G.P. (2002): Upper Carboniferous-Lower Permian (Bashkirian-Kungurian) mounds and reefs. – In: Kiessling, W., Flügel, E., Golonka, J. (eds.): Phanerozoic reef patterns. – SEPM Special Publication, 72, 271-338 Further reading: K129
10.2.1.7 Dasyclad Green Algae Dasyclads are the most important fossil calcareous algae used in microfacies analyses. These algae appeared in the Cambrian and were important rock builders during the Late Paleozoic, Mesozoic and Early Cenozoic. Because of their value as palecological proxies and for stratigraphical zonations, dasyclad algae are key elements in interpretating of ancient shallow-marine shelf carbonates. Excellent overviews on the biology of dasyclad algae and the methods used in their analysis were provided by Berger and Kaever (1992) and De Castro (1997). Deloffre (1988) presented a key for the distinction of suprageneric taxonomic units. Most dasyclads are recorded in thin sections as fragments. Generic and specific determinations demand the reconstruction of the morphology from random cuts and variously oriented sections (see Fig. 10.2). Definition, calcification and morphology: Dasyclads are a group of benthic calcified unicellular green algae. Living dasyclads are mm- to cm-sized, upright growing plants, which are attached to firm or soft substrates by rhizoids. Characteristically, the thallus consists of a long central stem (also called central cylinder or main
Dasyclad Algae
axis) which bears one or more whorls of lateral branches (laterals). These primary branches may subdivide distally forming one to three or more orders of laterals. Reproductive organs occur on or between branches, or within the central stem. Aragonite is precipitated around the stem (or parts of the stem) and between the branches. As a consequence, fossil dasyclads are represented as molds in which the area of the stem cell appears as a spar- or sediment-filled cavity, surrounded by a calcified zone exhibiting the branches as pores. The recognition of the diagnostic criteria is controlled by the extent of calcification. For major diagnostic criteria see Fig. 10.14. (1) Morphology: Most dasyclads form unbranched thalli; only a few form a branched thallus. Unbranched thalli exhibit various shapes including (a) rod-shaped (tubelike; e.g. Macroporella, Gyroporella, Velebitella) and spherical thalli (e.g. Mizzia), (b) clavate thalli (which increase in size during their development resulting in pear-shaped forms), and (c) stalked thalli with a spherical or barrel-shaped capitulum (e.g. Bornetella) or a cap formed by whorls of rays (e.g. Acicularia, Acetabularia). Most of these types may be unsegmented or segmented. Branched thalli are represented by dichotomous tube-like forms (e.g. Anthracoporella) and articulate forms appearing like beads on a string (e.g. Cymopolia). (2) External or internal configuration of the thalli expressed by undulation and articulation: Undulation refers to ring-shaped swellings at the thallus surface, annulation to ring-like lacing of the thallus surface not separated from each other (Neoteutloporella, Pl. 59/3; Salpingoporella, Pl. 62/2), perannulation to pronounced annular lacings in the calcareous skeleton connected by noncalcified segments (Eovelebitella, Pl. 59/1), and fissuration to annular fissures in the calcareous skeleton. Intusannulation is a term for more or less pronounced undulations of the main axis diameter engraved into the inner side of the central cylinder. (3) Arrangement of laterals (primary branches) at the main axis: Aspondyl describes an irregular arrangement at the main axis (Anthracoporella, Pl. 59/6; Teutloporella, Pl. 61/3), euspondyl a regular arrangement in whorls like the spokes of a wagon wheel (Neoteutloporella, Pl. 59/3; Salpingoporella, Pl. 63/2), and metaspondyl an arrangement of primary branches into clusters of regular whorls (Eovelebitella, Pl. 59/1, Pl. 60/3; Diplopora, Pl. 59/2). The arrangement of the laterals is one of the most effective criteria for defining dasyclad families. Text continued on p. 434
Dasyclads: Diagnostic Criteria
Fig. 10.14. Major diagnostic criteria of dasyclad green algae. After Berger and Kaever (1992).
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Reproductive organs of Dasyclads
Fig. 10.15. Position of the reproductive organs in dasyclad algae. After Berger and Kaever (1992).
Plate 59
Dasyclad Green Algae: Taxonomic Criteria
Main criteria in the differentiation of dasyclads are the growth form and shape of the thallus and the central stem, the arrangement, shape and the number of branching of laterals, and the position of reproductive organs (see Figs. 10.14 and 10.15). The plate shows examples of some of the diagnostic morphological criteria. 1 Eovelebitella occitanica Vachard. Cross section (CS), longitudinal (LS) and tangential sections (TS). Rod-shaped, annulate thallus. Primary branches (white arrows) are arranged in clusters of regular whorls (metaspondyl), forming a succession of subspherical articles, black arrows). The laterals increase in size towards the periphery (phloiophore type). The genus is locally abundant in Tournaisian and Viséan shelf carbonates. Early Carboniferous (Viséan). Southern Spain. 2 Diplopora phanerospora Pia. Cross section and longitudinal section. Club-shaped articulated thallus. Primary poikilophore and vesicular branches of one single order are grouped into tufts (metaspondyl). Calcified reproductive organs (cysts, arrow) are situated within the main axis (endospore type). Late Triassic (Zlambach beds, Rhaetian): Bavarian Alps, Germany. 3 Neoteutloporella socialis (Praturlon). Longitudinal section of densely spaced autochthonous dasycladacean thalli that form algal meadows and biostromes in outer platform settings of the Alpine-Mediterranean Tethys (Dragastan et al. 1987; Bodeur 1995). Cylindrical articulated thallus with short primary and long secondary branches (black arrows) arranged oblique to the main axis (white arrow) in close-set euspondyl whorls (W). Laterals become thinner at their distal ends (trichophore type). Stem cell with swellings and narrowings (annulation). The species is commonly associated with corals and ‘hydrozoans’ (e.g. Ellipsactinia, –> Pl. 81/5). Species range Oxfordian to Tithonian. Late Jurassic (Tithonian): Abruzzi National Park, central Apennines, Italy. 4 Mizzia velebitana Schubert. Oblique and tangential sections (TS). The thallus consists of spherical articulated segments. The central stem is barrel-shaped and has constrictions at the joints between segments (annulation). Only primary branches, arranged in concentric rows. Branches enlarge in diameter near the exterior, producing a honeycomb pattern on the surface. Late Permian: Bergama, western Turkey. 5 Mizzia wackestone. Isolated spherical segments. Note the difference in size, indicating transport of the algal fragments. Late Permian: Taurus Mountain Range, Turkey. 6 Anthracoporella spectabilis Pia. Longitudinal (LS), cross section (CS) and tangential section (TS). Rod-shaped, tube-like thalli. Closely spaced irregularly arranged (aspondyle) primary laterals. The thalli of this species are conspicuously large and are several centimeters long. Late Carboniferous (Lower Pseudoschwagerina beds, Gzhelian): Hüttenkofel, Carnic Alps, Austria. 7 Clypeina jurassica Favre. Star-shaped cross sections (CS) and chain-like oblique sections (OS) of isolated whorls. The cylindrical thallus consists of basket-shaped superposed whorls. The whorls are composed of jointed fertile branches bearing sporangia that almost touch each other, and broadening toward the periphery (arrow). Packstones and wackestones with mass accumulations of Clypeina jurassica are diagnostic of lagoonal inner platform environments with restricted circulation and sometimes also reduced marine salinity (–> Pl. 28/2, Pl. 62/5-6, Pl. 134/5). The species is widely distributed from the Late Kimmeridgian to the base of the Late Berriasian. Late Jurassic: Vrhnica, Slovenia.
Plate 59: Diagnostic Criteria of Dasyclads
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(4) Number and shape of laterals are essential in distinguishing species and genera of dasyclads: Primary branches arise from the main axis in different arrangements, numbers and forms. The primary laterals may branch distally, producing segments of second (Pl. 63/6, Pl. 64/4), third (Palaeodasycladus, Pl. 62/1) or higher order. The number of the segments per lateral and the number of laterals per whorl as well as the shape of the laterals are used as species- and genus-specific criteria. Terms characterizing the shape of the laterals are (a) phloiophore (branches distinctly expanded terminally; genus-specific; e.g. Bornetella; Eovelebitella, Pl. 59/1), (b) vesicular (laterals inflated, straight or slightly
Determination of Dasyclads
bent), (c) pirifere (laterals short, usually straight, inflated, widest diameter at their middle; nearly exclusively as a primary lateral), (d) trichophore (laterals long, pear-shaped, sudden decrease in diameter followed by very thin hairlike forms; restricted to higherorder segments of laterals; Neoteutloporella, Pl. 59/3), (e) acrophore (laterals thin, acute; mainly as higherorder laterals), and (f) umbrelloform (laterals inflated forming an umbrella-like whorl (Clypeina, Pl. 59/7) or separated pear-shaped or lanceolate rays; restricted to primary segments of laterals). (5) Position of the reproductive organs (Fig. 10.15): Gametes (sexual reproductive cells) are produced in
Box 10.3. How to determine dasyclad algae? The table should help in organizing the requisition of data that are necessary to separate dasyclads at least at a generic level. Basic considerations Look at the number of dasyclad sections in your sample. Are there many or just a few? Note that determining genera and taxa requires combining criteria seen in longitudinal, cross, oblique and tangential sections of variously shaped three-dimensional objects. Examine the preservation of the dasyclads. Are the algae well or poorly preserved? Did recrystallization destroy important taxonomic criteria (e.g. arrangement and shape of pores/laterals)? Are the central stem and pores filled with sediment or cement? Growth of cement may make it difficult to recognize the number of the orders of branches. Differentiate the sections into longitudinal, tangential, cross and oblique sections. Select those sections that exhibit a maximum of morphological criteria. Morphological criteria Evaluate the shape of the thallus using Fig. 10.14A. Differentiate between unbranched and branched thalli. Unbranched thalli may be tube-like, pear-shaped or stalked. Both unbranched and branched thalli may be segmented. Typify the subdivision of the thalli (Fig. 10.14B). The external or internal configuration (undulation, annulation, intusannulation) is important in separating genera. Describe the arrangement of the primary branches arising from the main axis (Fig. 10.14C). Are the branches irregularly arranged with respect to the main axis (aspondyl)? Are they regularly arranged like the spokes of a wheel (euspondyl), or do they form regular whorls consisting of clusters of primary branches (metaspondyl)? These criteria serve to distinguish families and genera. Look at the contact of the primary branches with the main axis. Laterals may form different angles with the axis; they may start from a small platform, or two or more branches may have a joint basis. Typify the shape of the branches/pores (Fig. 10.14E). Do the diameter and width of the pores increase distally (phloiophore and vesicular) or decrease distally (acrophore, trichophore)? These criteria are species-specific. Examine the number of higher-order branches arising from the primary branches (Fig. 10.14D). The pattern of distal branching of primary lateral into branches of second, third and higher orders is crucial in distinguishing genera and species. Look for reproductive organs within the central stem or in or between branches (Fig. 10.15). These criteria are important in distinguishing dasyclad groups. Quantitative data Collect measurements that characterize the range and variability of the dimensional criteria. Commonly used in distinguishing species are the diameter of the central cavity of the thallus, the ratio of the outer and inner thallus diameter (total diameter versus diameter of the main axis), the thickness of the ‘wall’ (representing the calcified zone around the central stem), the diameter of primary and other branches, the angle between branches and main axis, and the size of the whorls and the number of branching participating into whorl formation. Relevant references Start with the papers by Berger and Kaever (1992), De Castro (1997) and Bassoullet et al. (1975) in order to understand the terminology and the handling of the morphological criteria. In most cases you will have some idea about the approximate age of your sample. Look at the list given in ‘Basics’ at the end of this paragraph and select the relevant overviews that will lead you to other useful papers.
Distribution of Dasyclads
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Fig. 10.16. Frequency of Jurassic to Cenozoic dasyclad genera. New genera originated predominantly in the Late Jurassic and Paleogene. Data after Bucur (1999). Extinction rates of species were high at the end of the Early and Middle Jurassic, in the Late Aptian and near the Cretaceous-Tertiary boundary. Origination rates of Late Cretaceous species are somewhat controversial in comparison to generic origination rates, as shown by the increase in the number of species in the Cenomanian (Kuss 1994).
the main axis (endospore, e.g. Diplopora, Pl. 59/2), within the laterals, predominantly of primary laterals (cladospore), in cysts that are laterally or terminally attached to the primary laterals (choristospore), and in ray-shaped gametophors (umbrellospore; Acetabularia; Acicularia, Pl. 64/9). In addition, up to 30 quantitative parameters are used in discriminating species (e.g. exterior diameter, diameter of the stem cell, height of a segment, number of whorls per segment, distance between whorls, diameter of branches and pores). The biometrical treatment of these measurements assists in evaluating speciesspecific variations. Classification: Modern dasyclads are represented by only about 11 genera and 37 species (Berger and Kaever 1992), in contrast to ancient dasyclads comprising about 235 genera and up to 830 species (Bucur 1999). For a long time dasyclads were considered as one family (Dasycladaceae), and were promoted to the rank of an order (Dasycladales) comprising several families and
systematic subunits (tribi). In this book the term dasyclads is used as a name for the entire group. Following Deloffre (1988) and Berger and Kaever (1992) the dasyclads are now subdivided into six families comprising the • Seletonellaceae (laterals aspondyl. Reproductive organs endospore, seldom cladospore. Cambrian to Early Cretaceous, predominantly Early Paleozoic. 7 subgroups), • Diploporaceae (metaspondyl, endospore, seldom cladospore. Devonian to Early Cretaceous, predominantly Late Paleozoic and Triassic. 4 subgroups), • Triploporellaceae (euspondyl, cladospore. Devonian to Early Tertiary. 7 subgroups), • Dasycladaceae (euspondyl, choristospore. Jurassic to recent, predominantly Cretaceous and Early Tertiary. 4 subgroups), • Acetabulariaceae (euspondyl, umbrellospore. Jurassic to recent. 3 subgroups). • Beresellaceae (Late Devonian to Early Permian). Text continued on p. 438
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Late Paleozoic Dasyclads
Plate 60 Late Paleozoic Dasyclad Green Algae Dasyclad algae contributed significantly to the formation of extended algal carbonates (grain- and wackestones) and the formation of reef mounds and mud mounds in Carboniferous and Permian shelf environments. The plate displays some globally distributed algal limestones. Baffling effects of algae with small erect cylindrical thalli were important in the formation of Carboniferous mounds (–> 1, 2, 5). Some authors consider these algae as dasyclads (Beresellaceae), from which they differ in a segmentation of the cylinders by horizontal or vertical elements (–> 2); other authors regard them as a separate green algal group (Palaeosiphonocladales), known from the Devonian to the Early Permian and abundant in the Middle Carboniferous. The genera Dvinella, Donezella and Beresella, common in Bashkirian and Moscovian carbonates grew in dense low-lying thickets that baffled stabilized sediment, forming large low-relief mud-rich mounds or banks in relatively quiet normal marine lagoonal settings, landward of grainstone shoal deposits or of Chaetetes buildups (Riding 1979; Davies and Nassichuk 1989). Some of the banks and buildups form hydrocarbon reservoirs. Many Carboniferous and Permian platform and ramp carbonates were formed by the accumulation of dasyclad algae (–> 6) and their reworked and transported fragments (–> 3, 4, 7). Fig. 10.17. Distribution of algae and cyanobacteria on an idealized Carboniferous carbonate platform. Stacheinacean and solenoporacean algae (Pl. 55 and Pl. 56 ) indicate the deepest habitats in the photic zone. Higher up in open-marine settings red algae occur together with tubular palaeosiphonoclad algae. Shallow shoals and lagoonal environments are indicated by abundant dasyclads. Dasyclads associated with palaeosiphonoclads and cyanobacteria mark a semi-restricted very shallow environment. Similar associations occur in restricted nearcoastal environments together with algal cysts (calcispheres, see Sect. 10.2.1.9). Modified from Mamet (1991).
1 Pseudodonezella boundstone. Middle Carboniferous Beresellacean green algae were important in the formation of bindstones and bafflestones as well as of mounds on open-marine carbonate platforms. The algae are characterized by cylindrical ramified and constricted thalli (arrows). Middle Carboniferous: Southern Spain. 2 Beresella boundstone. Beresellids are characterized by a peripheral zone consisting of short, thin, blind-ending cortical branches, separated by ‘clear bands’ (arrow) of a non-perforated cortex and an outer cement layer (–> 5). Mid-shelf facies. Early Permian (Asselian-Sakmarian): Sverdrup Basin, Canadian Archipelago. 3 Coated dasyclad grainstone with Eovelebitella (see Pl. 59/1). The grains exhibit transitions from cortoids to oncoids. Open-marine platform. Early Carboniferous (Viséan): Spain. 4 Epimastopora grainstone. The algae are represented by plate-like fragments exhibiting numerous cylindrical or dumbbellshaped branches. Epimastoporids are widely distributed in Late Carboniferous and Early Permian outer shelf environments. Their assignment to dasyclads is a matter of discussion (Flügel and Flügel-Kahler 1980). Early Permian (Asselian): Carnic Alps, Austria. 5 Beresella herminae Racz. Note continuous calcite layer (arrow) around the thallus, interpreted as carbonate precipitation within the surficial mucilage sheat of the alga. Middle Carboniferous (Valdeteja Formation): Cantabrian Mountains, Northern Spain. 6 Mizzia wackestone. The thalli consist of spherical or cylindrical segments. The branches enlarge in diameter near the surface, producing a hexagonal pattern in tangential sections (arrows). Mizzia limestones were widespread in Middle and Late Permian shelf and back-reef environments of the tropical Paleotethys. Late Permian: Taurus Mountains, southern Turkey. 7 Dasyclad spores. Atractyliopsis carnica Flügel represents the fertile stage of a dasyclad alga that produced spores of various sizes and shapes within spherical sporangial cavities (arrows). Calcified spores occur in rock-building frequency in Late Paleozoic open-marine shelf carbonates. Early Permian: Carnic Alps, Austria. –> 2: Morin et al. 1994; 5: Eichmüller 1985
Plate 60: Late Paleozoic Dasyclads
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All the families except the Beresellaceae are nonseptate. Because the Beresellaceae differ from the typical dasyclads by the horizontal and vertical septation of the thalli and the absence of laterals, some authors consider the group as separate order of green algae. How to determine dasyclad algae? Many people studying the microfacies of carbonate rocks are quite satisfied to attribute algal bioclasts to the dasyclad group, and postulate that the sample indicates deposition in shallow or very shallow warm water of normal salinity. This is a very generalized statement that can be significantly enhanced if the dasyclads are differentiated at a genus and/or species level. Further systematic determinations are also necessary if algal-bearing platform carbonates are to be subdivided stratigraph-
Triassic Dasyclads
ically and correlated. Generic determinations require evaluation of the following criteria: The shape and subdivision of the thalli; aspondyl, euspondyl or metaspondyl arrangement of laterals; the shape of the laterals; and the position of reproductive organs. If these criteria and the extensive stratigraphic overviews of dasyclads are used, nonspecialists will also be able to determine genera and higher systematic groups. A determination of species should be left to specialists. Distribution of Phanerozoic dasyclads (Fig. 10.16): The first dasyclads are known from the Late Cambrian, but the group remains subordinate in contrast to cyanobacterians and other algal groups from the Ordovician to the Devonian. The first diversification took place during the Carboniferous and Permian, a total renewal Text continued on p. 442
Plate 61 Triassic Dasyclad Green Algae No calcareous algae are known from the Early Triassic. Dasyclad algae are common in Tethyan Middle and Late Triassic shallow-marine carbonates. Triassic dasyclads display environmentally controlled distribution patterns that are used in the spatial subdivision of lagoonal platform carbonates (Senowbari-Daryan and Schäfer 1979; Lobitzer et al. 1990). Middle Triassic dasyclads contributed significantly to the accumulation of grain- and packstones. Regional stratigraphical zonations of Tethyan carbonates are based both on algal assemblages and on the time-range of species (Ott 1972). A revised zonation has been worked out by Piros et al. (1994, 2002), who established a detailed zonation for the Middle Triassic and Carnian. 1 Bioclastic grainstone composed of dasyclad algae, aggregate grains (AG), a few solenoporacean algae (S) and corals (C). Dasyclads are Diplopora annulata (Schafhäutl) (D) and Teutloporella herculea (Stoppani) (T). See Bucur and Enos (2001) for a discussion of the taxonomy of Diplopora-like dasyclads. Identification of them requires evaluation of the criteria seen in cross sections (CS) and longitudinal sections (LS). Note the occurrence of several marine cement generations. The sample is a good example of SMF 18 (see Pl. 122). Paleoenvironmental interpretation: Deposition in highenergy environment (indicated by lack of micrite and grain-support fabric), water depth of a few meters (as compared with depth distribution of modern dasyclads and inferred from the occurrence of aggregate grains), warm water (related to the abundance of dasyclads), unstable sand bottom (high primary intergranular porosity, incipient marine cementation). The sample represents lagoonal platform limestones formed within a back-reef area. SMF 18-DASY. Middle Triassic (Wetterstein Limestone, Late Ladinian): Amtmanngalgen, Gesäuse Mountains, Styria, Austria. 2 Dasycladad grainstone. Cross sections of Heteroporella zankli (Ott) characterized by large sporangia arranged in rings. Dasyclads (D) are associated with porostromate cyanobacteria (C). Note high primary intergranular porosity now occluded by isopachous acicular marine-phreatic cements (MPC) and sparry coarse burial cements (CBC). Paleoenvironmental interpretation: High water energy (grainstone fabric, reworked cyanobacteria nodules), water depth a few meters (as compared with the average distribution of modern dasyclads and calcified cyanobacteria), warm water (chloralgal association, see Pl. 105/2), instable sand substrate. The sample comes from outer lagoonal carbonates deposited behind a major reef complex formed by coralline sponges and corals. Late Triassic (Norian to Rhaetian): Cuzzo di Lupo, Palermo Mountains, northwestern Sicily, Italy. 3 Dasyclad grainstone composed of Teutloporella nodosa (Schafhäutl). Diagnostic criteria are shown in longitudinal (LS) and oblique sections. The long thalli are annulated. The aspondyle branches become thinner toward the outside, forming a hairlike structure (white arrow). Note distinct circumgranular isopachous cements (black arrows) consisting of acicular crystals. These cements were formed in marine phreatic environments. Paleoenvironmental interpretation: Shallow-marine warm-water carbonate, formed in high-energy areas. The high and still open interskeletal porosity (P) can be explained by strong meteoric phreatic diagenesis. The sample comes from platform shoals formed by a parautochthonous accumulation of algae. Teutloporella is a facies-diagnostic fossil; it characterizes subtidal lagoonal deposits as well as proximal reef detritus areas (Ott 1967; Piros et al. 1994). Stratigraphic range of the species: Ladinian and Early Carnian. Middle Triassic (Late Ladinian): Calabria, Southern Apennines, Italy.
Plate 61: Triassic Dasyclads
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Jurassic Dasyclads
Plate 62 Jurassic Dasyclad Green Algae Biozonations developed for Tethyan Jurassic and Cretaceous shallow-marine carbonates are based predominantly on dasyclads and benthic foraminifera. These zonations use combinations of various stratigraphical concepts. Some zones correspond to taxon range and concurrent range zones, but many ‘coenozones’ are assemblage or abundance zones characterized by the association of several taxa or by a maximum abundance of specific taxa. Coenozone boundaries, therefore, are not necessarily synchronous time lines. The Early and Middle Jurassic represent periods of decline and renewal of dasyclads. Common Early Jurassic and Middle Jurassic dasyclads are Palaeodasycladus (–> 1) and Selliporella (–> 4). The Late Jurassic, particularly the Kimmeridgian and Tithonian, was characterized by highly diverse and widely distributed dasyclad associations. Common dasyclads in Late Jurassic-Early Cretaceous times were Salpingoporella (–> 2; Pl. 63/7), Clypeina (–> 5, 6), Cylindroporella (Pl. 63/8) and Triploporella. Dasyclads occur in rock-building abundance, as exemplified by Clypeina limestones (–> 5, 6) or limestones with Campbelliella (–> 2, 3) in lagoonal or protected shelf environments. 1 Palaeodasycladus mediterraneus Pia. Oblique (OS) and cross sections. The alga was common in Tethyan Early and Middle Liassic (Sinemurian to Domerian) intertidal lagoonal and in outer shelf settings (Sokac 2001) and is used in the biozonation of Liassic platform carbonates. It characterizes a Late Sinemurian acme zone. Elongated club-shaped thalli, up to third-order branches diverge with angles of about 45° from the main axis. The sediment consists of poorly sorted aggregate grains, peloids and micritized algae, sometimes forming nuclei for micrite oncoids (arrows). Note the sparry patches within the oncoids, indicating the contribution of microbial networks to the formation of ‘fenestral oncoids’. Early Jurassic (Pantokrator limestone, Sinemurian): Korfu Island, Greece. 2 Salpingoporella annulata Carozzi. Oblique (OS) and cross sections (CS). Cylindrical annulated thallus. Pores vertically flattened, arranged in whorls (arrow). CA: Transversal-oblique section of Campbelliella. Open-marine isolated platform. Late Jurassic (Tithonian): Sulzfluh Range, Graubünden, Switzerland. 3 Campbelliella striata (Carozzi). Arrows point to different sections. These microfossils were formerly regarded as incertae sedis, pteropod gastropods, large aberrant tintinnids, or pallets of boring teredinid bivalves. They are now recognized as dasyclad algae. The calcareous skeleton consists of a succession of articles interpenetrated into each other. Each funnel-shaped article (A) contains the pores of a single whorl (De Castro 1993). Campbelliella occurs in abundance with small benthic foraminifera and Favreina in restricted lagoonal platform carbonates. Late Jurassic (Tithonian): Sulzfluh Range, Graubünden, Switzerland. 4 Selliporella donzellii Sartoni and Crescenti. Oblique longitudinal section. Note double whorls (arrows) with pores thickening outward (phloiophore). Range of the species: Bajocian to Late Bathonian. Late Middle Jurassic: Southern Apennines, Italy. 5 Clypeina jurassica Favre. Cross sections. Fertile whorls shaped like a wide basket and composed of broadening sporangial chambers. Note (dark) pores in the center of these chambers. Species range: Late Jurassic to Berriasian, common in Kimmeridgian. Late Jurassic: Vrhnica, A B Slovenia. 6 Clypeina jurassica Favre. Isolated whorl fragments. Late Fig. 10.18. Reconstructions of Late Jurassic dasyclads. Jurassic: Sulzfluh Range, Graubünden, Switzerland. A: Clypeina jurassica Favre. After Pia (1920). B: Campbelliella striata (Carozzi). After De Castro (1993). –> 1: Flügel 1983
Plate 62: Jurassic Dasyclads
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at a species level during the Middle Triassic (Flügel 1985). Late Jurassic and Early Cretaceous platforms were characterized by highly diversified associations, both at a genus and species level. Starting in the Aptian, dasyclads decline significantly at a generic level during the Late Cretaceous before reaching high diversities and global distribution again in the Early Tertiary (Fig. 10.16). Critical lists of dasyclad species are available for the Permian and Triassic (Granier and Grgasovic 2001), Jurassic and Cretaceous (Bassoullet et al. 1978; Deloffre 1993), and Tertiary (Deloffre and Genot 1982; Deloffre and Granier 1992). Plates 60 to 64 show some typical and common dasyclads occurring in Carboniferous to Paleocene shallow-marine carbonates. Environmental constraints: Modern dasyclads occur predominantly in tropical and subtropical environments but are also found in temperate waters, e.g. the Mediterranean. Their distribution is controlled by water temperature (warm waters, mostly within the 20° isoclyne), substrate (sand and mud), salinity (normal
Significance of Dasyclads
marine; only a few modern and some Tertiary taxa became adapted to hypersaline and brackish waters), water energy (low, generally subtidal) and depth (below low tide down to about 30 m, commonly < 5 meter). Dasyclads are found in open lagoons, protected but open bays and in the shelter of ridges. The ecological requirements of post-Paleozoic dasyclads are regarded as similar to those of their living descendants (Elliott 1978, 1981). Deviations of distributional patterns of Permian and Triassic dasyclads from the recent shallow-water model have been explained by substrate control (Flügel 1981). Some algal remains attributed to ‘marine’ dasyclads have been reported from Early Permian lacustrine sediments and explained by wind drift of cysts from marine environments to terrestrial sites (Schneider and Gebhardt 1992; Pl. 131/4). Significance of dasyclad algae Sedimentological significance: Since the Carboniferous, the accumulation of dasyclad bioclasts contributed significantly to the formation of shallow-marine,
Plate 63 Early Cretaceous Dasyclad Green Algae Similar to the Late Jurassic, algal carbonates were also essential in the formation of Early Cretaceous platforms. More than 150 species have been described from Early Cretaceous deposits. The greatest wealth of species occurs in the Barremian. Berriasian-Valanginian and Barremian-Aptian assemblages differ in species composition and abundance. Dasylad assemblages allow regional biozonations (e.g. Bucur 1994; Bodrogi et al 1994). The large number of species decreased dramatically from the Aptian to the Albian. Dasyclads assist in differenting the paleoenvironments of various parts of platforms (Peybernes 1979; Fig. 10.19), and in evaluating cyclostratigraphic units (Schindler and Conrad 1994). Some widely distributed genera in the Tethyan region are shown on this plate. 1 Bakalovaella elitzae (Bakalova). Early Cretaceous (Late Barremian-Early Aptian): Southern Romania. 2 Salpingoporella ex. gr. muehlbergi (Lorenz). Oblique longitudinal section (OS) and transverse section (TS). Thallus cylindrical. Only primary branches are arranged perpendicularly to the main axis in euspondyl whorls (arrows). Salpingoporella is common in back-reef and open-lagoonal sediments. Early Cretaceous (Barremian-Aptian): ResitaMoldova Noua zone, southwestern Romania. 3 Salpingoporella ex gr. melitae Radoicic. Oblique longitudinal section. Elongated cylindrical thallus. Same locality as –> 2. 4 Pseudocymopolia jurassica (Dragastan). The oblique longitudinal section through two segments (S) exhibits primary and densely spaced secondary branches (arrow). Note the distinct annulation of the thallus. Same locality as –> 2. 5 Neomeris cretacea Steinmann. Arrows point to the sporangia characteristic of the species. Early Cretaceous: Southern Romania. 6 Pseudoactinoporella fragilis Conrad. Transverse section. Simple branches increasing in size outwardly (phloiophore). The distributional pattern of the branches indicates that the branches are arranged obliquely to the central stem. Early Cretaceous (Barremian): Romania. 7 Algal wackestone with Salpingoporella dinarica Radoicic. This facies is common in restricted parts of Aptian carbonate platforms. The association of the species with ostracods, charophytes and specific foraminifera in the Valanginian indicates pond environments in strongly protected lagoons. Early Cretaceous (Valanginian): Apennines, southern Italy. 8 Shallow-marine algal-foraminiferal limestone with miliolid foraminifera and dasyclad algae (Cylindroporella, note the spherical sporangia growing attached to the main axis, choristospore type; center). Inner shallow platform. Early Cretaceous: Subsurface, United Arabian Emirates.
Plate 63: Early Cretaceous Dasyclads
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sand-sized carbonates on platforms and ramps. Many dasyclad limestones (mostly grainstones and packstones) were formed in open-marine lagoonal or backreef environments. Examples are Permian Mizzia limestones (Pl. 60/6), Middle Triassic Diplopora limestones (Pl. 61/1), Liassic Palaeodasycladus limestones (Pl. 62/ 1), Late Jurassic Clypeina limestones (Pl. 62/5, 6), and Early Cretaceous Salpingoporella limestones (Pl. 63/ 7). Late Paleozoic and Mesozoic dasyclads contributed
Significance of Dasyclads
to the formation of biogenic mounds (e.g. Late Carboniferous Anthracoporella mounds, Pl. 59/6; Late Jurassic Neoteutloporella biostromes, Pl. 59/3). Recognition of shallow-marine depositional sites: Different algal groups prefer different depositional sites. Fig. 10.17 and Fig. 10.19 give a rough picture of the environmental distribution of Carboniferous dasyclads and other algal groups.
Plate 64 Late Cretaceous and Paleocene Calcareous Green and Red Algae Dasyclad algae suffered a significant decline during the Late Cretaceous. The number of genera per stage decreased from about 40 in the Early Cretaceous to less than 20 in the Late Cretaceous. A similar drop is shown by the number of species from the Cenomanian to the Campanian. Maastrichtian dasyclad assemblages exhibit slightly higher species numbers (up to 15). Late Cretaceous dasyclad species are of value in establishing composite biozonations (Kuss and Malchus 1989). Early Tertiary dasyclads are valuable time markers. The Early Cenozoic was a time of high dasyclad species diversity connected with a renewal of the algal associations. Paleocene and Eocene species numbers (81 and 95 respectively) surpass those of the Early Cretaceous. At the end of the Eocene more than 90% of the Paleocene-Eocene species had disappeared. The Neogene represents a marked decline for the dasyclads. 20 species were recorded from the Miocene, 8 from the Pliocene, only one species appears to have survived in recent assemblages, which contain about 35 species, predominantly of the group Acetabulariaceae. 1
Dasyclad green alga: Neomeris (Larvaria) sp. Tangential/axial longitudinal section. Black arrows point to distinctly differentiated whorls. White arrows indicate the dasyclad alga Aciculella. Paleocene (Thanetian): Mariazell, Styria, Austria. 2 Dasyclad green alga: Trinocladus perplexus Elliott. Longitudinal section. Shallow subtidal open platform. Late Paleocene: Priglitz, Lower Austria. 3 Udoteacean green alga: Halimeda nana Pia. Longitudinal section. Note the distinct differentiation of axial medulla (M) and cortex (C). Arrows point to filaments. Open platform. Paleocene: Priglitz, Lower Austria. 4 Dasyclad green alga: Cymopolia elongata (Defrance) Munier-Chalmas. Longitudinal section. Differences in size of the thalli and number of whorls allow low-energetic, moderate or high-energetic (this figure) environments to be differentiated. Same locality as –> 3. 5 Distichoplax biserialis (Dietrich). A widely distributed microproblematicum, restricted to Paleocene and Early Eocene carbonates. E: Echinoid spine. Paleocene (Thanetian): Lupina, Western Carpathians, Slovakia. 6 Dasyclad green alga: Cymopolia inflataramosa Segonzac. Oblique cross section. Diagnostic criteria are broadened primary branches (black arrows) and thin short secondary branches (white arrow). Paleocene (Late Thanetian): Hricovske Podhradie, Slovakia. 7 Composite crust consisting of red algae (top: Archaeolithothamnium, A; below: Pseudolithothamnium album, PA) and sessile foraminifera (Planorbulina, P). Deeper subtidal environment. Late Maastrichtian: Hochschwab, Styria, Austria. 8 Dasyclad green alga: Broeckella belgica Munier-Chalmas. Cross-section. Diagnostic criteria is the broadening of the primary branches (arrows). The species is an index fossil (Danian and Early Thanetian). Common in lagoonal and backreef sediments. Paleocene (Thanetian): Mooshuben, Styria, Austria. 9 Dasycladalean green alga: Acicularia cf. pavantina d’Achiarc. Note cross sections of ampullae. Same locality as –> 8. 10 Squamariacean red algae (Peyssonelia, P) and coralline red algae (Amphiroa, A; Lithophyllum, L). Criteria of squamariacean algae: Encrusting aragonitic thalli consisting of a basal layer (hypothallium, HT) from which vertical or arcuate rows of cells arise (perithallium, PT). Note the difference in the preservation of aragonitic squamariacean and High-Mg calcite corallinacean algae. Typical association of high energy shallow-marine environments. Late Maastrichtian: Hochschwab, Austria. 11 Articulated coralline red alga: Amphiroa sp. Longitudinal section. Criteria: Medullary segments (S) consisting of long cells. Distinctly layered cortical tissue (CT). Same locality as –> 10. 12 Crustose coralline red alga: Sporolithon sp. Note the linear arrangement of conceptacles (C). Late Paleocene: Priglitz, Lower Austria. –> 1–11: Courtesy of H. Tragelehn (Köln).
Plate 64: Late Cretaceous and Paleocene Dasyclads
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Dasyclad Associations
Fig. 10.19. Distributional patterns of dasyclad associations on Early Cretaceous carbonate platforms in the French and Spanish Pyrenees Mountains. Species composition of the association differs significantly in the various parts of the tidal and shallow subtidal shelf environments. A: Berriasian and Valanginian. B: Barremian to Aptian (Urgonian facies). Based on Peybernes (1979). See Masse (1993) for a discussion of facies development during time and the biostratigraphical value of dasyclad species associations. Dotted lines indicate rare occurrences.
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Basics Dasyclads
Similar to modern dasyclads, ancient dasyclads exhibit recurrent species assemblages that preferred particular depositional environments on shallow, welllighted shelves. The distribution of these assemblages follows regular patterns, at least for Mesozoic and Cenozoic dasyclads, and can be successfully used in recognizing and subdividing facies belts of platforms and ramps. Fig. 10.19 shows an example from Early Cretaceous platforms. Biozonations: Many dasyclad genera and species have long stratigraphical ranges, comprising more than one stage (Bucur 1999). However, notable exceptions exist that allow platform carbonates to be dated and subdivided. Platform carbonates are often poor in guide fossils. The only abundant fossils in these limestones are calcareous algae and foraminifera. Both groups are prone to strong facies controls that reflect the change of the major environmental constraints (e.g. sea-level fluctuations, clastic input) over time. These changes are recorded by vertical successions of rock units yielding specific dasyclads and other thin-section fossils that allow Mesozoic platform carbonates to be subdivided into variously defined ‘biozones’. These biozones are valuable stratigraphic tools in regional correlations, but their boundaries are often certainly not synchronous. Dasyclads play an important role in defining biozones. First recognized in 1920 by Julius Pia from Vienna, species ranges of dasyclads allow relatively short stratigraphic intervals to be differentiated for time intervals characterized by high origination rates (e.g. Middle and Late Triassic, parts of the Early Cretaceous, Paleocene and Eocene). Basics: Dasyclad green algae Barattolo, F. (1991): Mesozoic and Cenozoic marine benthic calcareous algae with particular regard to Mesozoic Dasycladalenas. –In: Riding, R. (ed.): Calcareous algae and stromatolites. – 504-540, Berlin (Springer) Bassoullet, J.-P., Bernier, P., Deloffre, R., Genot, P., Jaffrezo, M., Poignant, A.F., Segonzac, G. (1975): Réflexions sur la systématique des Dasycladales fossiles – Étude critique de la terminologie et importance relative des critères de classification. – Geobios, 8, 259-290 Bassoullet, J.P., Bernier, P., Conrad, M.A., Deloffre, R., Jaffrezo, M. (1978): Les algues dasycladales du Jurassique et du Crétacé. – Geobios, Mem. Spec., 2, 330 pp. Berger, S., Kaever, M.J. (1992): Dasycladales: an illustrated monograph of a fascinating algal order. – 274 pp., Stuttgart (Thieme) Bucur, I.I. (1999): Stratigraphic significance of some skeletal algae (Dasycladales, Caulerpales) of the Phanerozoic. – Palaeopelagos Special Publication, 2, 53-104 De Castro, P. (1997): Introduzione allo studio di sezioni sottili delle Dasicladali fossili (An approach to thin-section study of fossil Dasycladales). – Quaderni dell’Accademia Pontaniana Napoli, 22, 261 pp.
Deloffre, R. (1988): Nouvelle taxonomie des algues dasycladales. – Bull. Centres Rech. Explor.-Prod. ElfAquitaine, 12, 165-217 Deloffre, R. (1993): Inventaire critique des algues dasycladales fossiles. 2. partie – Les algues dasycladales du Jurassique et du Crétacé. –Revue de Paléobiologie, 12, 19-65 Deloffre, R., Genot, P. (1982): Les algues dasycladales du Cénozoique. – Mem. Centres Rech. Explor. Prod. ElfAquitaine, 4, 247 pp. Deloffre, R., Granier, B. (1992): Inventaire critique des algues dasycladales fossiles. – Revue de Paléobiologie, 11, 331356 Elliott, G.F. (1968): Permian to Paleocene calcareous algae (Dasycladaceae) of the Middle East. – British Mus. Nat. Hist. Bull., Geol. Suppl., 4, 111 pp. Elliott, G.F. (1978): Ecologic significance of post-Palaeozoic green algae. – Geological Magazine, 115, 437- 442 Elliott, G.F. (1981): The Tethyan dispersal of some chlorophyte algae subsequent to the Palaeozoic. – Palaeogeogr., Palaeoclimat., Palaeoecol., 32, 341-358 Flügel, E. (1985): Diversity and environments of Permian and Triassic dasycladacean algae. – In: Toomey, D.F., Nitecki, M.H. (eds.): Paleoalgology. – 344-351, Berlin (Springer) Genot, P. (1991): Cenozoic and recent dasycladales. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 131145, Berlin (Springer) Granier, B., Deloffre, R. (1993): Inventaire critique des algues dasycladales fossiles. 2. partie - les algues dasycladales du Jurassique et du Crétacé. – Revue de Paléobiologie, 12, 19-65 Granier, B., Grgasovic, T. (2000): Les algues dasycladales du Permien et du Trias. Nouvelle tentative d’inventaire bibliographique, géographique et de stratigraphique. – Geologica Croatica, 53, 1-197 Jaffrezo, M. (1973): Essai d’inventaire bibliographique des algues dasycladacées du Jurassique et du Cretacé inferieur. – Geobios, 6, 71-99 Pia, J. (1920): Die Siphoneae verticillatae vom Karbon bis zur Kreide. – Abhandlungen der zoologisch-botanischen Gesellschaft Wien, 11, 1-263 Further reading: K 129. See also Basics for 10.2.1.
10.2.1.8 Charophyta: Fresh-Water and Brackish-Water Algae Definition and morphology: Modern charophytes are regarded as a separate group of macrophytic algae or as a group of green algae. They live submersed and occur predominantly in shallow oligotrophic freshwater environments of lakes and rivers as well as in brackish waters of marginal seas. The erect and branched thallus is divided into a regular succession of nodes, with whorls of small branches, and internodes. The only calcified parts are the reproductive organs, developing from the nodes during fertile stages of the plant, and vegetative stem nodes and internodes. The female re-
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Plate 65
Charophyta
Charophyta
Charophyta are macrophytic green algae. The plant attains a size ranging from several centimeters to as much as a meter. The erect and branched thallus is divided into a regular succession of nodes with whorls of small branches, and internodes. The reproductive organs develop from the nodes during fertile stages. Ancient charophyte fossils are recorded by the female gametangia (oogonia) and by internode parts (–> 1). The oogonium consists of spiral cells (Fig. 10.20). These cells are partly calcified (Low- or High-Mg calcite). Only the calcified part of the oogonium (called gyrogonite) is preserved as a fossil. Many gyrogonites are elliptical in shape and 0.5 to 1.0 mm in size. Additional calcareous plates can form an envelope (utriculus) for the gyrogonites. The taxonomy of fossil charophytes is based on isolated oogonia. Generic differentiation of gyrogonites in thin sections is rarely possible and requires longitudinal sections exhibiting an apex, basal pore and basal plate. Charophytes have been known since the Silurian and became diversified in the Devonian and during the Late Jurassic, Early Cretaceous and Tertiary. Valuable charophyte-based biozonations have been established for the Late Mesozoic and the Tertiary. Fig. 10.20. Charophyta and their calcified parts. A: Charophyte plant. The erect branched thallus is divided into a regular succession of nodes, with whorls of small branches, and internodes. B: Calcification takes place at the nodes only with the elliptical female reproductive organ (oogonium) and near the stem nodes and internodes. C: Oogonium consisting of the internal egg cell, surrounded by spiral cells whose inner part is calcified (black). This part, called the gyrogonite, is preserved as a fossil. D: High-level systematic units are differentiated according to gyrogonite morphology and sculpture. From left to right: Sycidiales (Late Silurian to Early Carboniferous), Trochiliscales (Late Silurian to Permian), Charales (Devonian to Recent). A, B after Wray (1977). C, D after Schudack (1993).
1 Charophyte packstone. Cross sections (white arrows) and longitudinal section (black arrow) of internodia. The algae occur in association with thin-shelled ostracods. Brackish-water facies. LMF 7. Lowermost Cretaceous (Upper Münden beds, lowermost Berriasian): Lower Saxony, Germany. 2 Charophyte gyrogonites (G) and internodia (IN). Fine skeletal elements in the matrix are charophyte debris. Late Cretaceous (Turonian): Divaca, Slovenia. 3 Lacustrine charophyte limestone. Cross sections and oblique sections of internode parts of Characeae. Spinose surface structure (arrows) is caused by internode rips. Black grains are reworked micrite clasts (MC). White areas are open pores (P). Carbonate sands near green algal bioherms. Late Tertiary (Miocene): Marktoffingen, Ries, Southern Germany. 4 Gyrogonite. Cross section. The interior was originally occupied by the egg cell. The ‘wall’ consists of the calcified parts (calcine) of the spiral cells (SC) surrounding the egg cell. Early Cretaceous: Subsurface, southern Bavaria, Germany. 5 Gyrogonite. Oblique tangential section displaying the orientation of the spiral cells. Lowermost Cretaceous (Purbeck facies, Berriasian): Subsurface, Kinsau, southern Germany. 6 Gyrogonite. Lowermost Cretaceous (Purbeck facies, Berriasian): Same locality as –> 5. 7 Wackestone with gyrogonites. LMF 7. Same locality as –> 2. 8 Sycidium. Middle and Upper Devonian lagoonal carbonates can yield microfossils resembling charophyte oogonia and attributed to the genera Sycidium, Karpinskya or Trochiliscus. Middle Devonian: Eifel, Germany. –> 1, 6: Courtesy of M. Schudack (Berlin).
Plate 65: Charophyta
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Calcispheres and Algal Cysts
production organ (oogonium) is elliptical in shape, 0.2– 1.0 mm in size, and exhibits a wall composed of variously arranged tubular cells. The spiral or vertical arrangement of the enveloping cells is recorded in the ‘gyrogonites’ (Fig. 10.20) corresponding to the calcified parts of the oogonium and representing the essential criteria for the classification of fossil charophytes. Environment: Today the group is restricted to fresh and brackish waters (lakes, rivers and brackish marginal seas). Some species also occur in salt lakes. Charophytes are found in shallow lakes, predominantly in water depths of a few meters only, often exhibiting distinct shore-parallel distributional patterns. Modern charophytes are abundant in warm temperate climates but have a wide geographical range, and also occur in tropical and arctic settings. Major controlling factors on the distribution of charophyta are salinity, the oxygen content and chemistry of the waters as well as substrate type, salinity, nutrient supply and substrate type. Because oogonia can be transported by water and wind, the occurrence of abundant fossil oogonia, therefore, must not necessarily reflect biocoenoses. Many ancient charophytes are found in association with nonmarine or brackish water biota. Oogonia found in marine sediments are usually interpreted as allochthonous. However, Devonian forms (e.g. Sycidium, Pl. 65/8) are known from shallow-marine open shelf environments.
Plate 66
Thin sections: Taxonomic studies are predominantly based on isolated gyrogonites. Differentiating genera in thin sections is difficult and would need longitudinal sections displaying the apical and the basal part of the gyrogonites. The calcified spiral cells of the gyrogonites exhibit distinct microstructures (Schudack 1987, 1993). The calcified vegetative internodes of the charophycean stems correspond to cylindrical segments making up a hollow central channel. Cross sections exhibit an axial cavity surrounded by a peripheral zone of small, densely arranged circular structures corresponding to sections of external longitudinal grooves. Distribution: Modern charophytes represent a relict group comprising only six genera. Charophytes have been known since the Silurian. Important radiations took place during the Devonian, Late Jurassic and Early Cretaceous (‘Purbeck’ and ‘Wealden’ facies of Western Europe: Schudack 1993). Gyrogonites are useful in correlations between marine and nonmarine sediments; they are important zone fossils for the Early Cretaceous and Early Tertiary time interval. Paleoenvironmental significance: Charophyta are used in evaluating paleoclimate, paleolatitudes and paleoenvironmental constraints. Fossil gyrogonites are important paleosalinity indicators. They are often associated with nonmarine and brackish-water faunas (e.g.
Paleozoic and Mesozoic Calcispheres
Calcispheres is a term referring to spherical or egg-shaped microfossils which have been regarded for a long time as microproblematica with respect to their systematic position. Most authors now agree that most of these fossils are algal cysts, some perhaps also planktonic algae. Devonian and Carboniferous calcispheres are abundant in restricted and semi-restricted lagoonal environments and in back-reef settings. Most of the Jurassic and Cretaceous calcispheres occurring in pelagic limestones, and often described in thin sections as pithonellids or cadosinids represent calcareous dinoflagellate algae. 1 Paleozoic calcispheres. Peloidal calcisphere-foraminiferal grainstone. The sediment consists of calcispheres, peloids, small intraclasts and some endothyrid foraminifera. The calcispheres exhibit different shapes and different wall structures. These differences are used for taxonomic differentiations. This microfacies is common in Early Carboniferous inner ramp settings. SMF 16-NON-LAMINATED. Early Carboniferous (Viséan): Basin of Dinant, Belgium. 2 Paleozoic calcispheres. Variously sized thin-walled calcispheres. Note the distinct difference in the size of calcispheres and peloids. Early Carboniferous (Late Viséan): Nassfeld, Carnic Alps, Austria. 3 Paleozoic calcispheres. Radiosphera characterized by a radially structured wall. Back-reef facies. Late Devonian (Frasnian): Holy Cross Mountains, Poland. 4 Mesozoic calcispheres. Pithonella, a calcareous dinoflagellate. The SEM photograph shows the regular pattern of the calcite crystals forming the shell surface and the circular exterior pore. Late Cretaceous (Turonian): Switzerland. 5 Mesozoic calcispheres. Pelagic Pithonella packstone. Open-marine deep shelf environment. SMF 3. Late Cretaceous (Seewen limestone, Turonian): Switzerland. 6 Mesozoic calcispheres. Association of calcispheres and pelagic foraminifera (Globotruncana). SMF 3. Same locality as –> 5. 7 Mesozoic calcispheres. Strongly burrowed calcisphere packstone. Late Cretaceous (Pläner limestone, Cenomanian): Teutoburger Wald, Westphalia, Germany.
Plate 66: Paleozoic and Mesozoic Calcispheres
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ostracods, serpulids, ostreid pelecypods), but have also been reported from marine strata. Marine occurrences were explained by (a) transport of charophytes from coastal areas to the sea, (b) reworking of nonmarine beds with charophytes in connection with transgressions, or (c) adaption of some taxa to normal-marine conditions (e.g. some Devonian genera: Racki and Racki 1989). Common microfacies types associated with charophyte-bearing carbonates are ostracod wackestones and mudstones, lime mudstones with cyanobacteria and shell fragments, wackestones and grainstones with charophytes and a mixture of ooids, lithoclasts and skeletal grains, and carbonate breccia. Basics: Charophyta Feist, M., Grambard-Fessard, M. (1991): The genus concept in Charophyta: evidence from Paleozoic to recent. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 189203, Berlin (Springer) Mojon, P.O., Strasser, A. (1987): Microfacies, sédimentologie et micropaléontologie du Purbeckienne de Bienne (Jura suisse occidental). – Eclogae geol. Helvetiae, 80, 37-58 Racki, G., Racki, M. (1989): Ecology of the Devonian charophyte algae from the Holy Cross Mts. – Acta Geologica Polonica, 31, 213-222 Schudack, M. (1987): Charophytenflora and fazielle Entwicklung der Grenzschichten mariner Jura/Wealden in den nordwestlichen Iberischen Ketten (mit Vergleich zu Asturien und Kantabrien). – Palaeontographica, B, 204, 1-180 Schudack, M. E. (1993): Die Charophyten in Oberjura und Unterkreide Westeuropas. Mit einer phylogenetischen Analyse der Gesamtgruppe. – Berliner Geowissenschaftliche Abhandlungen, E, 8, 209 pp. Further reading: K129
10.2.1.9 ‘Calcispheres’ and Algal Cysts Many Paleozoic and Mesozoic shelf limestones yield small-sized (diameter commonly < 500 µm), hollow spherical microfossils exhibiting calcitic ‘walls’. Many of these fossils, often designated as ‘calcispheres’, are interpreted as algal remains, particularly as algal cysts. True algal spores are very rare in thin sections but red algal spores have been recorded from Jurassic limestones (Fay and Dirsche 1975). Paleozoic calcispheres occur in shallow-marine platform and ramp carbonates. Jurassic and Cretaceous spherical microfossils are predominantly found in pelagic limestones. Dasyclad cysts In living, calcifying dasyclads cysts are formed in reproductive branches or chambers. Some dasyclads are poorly calcified except around the cysts (e.g. Hali-
Calcispheres
coryne). Isolated cysts and aggregates of cysts can be common in limestones. Modern dasyclad cysts are represented by calcareous, thin-walled spherical hollow bodies (Marszalek 1975; Berger and Kaever 1992). Morphologically comparable spheres and aggregates of spheres and egg-shaped bodies, ranging in size between 100 and 500 µm, are known from Permian, Mesozoic and Paleogene carbonates (e.g. Flügel 1966; Elliott 1986; Pl. 60/7). Devonian and Carboniferous calcispheres Hollow spheres with simple or differentiated calcitic walls are common constituents of Middle and Late Devonian restricted lagoonal and back-reef carbonates as well as early Carboniferous ramps. Sizes between 100 and 500 µm. The spheres have been taxonomically differentiated according to wall structure, ornamentation and size variation (Pl. 66/1-3). The fossils have been regarded as dasyclad algae (Marszalek 1975) or as planktonic green algae (Volvocales: Kazmierczak 1975), but also as single-chambered foraminifera. The spheres are commonly associated with peloids, amphiporid stromatoporoids (Devonian) or smaller foraminifera (Carboniferous). Paleozoic calcispheres can be used as indicators of paleosalinity. Mesozoic and Cenozoic ‘calcispheres’: Calcified dinoflagellate algae Thin sections and SEM studies of Late Jurassic, Cretaceous and Tertiary limestones reveal hollow spherical and ovoid microfossils in rock-building abundance (Pl. 66/5), often associated with pelagic foraminifera, calpionellids and radiolarians. The sizes of the spheres range between about 10 and 100 µm, usually about 40 µm). The spherical or egg-shaped bodies are characterized by differentiated calcitic walls, with or without openings and pores. Differences in wall microstructure have been used to distinguish genera (e.g. Pithonella, Stomiosphaera, Cadosina, Calcisphaerula; Rehanek and Cecca 1993). SEM studies indicate that many taxa separated by thin-section criteria represent diagenetic forms only and that other ‘taxa’ refer to the high morphological variation of these fossils. Many single-layered Mesozoic calcispheres are cysts of planktonic algae (Calcidinoflagellata: see Keupp 1981). Calcareous dinoflagellate cysts have existed probably since the Triassic. Mesozoic calcispheres exhibit a wide climatic range and were common in temperate regions. They have been found predominantly in sediments of the deeper shelf to slope and basinal settings, sometimes in inner shelf environments. These calcispheres are abundant in association with other planktonic algae and occur in pelagic upper slope
Introduction to Foraminifera
and basinal carbonates as well as in outer shelf carbonates of low and mid latitudinal settings. The fossils can be used as bathymetric indicators of shelf margin and slope environments and in the stratigraphic subdivision of Late Jurassic to mid-Cretaceous Tethyan and subboreal pelagic carbonates (Borza 1969; Bucur 1992). Basics: Calcispheres and algal cysts Dasyclad cysts Elliott, G.F. (1986): Isolated cysts in the Palaeocene-Lower Eocene of Kurdistan. – Palaeontology, 29, 739-741 Flügel, E. (1966): Algen aus dem Perm der Karnischen Alpen. – Carinthia II, Sonderheft, 25, 3-76 Marszalek, D.S. (1975): Calcisphere ultrastructure and skeletal aragonite from the alga Acetabularia antillana. – J. Sed. Petrol., 45, 266-271 Paleozoic calcispheres Flügel, E. and Hötzl, H. (1971): Foraminiferen, Calcisphären und Kalkalgen aus dem Schwelmer Kalk (Givet) von Letmathe im Sauerland. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 137, 358-395 Kazmierczak, J. (1975): Colonial Volvocales (Chlorophyta) from the Upper Devonian of Poland and their paleogeographical significance. – Acta Palaeont. Polonica, 29, 73-85 Mesozoic calcispheres Borza, K. (1969): Die Mikrofazies und Mikrofossilien des Oberjuras und der Unterkreide der Klippenzone der Westkarpaten. – 301 pp., Bratislava (Slov. Akad. Znanosti) Keupp, H. (1981): Die kalkigen Dinoflagellaten-Zysten der borealen Unter-Kreide (Unter-Hauterivium bis UnterAlbium). – Facies, 5, 1-190 Keupp, H. (1991): Fossil calcareous dinoflagellate cysts. – In: Riding, R. (ed.): Calcareous algae and stromatolites. – 267-286, Berlin (Springer) Rehanek, J., Cecca, F. (1993): Calcareous dinoflagellate cysts, biostratigraphy in Upper Kimmeridgian-Lower Tithonian pelagic limestones of Marches Apennines (Central Italy). – Revue de Micropaléontologie, 36, 143-163, Villain, J.M. (1977): Les Calcisphaerulidae: Architectures, calcification de la paroi et phylogenèse. – Palaeontographica, A, 159, 139-177 Further reading: K139, K140
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the class Rhizopoda. Similar to calcareous algae, foraminifera are of prime importance in microfacies analyses. Foraminifera provide time markers for biozonations of shallow and deep marine carbonates, are excellent environmental proxies and permit ancient depositional systems to be reconstructed. The shells of benthic foraminifera have been major constituents of shelf carbonates at least since the Late Paleozoic and those of planktonic foraminifera formed pelagic limestones from the Late Mesozoic on. Foraminiferal shelf limestones are important reservoir rocks for hydrocarbons. Most foraminiferal shells typically range from 0.1 to 1 mm in size, but some groups developed rather large tests. For study purposes a distinction is made between smaller foraminifera and larger foraminifera (larger than about 0.6 mm). Smaller foraminifera with test sizes of approximately 100 µm to 600 µm form the majority of benthic foraminifera. They are usually identified by their external characters, but can be also typified in thin sections. Larger foraminifera are usually large enough that geologists can spot them in the field. Their tests have complex internal structures that need to be studied in thin sections. The two major groups of foraminifera are benthic, living in or on sediments on the sea floor, and planktonic, living in the upper 100 m of the oceans.
10.2.2.1 Foraminifera
Morphology of foraminifera The living cell is made up of protoplasma. Extensions of the protoplasma (pseudopodia) protrude through a single foramen or through minute pores in the test. Pseudopodia assist in locomotion, capturing of food, and removal of waste products. The cell is enclosed by a test consisting of a single chamber or several chambers each connected by one or several openings. Most foraminifera are multichambered. The chambers are arranged in whorls. They may be uniserial (forming a single linear series), or multiserial (two or more chambers per whorl). If the test coils in a single plane around the growth axis, it is called planispiral. In trochospiral tests the chambers are arranged in a helical coil. A common arrangement in which five chambers are visible is called quinqueloculine. Internally the chambers are separated by walls called septa. An opening (aperture) in the wall of the final chamber connects external pseudopodia and internal endoplasm. The external surface of the tests is smooth or variously ornamented (e.g. spines and keels).
Foraminifera are small, predominantly marine heterotrophic protists that construct chambered shells (tests). The group belongs to the phylum Sarcodina and
Diagnostic criteria, classification and major groups Diagnostic criteria of foraminifera are (1) wall composition and microstructure, (2) chamber arrangement
10.2.2 Foraminifera and other Protozoa The second prominent category of fossils found in thin sections of carbonate rocks includes small-sized singlecelled protozoan organisms : foraminifera, radiolarians and the calpionellids (an extinct group of the ciliates).
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Wall Structures of Foraminifera
Major foraminiferal wall structures seen in thin sections and studied in SEM are agglutinated, microgranular calcareous, porcelaneous calcareous and hyaline calcareous (Fig. 10.22). Wall composition and microstructures are the most important features for classifying foraminifers on a suprageneric level, and are the subjects of current critical evaluations (Piller 1990). Note that the groups mentioned below are those of more conventional classifications. Loeblich and Tappan (1988) recognized 12 suborders acknowledging specific skeleton mineralogy (e.g. Involutinina originally with aragonitic walls, Fig. 10.27, see Pl. 111/15) or specific lamellation pattern (e.g. Globigerinina, Pl. 73/9).
Fig. 10.21. Major wall structures of foraminiferal groups. Agglutinated walls are characterized by individual grains from the sea floor bound together in a cement, which may be calcareous, ferruginous or organic. Fusulinid foraminifera have a microcrystalline calcitic wall, that is often differentiated into several units. The porcelaneous walls of miliolinid foraminifera are submicroscopical crystals of magnesian calcite, usually in random orientation. Rotaliinid foraminifera have glassy-looking (hyaline) walls, which may be differentiated into layers. AG - agglutinated grains, ALV - alveoles, BL basal layer, C - cement, IL - inner layers, OL - outer layer, ORL - organic lining. Arrows indicate the optical orientation of calcite crystals. Modified from Brasier (1980) and Bignot (1985).
and septal addition, (3) apertural characters and modifications, and (4) chamber form and ornamentation. These criteria are by far the most important characteristics used in the classification and taxonomic differentiation. Further properties particularly important for discriminating modern foraminifera are life habits and habitats, protoplasma features, ontogenetic changes, and genetics. The last criterion has become increasingly more meaningful for higher-level classification of ancient foraminifera as well.
The wall of the agglutinated textulariid foraminifera consists of particles held together by organic material and cement of different composition. The particles can be coarse (see Ammobaculites, Pl. 69/7; Haddonia, Pl. 75/9; Reophax; Tolypammina, Pl. 51/2, Pl. 111/18) or fine (e.g. Alpinophragmium, Pl. 111/1; Chrysalidina, Pl. 72/3). In more complex Mesozoic taxa an outer microgranular wall covers a complex inner layer that may be folded inward into the chambers, producing labyrinthic or alveolar structures (see Alveosepta, Pl. 69/8; Orbitolina, Pl. 70, Pseudocyclammina, Pl. 69/1). Textulariid foraminifera occur in various marine environments. They are often dominant in transitional environments with brackish and marsh conditions as well as at bathyal and abyssal depths. Fusulinid foraminifera with microgranular calcareous wall structures are represented by smaller and larger foraminifera. The wall of the smaller foraminifer consists of microgranular calcite cemented by microcrystalline calcareous cement (see Earlandia, Pl. 67/ 2; Endothyranopsis, Pl. 67/1; Cribrogenerina, Pl. 67/ 10; Tuberitina, Pl. 67/2). Larger fusulinid foraminifera, (Fusulinacea; Viséan to Late Permian; Pl. 68) are morphological highly differentiated. Miliolinid foraminifera exhibit a porcelaneous calcareous wall structure without pores, characterized by a shiny, smooth appearance of the test. Examples are Austrocolomia, Pl. 111/23; Alveolina, Pl. 75/2; Galeanella, Pl. 111/4-5; Hemigordius, Pl. 67/7). The rotaliinid foraminifera are characterized by hyaline calcareous wall structures. In cross sections organic layers or calcareous laminations may be visible. Lamellar tests can be recognized by thicker early chambers, thinner later chambers and a single-thickness final chamber. Lamellar walls are characteristic for taxa
Environment of Foraminifera Wall and microstructure
455 Appearance
Criteria
Group
Time range
Organic, proteinaceous material
Poor preservation potential
Tubular, single-chambered round or flask-shaped forms
Allogromiina
Cambrian to Holocene
Various foreign grains glued together by organic, calcareous, ferrugineous or siliceous cement
Dark in transmitted light
Simple agglutinated (arenaceous) walls. Advanced Mesozoic forms with outer layer covering complex inner layer forming chambers, alveolar and labyrinthic structures
Textulariina
Early Cambrian to Holocene
Microgranular calcite, randomly or normally oriented crystals of unequal size
Dark in transmitted, opaque in reflected light
Advanced forms show differentiation of the wall into two or more layers
Fusulinina (include Carboniferous and Permian larger foraminifera, Fusulinacea)
Early Silurian to Late Permian
Calcite crystals oriented randomly or in a brick-like pattern, Low-Mg and High-Mg calcite
Dark or amber-colored in transmitted light, milky white and brilliant in reflected light. Under crossed nicols calcite crystals appear as tiny multicolored specks
Porcelaneous. Imperforate wall Miliolina with inner and outer surface layers
Crystals arranged in prisms Light-colored, clear or gray in with c-axes normal to test transmitted light, glossy and surface radial hyaline or dark in reflected light randomly granular hyaline. Low- or High-Mg calcite. Few groups with aragonite
Hyaline, finely perforate wall. Non-lamellar, mono- or multilamellar
Carboniferous to Holocene
Rotaliina, including Triassic to benthic Rotaliacea (many Holocene multilamellar and Cenozoic larger foraminifera) and planktonic Globeriginacea (predominantly Cretaceous and Cenozoic)
Fig. 10.22. Microstructure, wall composition and major systematic categories of foraminifera. Note that the suprageneric high-level classification of the foraminifera is much more complicated than shown in this simple graph.
of the benthic Rotaliacea with multichambered spiral shells (Orbitoides, Pl. 72/7-10; Nummulites, Pl. 75/2) and the planktonic Globigerinacea with bubble-shaped chambers arranged in expanding spirals (see Globigerinoides, Pl. 73/9). Environmental constraints of foraminifera Foraminiferal environments include shelves, platforms, ramps, reefs, slopes and basins. Most modern foraminifera are marine and benthic and live from the intertidal zone down to the deep sea. Some benthic foraminifera inhabit fresh and brackish waters. Planktonic foraminifera live in the upper part of the open oceans. (A) Benthic foraminifera are restricted to a narrow range of depths and substrates, and have characteristic tolerances to salinity and temperature: Salinity: Most foraminifera are found in water with a normal marine salinity of 35‰. Diversity decreases with decreasing salinity. Low salinities (e.g. in coastal lagoons) favor low-diversity assemblages, rich in agglutinated forms. High carbonate concentrations in hypersaline marine waters favor the occurrence of porcelaneous Miliolina. Triangular plots of the relative proportions of Textulariina, Miliolina and Rotaliina have proved to be valuable indicators of paleosalinity and paleotemperatures and can also be derived from thinsection analysis of foraminiferal limestones. Temperature: Benthic foraminifera enable warmwater, temperate and cool-water environments to be
distinguished. Cold-water foraminiferal associations inhabiting the shallow waters of high latitudes or the cooler waters of the deep ocean are fairly uniform. Because of the lower solubility of calcium carbonate in warmer than in cool waters, accumulations of foraminifera and foraminiferal sediments are more frequent in carbonate platforms of low latitudes. The increasing solubility of carbonate with depth results in a distinct and final drop of calcareous foraminifera at the lysocline and the CCD (see Box 2.6) and results in a dominance of agglutinated foraminifera in deep-water environments. Oxygen and nutrients: Most benthic foraminifera occur in areas of normal oxygen concentration, although deep-water forms may tolerate very low oxygen levels. The composition of slope and outer shelf foraminiferal faunas is considerably influenced by upwelling oxygen-poor and nutrient-rich waters. Specific faunas occur in the oxygen-minimum zone of the continental slope. Substrate: Various types of sea bottom contain distinctive associations differing in shell morphology and taxonomic composition. Microenvironments around the sediment-water interface provide distinct substrate-controlled microhabitats for benthic foraminifera living free or attached to sedimentary grains, algae and sea grass, hard substrates or as infauna at depths between 0.5 to about 10 cm (Langer 1988; Yamano et al. 2000). Water depth: Many foraminifera are limited to the photic zone because of their endosymbiotic associa-
456
tions, or because they prefer habitats within algal and sea grass meadows. Depth distribution of endosymbiotic foraminifera depends on the type of endosymbionts. Foraminifera with green or red symbionts are restricted to shallower depths than those with diatoms that can occupy greater depths. Near-coastal, lagoonal, and inner and outer shelf environments as well as upper and lower slope and basinal environments differ in foraminiferal association patterns. Shallow nearshore and lagoonal environments, down to about 50 m, are characterized by porcelaneous miliolid foraminifera. Planktonic foraminifera are absent or rare in shallow-marine environments. Hyaline rotaliinids occur in shallow and deeper waters. Agglutinated tests are ubiquitous but common on the slope and also occur at depths of several thousand meters. In outer shelf environments (about 50 to 200 m) the faunas are dominated by perforate rotaliinid species. Bathyal depths tend to be dominated by biserial and
Planktonic Foraminifera
triserial agglutinated foraminifers along with some hyaline rotaliinids and porcelaneous miliolinids. (B) Planktonic foraminifera are adapted to different oceanic layers with particular temperatures and densities. Temperature-controlled planktonic species exhibit specific morphological characteristics (shape, size and coiling direction). These adaptions are used together with geochemical and isotope data as proxies for paleoceanographic interpretations of Cretaceous and Cenozoic pelagic carbonates. Most planktonic foraminifera are restricted to marine salinities of 30–40‰ and live in defined parts of the water column. The majority occurs in the upper part of the water column within the upper 50 to 100 m near the surface (photic zone). Below 200 m they are virtually absent. The thin-walled spherical, symbiont-bearing globigerinids prefer the surface waters where there is light, while those with thicker walls and keeled shells (lacking photosymbiosis) tend to be
Fig. 10.23. Different sections through foraminiferal tests of different shapes. A. Single chambered: 1 - axial section, 2 -longitudinal, 3 - tangential, 4 -transversal, 5 - oblique. B. Multichambered uniserial: 1 - axial, 2 - longitudinal, 3 - transversal, 4 - oblique passing through two chambers. C. Multichambered biserial: 1 - axial sagittal, 2 - axial frontal, 3 - longitudinal, 4 - axial oblique, 5 - transversal, 6 - transversal oblique. D. Multichambered planspiral (a, lateral view; b, peripheral view): 1 - axial, 2 - transversal equatorial. E. Multichambered triserial (a, lateral view; b, oral view): 1 - axial, 2 - longitudinal passing through two chambers, 3 - longitudinal passing through one chamber, 4 - transversal. F. Multichambered trochospiral: 1 - axial, 2 - transversal dorsal, 3 - transversal, passing through an embryonal chamber, 4 - transversal ventral. After Neumann (1967).
References on Foraminifera
found at slightly greater depths. Planktonic foraminifera avoid terrigenous input. Their shells, therefore, accumulate in the bathyal depths of open-marine basins. Foraminifera in thin sections Relevant criteria for the taxonomic discrimination of foraminifera in thin sections are wall structure, arrangement of chambers and the form and position of the apertura (Arnaud-Vanneau 1980). Foraminifera should be studied in plane and polarized light.
457
Taxonomic determination of foraminifera in thin sections of limestones may be difficult because only randomly oriented sections are commonly available, and wall structures are often obliterated by diagenesis. Some of these problems can be overcome by the following steps: • Compare foraminifera dissolved from limestones by acids with those found in thin sections of the limestone (see Pl. 9/1, 2); • compare hypothetical figures of geometrically dif-
Box 10.4. Selected references on foraminifera in thin sections. The list includes papers accompanied by plates that are useful for determining foraminifera and that assist in interpreting depositional environments of limestones with foraminifers. Planktonic Cretaceous and Tertiary foraminifera: Atlas 1979; Caron 1985; Lipps 1966; Noé 1993; Robaszynski et al. 1984; Sliter 1989; Van Konijnenburg et al. 1998; Radoicic 1995; Weidich 1984, 1990 Tertiary: Bonnefous and Bismuth 1982; Charproniere 1975; Drobne 1974; Drobne et al. 1977; James et al. 1962; Geel 2000; Ghose 1977; Hallock and Glenn 1986; Hallock et al. 1988; Hottinger 1997; Hottinger and Drobne 1980, 1988; Kenawy 1966; Less 1998; Luterbacher 1984; Moussavian and Höfling1993; Perrin 1992, 1994; Reiss and Hottinger 1984; Ross 1979; Schaub 1981; Tourmakine and Luterbacher 1985 Late Cretaceous: Beckmann et al. 1982; Berthou 1973; Berthou and Schroeder 1978; Bilotte 1984, 1985; Carbone et al. 1971; Caus 1988; Chiocchini et al. 1983; De Castro 1971, 1980, 1988; Eames et al. 1962; Fleury et al. 1985; Gargouri 1988; Gusic et al. 1988; Hamaoui and Brun 1974; Hamaoui and Saint-Marc 1970; Hart 1988; Höfling 1988; Hottinger 1966; Krijnen et al. 1993; Kuss 1994; Kuss and Conrad 1991; Kuss and Malchus 1989; Papp 1955; Saint-Marc 1982; Septfontaine 1981; Velic and Vlahovic 1994 Early Cretaceous: Arnaud 1981; Arnaud-Vanneau 1980, 1986; Arnaud-Vaunard and Arnaud1978; Arnaud-Vanneau and Masse 1989; Arnaud-Vanneau et al. 1988; Bassoullet et al. 1985; Becker 1999; Berthou and Schroeder 1978; Bodrogi et al. 1994; Bucur 1988; Cherchi et al. 1981; Chiocchini and Mancinelli 1978; Chiocchini et al. 1983, 1988; Conrad 1977; Darga and Schlaginweit 1991; De Castro 1988; De Castro and Peybernes 1983; Douglass 1960; Dragastan et al. 1975; Gallagher 1997; Gusic 1975, 1981; Hottinger 1967; Husinec et al. 2000; Kaever 1967; Masse 1976; Moullade et al. 1985; Neumann 1958; Peybernes 1979; Ramalho 1971; Sanchez et al. 1991; Schroeder et al. 1968; Schroeder and Neumann 1985; Septfontaine 1981; Velic 1988 Late Jurassic: Allemann and Schroeder 1972; Altiner and Septfontaine 1979; Bassoullet et al. 1985; Chiocchini and Mancinelli 1978; Chiocchini et al. 1988; Darga and Schlaginweit 1991; Dragastan et al. 1975; Farinacci 1964; Grossowicz et al. 2000; Gusic 1969, 1977; Hottinger 1967, 1973; Misik and Sotak 1998; Pelissie and Peybernes 1984; Peybernes 1979; Ramalho 1971; Schmalzriedt 1991; Septfontaine 1988; Stam 1986 Middle Jurassic: Bernier and Neumann 1970; Blau and Wernli 1999; Cherchi and Schroeder 1981; Ruggieri and Giunta 1965; Furrer and Septfontaine 1977 Early Jurassic: Blau 1987a, 1987b; Gazdzicki 1983; Gusic 1975; Hottinger 1967; Pirini 1965; Septfontaine 1984, 1985 Triassic: Gazdzicki 1983; Hohenegger and Lobitzer 1971; Hohenegger and Piller 1975a, 1975b, 1975c, 1977; KoehnZaninetti 1969; Oravecz-Scheffer 1987; Piller 1978; Rettori 1995; Salaj et al. 1983; Schäfer and Senowbari-Daryan 1978; Schott 1983; Zaninetti 1976 Permian: Groves and Wahlman 1997; Henbest 1963; Kahler 1983; Leven, 2003; Okimura et al. 1985; Pasini 1965; Ross 1982, 1995; Wilde 1990 Late Carboniferous: Flügel 1971; Forke et al. 1998; Groves and Wahlman 1997; Henbest 1963; Pasini 1965; Ross 1982; Toomey 1974; Vachard and Krainer 2001a, 2001b Early Carboniferous: Conil 1964; Conil et al. 1979, 1990; Cossey and Mundy 1990; Fewtrell et al. 1981; Gallagher 1997, 1998; Horbury and Adams 1996; Krainer and Vachard 2002; Mamet 1977; Mamet et al. 1986; Mamet and Pinard 1990; Vachard 1988 Devonian: Chuvashov 1965; Flügel and Hötzl 1971; May 1994; Petrova 1981; Poyarkov 1979; Toomey 1965; Vachard 1988; Vachard and Massa 1989; Zadorozhnyi and Yuferev 1981
458
ferent tests with the random figures displayed in thin sections (Fig. 10.23; Neumann 1967; Hottinger 1971); • compare free tests and line drawings of possible sections using a defined coordinate system (Cummings 1958); • reconstruct three-dimensional tests enclosed in limestones using serial peels and computer-assisted designs; • use X-ray microradiographic techniques in taxonomic differentiation of pyritized foraminifera occurring within micritic limestones (Mehl and Noé 1990).
Larger foraminifera These foraminifera are significant constituents of Late Paleozoic, Jurassic, Cretaceous and Early Tertiary limestones. These foraminifera include systematic groups that are not closely related taxonomically. They are characterized by tests with sizes between about 0.5 mm and several centimeters (up to 20 cm), complicated internal structures, and the often distinct differentiation into sexual and asexual reproductive stages that can be distinguished by the size of the initial chambers and the shells. Larger foraminifera are thin-section fossils per se because of the necessity of studying their endoskeletons in oriented sections. This taxonomic differentiation requires sections exhibiting the size and shape of the test, the morphology of the initial and embryonal chambers, the character of the whorls, the shape and number of septa as well as dimensional criteria. Modern larger foraminifera: About ten recent families of the tropics and subtropics and thirteen additional fossil families include taxa with large calcareous tests, belonging to widely divergent taxa of the suborders Textulariina, Fusulinina, Miliolina and Rotaliina. These foraminifera attain a volume of at least 3 mm3 before reaching reproductive maturity and many reach volumes of more than 300 mm3 (Ross 1979). Larger foraminifera are typically associated with tropical and subtropical shallow-water carbonate sediments (Hottinger 1983). They can form a considerable portion of the skeletal debris of reef and platform environments. Their carbonate production rates are comparable to those of calcareous algae and corals (Hallock 1981). Most larger foraminifera live in symbiotic association with unicellular algae. In tropical carbonate shoal and reef environments, calcareous larger foraminifera house various algae, including dinophyceans, diatoms, chlorophyceans and rhodophyceans. The sym-
Paleozoic Foraminifera
bioses provide nutrients from photosynthesis, facilitate life in different water depths and favor maximum carbonate precipitation by carbon dioxide uptake. Differences of light attenuation by the water column and of water energy are reflected by different wall structures (Hohenegger et al. 1999). Most larger foraminifera are restricted to relatively shallow well-lighted sea bottom, and if untransported, their presence is generally indicative of depths less than 30 m (Hallock 1986). Symbiont-bearing larger miliolinid and rotaliinid foraminifera are important constituents of modern reef flat and forereef biota (Hottinger et al. 1993; Hohenegger 1994). In addition to symbioses, the distribution of modern larger foraminifera is controlled by surface water temperature (within the 15° isotherm) and nutrients (oligotrophic, nutrient-poor waters; Langer and Hottinger 2000). Ancient larger foraminifera: The first larger foraminifera appeared late in the Early Carboniferous. The fusulinids (Pl. 68) are useful for biostratigraphically subdividing Late Carboniferous and Permian carbonate rocks and for biogeographical interpretations. Agglutinated larger foraminifera are important index fossils used for regional and global correlations in Jurassic (Pl. 69) and Cretaceous shelf carbonates, particularly in the Early Cretaceous (orbitolinids, Pl. 70, Pl. 71). Late Cretaceous larger foraminifera are represented by the rotaliinid orbitoideaceans (Pl. 72/7-10), Tertiary larger foraminifera by the miliolinid alveolinid and rotaliid nummulitid foraminifera (Pl. 74, Pl. 75). Late Paleozoic Fusulinacea as well as other fossil Mesozoic and Cenozoic larger foraminifera are believed
Fig. 10.24. Early Paleozoic foraminifera are characterized by agglutinated or calcareous microgranular, usually singlechambered tests. Parathuramminid foraminifera are common in Middle and early Late Devonian lagoonal limestones. The sample shows Parathurammina dagamarae (Suleimanov). Middle Devonian (Givetian): Letmathe, Sauerland, Germany. Scale is 500 µm.
Paleozoic Foraminifera
to have had symbionts (Ross 1972; Hallock 1985). Many larger foraminifera have chambers which are subdivided into smaller chamberlets. Most fusulinids subdivided the chambers by septal fluting (folding of septal walls) similar to Tertiary alveolinids. Other larger foraminifera have small bar-like elements starting at the septa and forming additional subdivisions within the shell. These exoskeletal structures are known from the Late Permian verbeekinid fusulinids, but are also very common in Jurassic, Cretaceous and Tertiary agglutinated foraminifera (e.g. Pl. 69/2). Endoskeletal structures are arranged from one septum to the other and reflect the direction of the protoplasma flow. Some Late Cretaceous and Cenozoic larger foraminifera (Rotaliina) have a channel system and a lamellar-perforated wall consisting of inner and other lamellae. Distribution and fossil record of foraminifera Examples of foraminifera common in thin sections of Late Paleozoic, Jurassic, Cretaceous and Tertiary limestones are shown on Pl. 67 to Pl. 75. For Triassic foraminifera see Pl. 111. Box 10.4 lists relevant papers. Benthic foraminifera have been known since the Early Cambrian. During the Carboniferous and Permian an explosive radiation took place, specially among larger foraminifera (Fusulinacea). Carbonate rocks with abundant benthic foraminifera have been common since the Early Carboniferous and became widely distributed in the Cretaceous and Tertiary. Planktonic foraminifera evolved in the Middle Jurassic and became major constituents of the marine plankton during the Cretaceous. Early Paleozoic: An overview of Paleozoic foraminifera was given by Ross and Ross (1991). Cambrian foraminifers are poorly documented (e.g. Cherchi and Schroeder 1984). Evolution was slow during the Ordovician. Silurian populations were rather small and diversity very low. Nearly all Cambrian to Silurian foraminifera are characterized by agglutinated single-chambered tests. The first major radiation took place in the Middle Devonian and culminated in the Late Frasnian. Devonian microgranular and agglutinated foraminifera occur in shallow-marine and basinal facies types. Fusulinid Parathuramminacea were common in back-reef and lagoonal carbonates (Fig. 10.24). The Devonian fauna became almost extinct in the Famennian and was replaced by the endothyrid group. This group had a major outburst in the Tournaisian and the Viséan. Late Paleozoic: Nineteen Tournaisian, Viséan and lower Middle Carboniferous biozones have been established based on assemblages of benthic smaller fora-
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Fig. 10.25. Smaller foraminifera are used for the biozonation and correlation of carbonates deposited on Carboniferous shelves and ramps. The biozones can be correlated both on a regional and a global scale. The sketch exhibits a grainstone with different sections of Endothyranopsis and other foraminifera, indicating an Early Viséan age. After Sorby (1879).
minifera taken from different phylogenetic lines (Conil and Lys 1964; Mamet 1977; Conil et al. 1981; Lys 1988; Fig. 10.25). These zones were recognized in limestones formed in shallow subtidal parts of platform and ramp environments. Smaller foraminiferal assemblages also assist in the zonation of Late Permian shelf carbonates (Okimura et al. 1985; Noé 1988). Smaller foraminifera often comprise only less than 5% of the rock volume, but may increase to more than 20% in grainstones. Common shallow-marine facies types of bedded Viséan platform carbonates yielding benthic smaller foraminifera are (a) grain-, pack- and wackestones with crinoids, bryozoans, brachiopods, dasyclad green algae and ungdarellacean red algae (Pl. 56/7), (b) fine-grained peloidal grainstones, and (c) lagoonal wackestones and packstones. Foraminifera are rare to absent in oolitic limestones (but see Fig. 4.22), deeper-water crinoid limestones and basinal sediments. In graded calciturbidites intercalated within basinal sediments, smaller foraminifera are differentiated by size, indicating hydraulic sorting of microbiota derived from all the parts of adjacent platforms (Herbig and Mamet 1994). Statistically-based association patterns of Early Carboniferous and Permian foraminifera are indicative of specific shelf habitats (Gallagher 1998). Distribution of taxa, frequency patterns and association types of Paleozoic smaller foraminifera permit various parts of platforms, ramps and mud mounds to be differentiated Text continued on p. 464
460
Late Paleozoic Smaller Foraminifera
Plate 67 Carboniferous and Permian Smaller Foraminifera Late Paleozoic foraminifera shown on this plate comprise both ‘smaller foraminifera’ with approximately 0.1 and 0.6 mm tests, and millimeter- to centimeter-sized ‘larger foraminifera’ (see Pl. 68). Foraminiferal assemblages enable biostratigraphic zonations of Late Paleozoic, particularly Early Carboniferous ramps and platforms (–> 1; Pl. 108/2) and Permian shelf deposits to be biostratigraphically zoned. Distribution of taxa, frequency patterns and association types of Paleozoic smaller foraminifera are used in differentiating the facies depositional settings of ramps and mud mounds as well as recognizing environmental conditions (water energy, paleobathymetry and sedimentation rates). Association patterns are indicative of local shelf habitats. This plate displays taxa common in Carboniferous (–> 1, 6, 11) and Permian (–> 2-5, 7-10) shelf and reef carbonates. Agglutinated foraminifera (–> 9) as well as foraminifera with calcitic microgranular (–> 1, 5, 6) or porcelaneous walls (–> 7) are used in paleoecological analyses. Note the different growth forms, including unilocular tests, shells consisting of a proloculus followed by a tubular non-divided chamber (–> 2), multichambered uniserial (–> 4) or coiled tests (–> 11) and conical trochospiral tests (–> 1). 1
Biozonation based on smaller foraminifera. The succession of foraminiferal assemblages permits zonation and global correlation of Tournaisian and Viséan ramps and platforms. The sample shows a foraminiferal grainstone with axial sections of Forschia prisca Mikhailov (center left) and equatorial sections of Endothyranopsis crassa (Brady), top right, indicating the Late Viséan foraminiferal zone 15. Early Carboniferous: Basin of Dinant, Belgium. 2 Encrusting foraminifera and very small tubular microfossils, usually also assigned to foraminifera, are common constituents of Late Paleozoic shelf carbonates. Some of these foraminifera encrusted algal blades, others hard bottoms. Note that recrystallization only affected the matrix, but not the microgranular foraminiferal tests. Abundant taxa are Tuberitina Galloway and Harlton (T) belonging to the parathuramminacean Fusulinina, and Earlandia Plummer (E) assigned to the earlandiacean Fusulinina. Both genera occurred in inner and outer shelf environments. The sample is from the top of a meter-scaled cyclothem. Early Permian (Asselian): Carnic Alps, Austria. 3 Miliolinid attached foraminifera. The initial chamber of Apterinella Cushman and Waters is encircled by uncoiled chambers planspirally enrolled and later arranged in zigzag bends. The wall appears dark in transmitted light. The sample is from back-reef carbonates deposited in the leeward position of a platform margin reef. Middle Permian (Capitanian): Permian Reef Complex, Texas, U.S.A. 4 Differentiated wall structure. The wall of Pachyphloia/Langella exhibits a thin dark inner layer (arrow) and a thick hyaline outer layer. The uniserial elongate test consists of strongly overlapping chambers. Middle Permian: Pietra di Salomone near Sosio, western Sicily. 5 Differentiated test. Axial sections of Lasiotrochus tatoensis Reichel (left) and Lasiotrochus minor Reichel. Note initial chamber and microgranular wall. Arrows point to hyaline pillars composed of transparent calcite and tubular chamberlets. Genus range: Permian. Same locality as –> 4. 6 Differentiated test. Oblique axial section of Howchenia Cushman. Conical test. The undivided tubular chamber is coiled in a high spire around an umbilical region filled with fibrous calcite. The genus is restricted to Viséan to Moscovian and common in intraclastic wackestones deposited in deep shelf environments affected by storm waves. Early Carboniferous (Badstuben Breccia, Viséan): Nötsch, Carinthia, Austria. 7 Miliolinid foraminiferal grainstone with Hemigordius Schubert. The porcelaneous discoidal tests appear dark in transmitted light. The dark tint is intensified by early pyritization of the tests. The sample is from black limestones representing bioclastic shoals within a restricted lagoon. Late Permian (Bellerophon Formation, Dorashamian): Comelico, northern Italy. 8 Tetrataxis Ehrenberg, a common Carboniferous and Permian genus occurring in open-marine subtidal crinoid and bryozoan thickets at or below the fair-weather wave base, which were reworked by currents and storms. The wall consists of an outer microgranular layer and hyaline radiate inner layer. Tetrataxis was a loosely attached mobile foraminifer that responded to environmental extremes by attaching onto substrates (Cossey and Mundy 1990). The test is encrusted by Tuberitina, (T). Genus range: Early Carboniferous to Triassic. Early Permian: Carnia, Italy. 9 Agglutinated textulariid foraminifera. Tangential and more central longitudinal sections. The wall is composed of fine sand grains. The chambers are uniserially arranged. Early Permian: Southern Karawanken Mountains, Austria. 10 Cribrogenerina Schubert, a common Permian palaeotextulariid foraminifera. The genus is characterized by initially biserial, later uniserial, broad chambers with multiple and cribrate roofs. The oblique section displays the uniserial part of the test. Late Permian (Zazar Formation): Idrija, Slovenia. 11 Bradyina Möller, an endothyracean fusulinid foraminifer which was common in Carboniferous shallow-marine algal meadows. Planispiral involute test, sutures deeply incised. Calcareous wall, with microgranulate outer layer (tectum, T) and inner perforate structure (P). Late Carboniferous (Gzhelian): Nassfeld, Carnic Alps, Austria. –> 4, 5: Flügel et al. 1991; –> 7, 10: Noé 1987
Plate 67: Carboniferous to Permian Smaller Foraminifera
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Larger Foraminifera: Fusulinids
Plate 68 Late Paleozoic Larger Foraminifera: Fusulinids This plate exhibits characteristics in the morphology and wall structures of Fusulinids of different ages. Fusulinids were benthic rock-forming larger foraminifera with a restricted geological range (Viséan to latest Permian), a wide geographic spread and distinct evolutionary lines allowing global biostratigraphic zones to be defined. The mm- to cm-sized tests have spindle-like, subcylindrical or globose shapes (Fig. 10.26). The complex internal structure is caused by variations in coiling, folding patterns of the partitions between chambers (septa) and additional elements projecting from the walls into the chamber interior (–> 9). The walls are microgranular and appear dark in transmitted light and opaque in reflected light. They consist of one to four layers (–> 2, 4, 6). Pores occur within walls and septa (–> 4). The determination of fusulinid species requires sections in the axis of coiling and passing through the initial chamber (–> 1, 3, 5) and cooperation with micropaleontologists. However, even nonspecialists can use these foraminifera for a very rough dating of shelf carbonates, if they look at the wall structure type. The wall structures increase in complexity over time, starting with relatively simple multilayered walls (–> 2) in the Early and Middle Carboniferous, and continuing with double-layered walls (–> 4) characteristic of Late Carboniferous and Early Permian Schwagerinids (–> 3). Middle and Late Permian Neoschwagerinids have two-layered walls and shorter and longer projections into the chambers forming small chamberlets (–> 6). Most fusulinid families became extinct near the end of the Middle Permian (Guadalupian), the last Fusulinids at the PermianTriassic boundary. Late Permian taxa are usually small and often characterized by distinct deviations in the coiling planes during individual growth (–> 7, 8). Fig. 10.26. Mixture of spherical and spindle-shaped fusulinid shells derived from different subenvironments. Poorly sorted bioclastic grainstone of an outer carbonate platform. The large spherical test (Zellia heritschi mira Kahler and Kahler) is cut through the initial chamber. Spindle-shaped forms are represented by two sections cutting the tests tangentially (bottom) and transversially (top). Spindleshaped tests are peripherally micritized and often broken. Early Permian (Asselian): Carnic Alps, Austria. Scale is 1 mm. 1
Beedeina Galloway. Axial section of the fusiform test. Note the initial chamber (proloculus; center) and the strongly fluted septa. Generic range: Middle to Late Moscovian. Middle Carboniferous (Moscovian): Kalinovo, Donets Basin, Ukraine. 2 Beedina sp. The wall structure exhibits four parts: Dark tectum (t), light diaphanotheca (d) bordered by the upper tectorium (ot) and lower (ut) tectorium. Same locality –> 1. 3 Sphaeroschwagerina carniolica Kahler and Kahler. Axial section of the spherical test. Note difference between the dimensions of inner and outer whorls. The thin septa are weakly fluted. Wall thin, thickening as the test enlarges. Species range: Early Permian. Asselian (Grenzland Formation): Rudnigsattel, Carnic Alps, Austria. 4 Sphaeroschwagerina carniolica Kahler and Kahler. Wallstructure: Tectum (t) and ‘honeycomb-like’ alveolar keriotheca (k). The septum (s) is pierced by septal pores (sp). Same locality –> 3. 5 Yabeina syrtalis (Douvillé). Axial section of the inflated fusiform test. Note the abundance of whorls, numerous relatively thin septa and short septula (–> 6). Middle Permian sponge-algal reef (Wordian): Bati Beni section, Djebel Tebaga, southern Tunisia. 6 Yabeina syrtalis (Douvillé). Wall structure consisting of two parts: tectum and keriotheca. The chambers are subdivided by parachomata (pch) and transverse septula (ts) and filled with granular carbonate cement. Same locality –> 5. 7 Reichelina Erk characterized by a marked difference between an early planispirally enrolled stage and uncoiled final whorl with septa. Late Permian (Wargal Formation, Wuchiapingian): Chidru Nala, Salt Range, Pakistan. 8 Codonofusiella Dunbar and Skinner. Note sharp change in coiling axis. Same locality –> 7. 9 Random cuts: Microfacies thin sections commonly yield only randomly cut tests which are difficult to treat taxonomically. But sometimes specific structures are sufficient for a rough determination. Example: The broad and regularly spaced transverse septula indicate a neoschwagerinid form and points to a Middle Permian age of the sample. Middle Permian: Bohinska Bela near Bled, Slovenia. 10 Random cuts: Sagittal section (perpendicular to the axial section and passing though the initial chamber) and tangential section of fusulinids. Note deformation of the last whorl of the left specimen (arrows) pointing to a pathological disturbance of shell growth rather than a mechanical destruction (Kahler 1988). Same locality as –> 9. –> 1–8: Courtesy of H. Forke (Erlangen)
Plate 68: Larger Foraminifera: Fusulinids
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Triassic Benthic Foraminifera
as well as environmental conditions (water energy, paleobathymetry and sedimentation rates) to be recognized. The use of Late Paleozoic smaller foraminifera in microfacies studies requires • closely spaced sampling (within a distance of < 10 m, often within a distance of less than 1 m), • samples displaying the diversity of the foraminifera (e.g. Pl. 67/1), • systematic determination based on well-defined section plains (Conil et al. 1979), • quantitative analysis using the number of different taxa (e.g. per square centimeter of thin-sections), presence/absence patterns, and • statistical treatment of the quantitative data (Gallagher 1997). The millimeter- to centimeter-sized fusulinid foraminifera (Pl. 68) are excellent biostratigraphic markers on stable, shallow carbonate shelves. The global and regional zonation of the Middle Carboniferous to Late Permian shelf carbonates relies on a succession of genera and species (Ross and Ross 1987). Fusulinids are multichambered. Growth starts with an initial chamber (proloculus). Specimens of sexual generation have a small microspheric proloculus, those of asexual generations a large megalospheric proloculus. Chambers are added to form a coiled shell by extension of the spiral wall (spirotheca). After several chambers are added, a central tunnel is resorbed through the septa allowing free communication between the chambers. Main evolutionary trends are an increase in the size of the initial chamber, an increase in overall size, and an increase in the complexity of the test wall. Note that present classifications differentiate a suborder Fusulinina that includes the large-sized typical Fusulinacea as well as various groups of smaller foraminifera. Fusulinids were adapted to shallow warm and warmtemperate waters, usually in tropical and subtropical regions on shelves, platforms and reefs. They lived in normal marine, well-oxygenated environments on subtidal open shelves, in depths between a few meters to a few tens of meters. Variations in test shape indicate specialization to different shallow water subenvironments (e.g. sand bars). Spindle-shaped fusulinids with pole torsions were concentrated in high-energy stress environments, spherical forms in low-energy lagoonal environments (Ross 1982). The life habits and taphonomic history of fusulinids can be derived from shell construction, encrustation patterns, shell damage, and
Fig. 10.27. Involutinid foraminifera are abundant elements of reefal Late Triassic carbonate platforms (see Pl. 111). A: Triasina hantkeni Majzon defines a Rhaetian biozone. The species occurs in open-marine platform environments. Allgäu, southern Bavaria, Germany. B: Angulodiscus permodiscoides Oberhauser and related taxa are common in muddy near-reef sediments. Left subaxial section, right oblique-horizontal section. Note the distinct fibrohyaline wall structure. Late Triassic (Rhaetian): Kendelbachgraben, Austria. Scale is 500 µm. Further Triassic foraminifera are depicted on Pl. 111.
sedimentary infillings (Kahler 1988). Worn and rounded fusulinid tests as well as mixtures of juvenile and adult specimens are indicative of transport, erosion and redeposition processes that may sometimes produce mixtures of fusulinids derived from different subenvironments (Fig. 10.26), or even different stratigraphic ages. Carboniferous and Early Permian fusulinids are particularly common in bedded medium to light gray grainstones, pack- and wackestones deposited on shallow carbonate shelves, banks and platforms, and in calcareous shales. They mark specific levels in cyclic successions. Middle and Late Permian verbeekinid and neoschwagerinid fusulinids are often associated with platform margin reefs (Pl. 68/5, 9). Some species are restricted to back-reef lagoons. Fusulinids also occur as clasts in debris flows and turbidites deposited in deeper-water slope and toe-of-slope settings.
Cretaceous Foraminifera
Fusulinid coquinas (grain- and packstones) may represent parautochthonous assemblages, allochthonous downslope accumulations between buildups, concentrations through storms, or transgressive deposits during times of reduced mud deposition, or they may just be diagenetic packstones caused by intergranular solution. Syn- and postsedimentary changes of test shapes and internal structures (e.g. broken or distorted shells, shells with breaks of septa) are indications of late diagenesis and burial history (see Pl. 36/1, 3). Triassic: Benthic foraminifera are common in thin sections of Middle and Late Triassic platform and reef carbonates (see Pl. 111). Systematic determinations must consider diagenetic changes of wall structures that are particularly frequent in involutinid foraminifera (Hohenegger and Piller 1975). Some taxa are good index fossils and are used for biozonations (Gazdzicki 1983, Salaj et al. 1983; Fig. 10.27). Open and restricted platforms as well as different parts of reefs can be discriminated using distinct foraminiferal associations defined by the occurrence or quantitative predominance of specific taxa (e.g. Schäfer and Senowbari-Daryan 1978; Schott 1983). Examples from various parts of the Tethys (Alps, Sicily, Greece, Oman) exhibit corresponding association patterns, proving the usefulness of foraminifera for environmental analysis. Foraminiferal distribution was controlled by substrate conditions, specific life behavior, the topography of the depositional environment and water depth. Common Late Triassic foraminifera include taxa of Textulariina (Pl. 111/1, 9, 10, 14, 18, 19, 20), Miliolina (Pl. 111/2-8, 1517, 22, 23), Rotaliina as well as microgranular foraminifera (Pl. 111/21). Note that higher systematic groups of Late Triassic foraminifera have been differentiated more deliberately using wall structure types (Hohenegger and Piller 1975, Piller 1978). Involutinid foraminifera offer the possibility for subdividing restricted and open shelf environments. Jurassic: Benthic foraminifera (Pl. 69) and dasyclad calcareous algae assist in biozonations of Jurassic platform, ramp and upper slope carbonates in the Mediterranean Tethys (Sartoni and Crescenti 1962; Jaffrezo 1980; Sartorio and Venturini 1988; Fig. 10.28A). These ‘coenozones’ are based on various bioevents (e.g. first occurrence and disappearance and total range of taxa, maximum frequency and assemblage patterns; see Sect. 10.3). The first planktonic foraminifera appeared in the Middle Jurassic (Pl. 69/12). Particularly common in thin sections of Jurassic platform carbonates are large lituolinid foraminifera characterized by complex inner structures (Septfontaine
465
1988). In contrast to lituolinid foraminifera without wall differentiation and internal partitions of the chambers, (Pl. 69/8) which were common in open-marine environments, lituolinid foraminifera with complex inner structures were restricted to warm lagoonal environments. Widely distributed examples of these Lituolacea are Lituosepta (Early Jurassic), Orbitopsella (Early Jurassic), Haurania (Early and Middle Jurassic), Meyendorffina (Middle Jurassic), Alveosepta (Late Jurassic, Pl. 69/8), Kurnubia (Middle and Late Jurassic, Pl. 69/ 2, 3, and Middle Jurassic to Early Cretaceous; Pl. 69/ 1). Microforaminifera are linings of juvenile parts of foraminiferal tests corresponding to chitinous membranes, and often stained red by diagenetic impregnation of Fe-oxides (Stancliffe 1989; Misik and Sotak 1998). These very small fossils (<150 µm) are abundant in red Jurassic basinal and deep shelf limestones. They are known from the Permian to the Quaternary. Cretaceous: From the Cretaceous foraminifera spread into the multitudinous number of niches that they also occupy today. Thus, by the Cretaceous, environmental factors can be reasonably delimited from foraminiferal patterns. Benthic foraminifera are important index fossils of the extended Cretaceous platform carbonates. The evolution of planktonic foraminifera during the Cretaceous and in the Cenozoic forms the basis for excellent biostratigraphic zonations and worldwide correlations. Cretaceous benthic foraminifera allow the paleoenvironment to be biostratigraphic zoned and parts of platforms and ramps to be differentiated. Stratigraphic subdivisions are predominantly based on biozones derived from the distribution, abundance and ranges of benthic foraminifera. Oriented sections are important for determining larger foraminifera such as orbitolinids and orbitoidaceans. Agglutinated foraminifera (cuneolinids, dicyclinids, valvulinids and various ataxophragmiaceans; Pl. 71) are common in lagoonal and shallow inner platform environments. The agglutinated orbitolinids (Orbitolinacea; Pl. 70) were common in inner and outer platform settings. They permit particularly Early and Middle Cretaceous shallow-water limestones to be stratigraphically zoned in detail (Schroeder and Neumann 1985; Velic 1988; Clavel et al. 1994; Becker 1999; Husinec et al. 2000). The time intervals correspond to taxon-range zones and assemblage zones. In the Tethys a significant spreading took place during latest Barremian to Early Aptian times, associated with the development of the ‘Urgonian’ platforms (Masse 1976; Peybernes 1976; Arnaud 1981; Arnaud-Vanneau
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et al. 1988), and again in the latest Albian. In this region the group becomes extinct in the Middle Cenomanian. The group includes valuable paleoenvironmental indicators. Orbitolinids lived in shallow warm waters of normal salinity. They occur in different microfacies types formed on shallow lagoonal platforms, sometimes in close association with reefs, and in upper slope settings. Late Cretaceous shelf and shelf slope deposits yield discoidal or lenticular tests of orbitoidacean rotaliinid foraminifera (Pl. 72/5-10). The orbitoidaceans are char-
Jurassic Foraminifera
acterized by radially hyaline and perforated walls, cyclically arranged equatorial chambers and regular lateral chambers. Orbitoidacean foraminifera are abundant in Maastrichtian shallow-water shelf carbonates (see Pl. 72/7-10). This group is used for biostratigraphic zonations (Berthou 1984; Caus 1988; De Castro 1990). Planktonic Cretaceous and Cenozoic foraminifera Planktonic foraminifera (Globigerinacea) are abundant in fine-grained Cretaceous and Cenozoic carbonates deposited at the deep-sea bottom of open-marine Text continued on p.478
Plate 69 Jurassic Foraminifera Many benthic Jurassic foraminifera as well as dasyclad calcareous algae (see 10.2.1.7) are facies-diagnostic fossils and assist in subdividing platform and reefal platform-edge carbonates (Chiocchini et al. 1988). Lituolinid larger foraminifera with complex inner structures (–> 1, 2, 3, 8) are excellent index fossils (Hottinger 1971, Septfontaine 1981) and form the basis for a biozonation of lagoonal platform carbonates. The plate displays some common taxa comprising Textulariina (–> 1-3, 7, 8), porcelaneous Miliolinina (–> 4, 6, 10) as well as Involutinina (–> 5, 6). 1
Pseudocyclammina lituus (Yokoyama). Lituolinid larger foraminifer with complex internal structures. Axial section. Test planspirally enrolled. Note thick honeycombed labyrinthic wall with alveoli and thick septa perforated by large openings. Species range: Middle Jurassic (Domerian) to Early Cretaceous (Valanginian). Late Jurassic (Tithonian): Ernstbrunn, Austria. 2 Kurnubia palastinensis Henson. Lituolinid foraminifera with complex internal structures. Oblique section. Important criteria are the subepidermal network (arrow) and the central column (CC). Species range: Oxfordian to Tithonian, abundant in the Oxfordian. Late Jurassic (Kimmeridgian): Central Apennines, Italy. 3 Kurnubia palastinensis Henson. Transverse section. Diagnostic criteria of the endoskeleton are best shown in sections like this perpendicular to the coiling axis. Late Jurassic (Kimmeridgian): Western High Atlas Mountains, Morocco. 4 Trocholina alpina Leupold. Oblique axial sections of conical tests. Common in platform and reef carbonates. Early Tithonian: Sulzfluh Mountains, Graubünden, Switzerland. 5 Spirillinid foraminifer. The characteristic feature of the involutinid group Spirillilina, particularly common in Liassic deep shelf limestones (Blau 1987), are monocrystalline shells (Hohenegger and Piller 1977). Transverse section of the discoidal test. Late Jurassic (sponge-Tubiphytes reef, Tithonian): Altensteig 1 well, western Molasse Basin, southern Germany. 6 Conicospirillina basiliensis Moehler. Oblique section of the spirillinid foraminifera. Species range: Tithonian to Valanginian. Same locality as –> 4. 7 Ammobaculites Cushman. Textulariina. Lituolacean foraminifera without complex internal structures. Longitudinal section. Free, elongated test. Early portion coiled, later rectilinear. Note coarsely agglutinated wall structure. Genus range: Mississippian to Holocene. Same locality as –> 5. 8 Alveosepta Hottinger. Lituolinid foraminifer with complex internal structures. Section slightly oblique to the coiling axis. Exoskeleton with subepidermal network lining the entire chamber including the septa (arrow). Genus range: Oxfordian to Kimmeridgian. Same locality as –> 3. 9 Lenticulina Lamarck. Rotaliina. Oblique section. Involute planispiral test. Wall calcareous, hyaline, finely perforate. Same locality as –> 5. 10 Protopeneroplis striata Weynschenk. Involutinina. Test enrolled in about two rapidly enlarging and loosely coiled whorls. Wall calcareous lamellar. The species is common in Middle and early Late Jurassic carbonates deposited in shelf-margin and upper slope positions. Same locality as –>4. 11 Frondicularia Defrance. Miliolina, Nodosariacea. Simple uniserial, elongated test consisting of broad, uniform chambers. Sutures are highly arched to angled at the midline of the test. Same locality as –> 5. 12 Planktonic foraminifera: ‘Protoglobigerinid’, an informal name for probably the oldest planktonic foraminifera, are regionally abundant in Callovian and Oxfordian basinal carbonates (Riegraf 1987). Middle Jurassic: Unken, Austria. –> 8: Hüssner 1985
Plate 69: Jurassic Foraminifera
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Cretaceous Larger Foraminifera
Plate 70 Cretaceous Larger Foraminifera: Orbitolinids Orbitolinids (Textulariina, Orbitolinacea) are large agglutinated foraminifera. They comprise important biostratigraphic markers and are used for biozonations, particularly for Barremian, Aptian, Albian and Cenomanian platforms and ramps. These zonations are based on taxon ranges, assemblage composition and the abundance of species. Identifying orbitolinids requires oriented thin sections. The tests consist of calcareous or non-calcareous particles cemented with a mineral cement (–> 3). The conical forms have uniserial stacks of low, saucershaped chambers following an early trochospiral stage. The chambers grow rapidly in diameter and are subdivided by septula on the periphery, often exhibiting a thin marginal zone of minute cellules, and a reticulate zone in which the radial chamberlets are further subdivided by vertical pillars (-> 8, 9). Of diagnostic importance are the shape of the initial and subsequent embryonic chambers (–> 4, 7, 13; see figures at the bottom of this page). 1
Preservation: Tests of orbitolinid foraminifera were frequently transported as shown by the various orientations (see –> 15). A common case is downslope transport from subtidal shelf environments into deeper ramp or basinal settings. Cretaceous (Mishrif Formation, Cenomanian): Subsurface, Ras al Khaimah, United Arabian Emirates. 2 Recognition and composition: Reworked or worn, finely agglutinated orbitolinid tests can be misinterpreted as micrite clasts if the microcrystalline calcite cement is more conspicuous than the fine-grained particles building the wall. Middle Cretaceous (Aptian): Subsurface, Saudi Arabia. 3 Composition of the test: Agglutinated test composed of angular quartz grains (white) cemented by calcite. Albian: Bavarian Alps, Germany. 4 Diagnostic criteria: Taxonomic differentiation of orbitolinids requires sections showing the initial and embryonic chambers that changed in shape and size from the Albian to the Cenomanian. The axial section of the macrospheric embryonic stage of Orbitolina (Orbitolina) duranddelgai Schroeder. Orbitolina (Orbitolina) occurs from the Late Albian to Late Cenomanian. The sample is from the Late Albian of the Sierra de Aulet, Huesca Province, Spain. 5 Orbitolinopsis capuensis (de Castro). Axial section. Genus range: Barremian to Early Aptian. Early Barremian: Djebel Debar, Constantinois, northeastern Algeria. 6 Palorbitolina lenticularis (Blumenbach). Axial section exhibiting the initial chamber. Compare this figure with –> 7 and note the difference in the shape of the initial chamber and the arrangement of the chambers following the initial chambers. The increase in the size of the initial chamber through time can be used for dating (Gusic 1981). Genus and species range: Barremian to Aptian. Early Aptian: Resita Zone, Apuseni Mountains, Romania. 7 Macrospheric embryonic stage of Palorbitolina lenticularis (Blumenbach). Axial section. Early Aptian: Bari, Puglia, Italy. 8 Palaeodictyoconus arabicus (Henson). Axial section. Genus range: Valanginian to Bedoulian. Axial section. BarremianEarly Aptian: Resita Zone, Apuseni Mountains, Romania. 9 Palaeodictyoconus arabicus (Henson). Transversal section. Late Barremian-Early Aptian: Grand Banks, Eastern Canadian Shelf (DSDP leg 43). 10 Palaeodictyoconus arabicus (Henson). Oblique section. Aptian: Subsurface, Ras al Khaimah, United Arabian Emirates. 11 Neoiraqia convexa Danilova. Axial section of the macrosphere. Genus range: Cenomanian. Early Cenomanian: Caserta, Campania, Italy. 12 Orbitolina (Mesorbitolina) texana Roemer. Tangential axial section. Range of Mesorbitolina: Late Aptian to Early Cenomanian. Late Albian: Padurea Craiului, Apuseni Mountains, Romania. 13 Orbitolina (Mesorbitolina) aperta Schroeder. Oblique section of the embryonic macrospheric stage. Lowermost Orbitolina (Mesorbitolina) texana (Roemer). Axial and Cenomanian: Montagne de Tauch, Département Aude, transversal basal sections. France. 14 Orbitolina (Orbitolina) duranddelgai Schroeder. Tangential section of the macrospheric stage exhibiting the embryonic chamber. Same locality as –> 4. 15 Sandy orbitolinid limestone. Ramp environment. Foraminifera were reworked and eroded (compare –> 1 and 2). Late Aptian: Calizas del Busco, Barranco de las Casillas, Requena, Valencia, Spain. –> 4, 5, 9, 11: Courtesy of R. Schröder (Frankfurt).
Left figure: Differences in the morphology of initial and embryonic chambers of orbitolinid foraminifera. After ArnaudVanneau (1980).
Plate 70: Cretaceous Larger Foraminifera
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Cretaceous Benthic Foraminifera
Plate 71 Cretaceous Platform Carbonates: Benthic Foraminifera Thin sections of many Cretaceous platform carbonates are rich in benthic foraminifera, allowing generic or specific identification. Benthic foraminifera are particularly useful for biostratigraphic zonations of Tethyan shallow-marine carbonates (Becker 1999). These biozones exhibit a general trend with regard to the superposition of the zones, but may include different parts of chronostratigraphic units, owing to different paleoenvironmental conditions and regional fluctuations of sea level and subsidence. Conical tests exhibiting chambers subdivided by many septulae and belonging to the textulariid family Cuneolinidae (superfamily Ataxophragmiacea) are index fossils in Early Cretaceous carbonates (–> 1–5, 8, 11). 1
Vercorsella camposauri (Sartoni and Crescenti). Textulariina, Ataxophragmiacea. Transversal-oblique section. The chamber interior is subdivided by radial vertical partitions. Wall agglutinated with fine calcite cement. Vercorsella camposauri and Vercorsella scarsellae (–> 2) characterize an Early Aptian biozone. Genus range: Late Hauterivian to Early Aptian. Barremian-Early Aptian (Berdiga limestone): Kircaova south of Trabzon, northern Turkey. 2 Vercorsella scarsellai (de Castro). Transverse section. Vercorsella characterizes the inner platform environment. Same locality as –> 1. 3 Vercorsella laurentii (Sartoni and Crescenti). Longitudinal-tangential section. Same locality as –> 1. 4 Vercorsella laurentii (Sartoni and Crescenti). Axial section. Same locality as –> 1. 5 Sabaudia minuta (Hofker). Textulariina, Ataxophragmiacea. Subaxial longitudinal section. Wall agglutinated, imperforate. Small conical test. Trochospiral initial stage with inner microgranular and outer hyaline layer. The biserial chambers are subdivided by vertical partitions forming a network of chamberlets. Genus range: Late Hauterivian to Aptian. Late Aptian: Padurea Craiului, Apuseni Mountains, Romania. 6 Dictyoconus algerianus Cherchi and Schroeder. Textulariina, Orbitolinacea. Axial section. Conical test, dish-like chambers with typical convex base. Late Aptian-Early Cenomanian: Gebel Maghara, Sinai, Egypt. 7 Dictyoconus algerianus Cherchi and Schroeder. Oblique section showing the vertical partitions of the marginal zone (arrows). Same locality as –> 6. 8 Cuneolina ex gr. pavonia d’Orbigny. Textulariina, Ataxophragmiacea. Conical test. As shown in this oblique axial section the broad and low chambers are subdivided into chamberlets. Cuneolinid foraminifera are abundant in shallow inner platform deposits associated with miliolinid foraminifera. Genus range: Valanginian to Coniacian. Same locality as –> 6. 9 Reophax giganteus Arnaud. Textulariina, Hormosinacea. Longitudinal section. Free, elongated test with few chambers. The coarsely agglutinated wall consists of ooids (O) and foraminiferal debris (FD). The genus is known from the Middle Ordovician to the Holocene. Early Cretaceous: Subsurface, Hohenlinden-1 well, Molasse zone, Bavaria, Germany. 10 Choffatella decipiens Schlumberger. Textulariina, Loftusiacea. Transversal equatorial section. Planispiral, enrolled test. The wall is differentiated into a peripheral network (top right). The endoskeleton exhibits thick septa pierced by pores. Genus range: Late Jurassic (Oxfordian) to Late Cretaceous (Cenomanian). Common in the Barremian to Aptian time interval. Early Cretaceous: Spain. 11 Dicyclina Munier-Chalmas. Textulariina, Ataxophragmiacea. Axial section. The interior of the discoidal test is subdivided by numerous thin radial partitions, perpendicular to the outer wall and aligned from chamber to chamber. The maximum spreading of the Dicyclina microfacies took place during the Coniacian and the Turonian. Senonian: Subsurface, Ras al Khaimah, United Arabian Emirates. 12 Reophax giganteus Arnaud. Transverse sections. Walls are coarsely agglutinated. Early Cretaceous: Subsurface, Hofolding-1 well, southern Bavaria, Germany. 13 Rhapydionina liburnica (Stache). Miliolina, Milioliacea. Longitudinal section. Genus range: Cenomanian to Maastrichtian. Senonian: Mesarlek, southern Turkey. 14 Rhapydionina liburnica (Stache). This species often dominated an oligotypic assemblage of upper Late Cretaceous inner platform environments characterized by restricted water circulation. Axial section. Same locality as –> 13. Left figure: Schematic sections of Dicyclina. Center: planispiral test. A: Axial section passing through the initial chamber. B: Subaxial section exhibiting several peripheral chambers and radial partitions in the central part of the shell. After Neumann (1967). –>1–4: Kirmaci et al. 1996; 5: Courtesy of I. Bucur (Cluj-Napoca); 6–8: Kuss and Schlagintweit 1988
Plate 71: Cretaceous Benthic Foraminifera
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Late Cretaceous Benthic Foraminifera
Plate 72
Late Cretaceous Benthic Foraminifera
Benthic foraminifera allow the biostratigraphic zonation as well as the facies differentiation of platforms and ramps. Middle and Late Cretaceous miliolinid alveolinids (e.g. Praealveolina, Ovalveolina, Cisalveolina, –> 2) are abundant in inner parts of carbonate platforms (De Castro 1980, 1988). The large discoidal or lenticular tests of the orbitoidacean foraminifera (–> 5-10) are characterized by radially hyaline and perforated tests, cyclically arranged equatorial chambers and more or less regular lateral chambers arranged in annular cycles (–> 7). Orbitoidacean foraminifera are abundant in Campanian and Maastrichtian outer platform and upper slope carbonates. The plate exhibits agglutinated textulariid, miliolinid and rotaliinid foraminifera. 1
2
3
4
Nezzazata Omara. The systematic position of the genus is doubtful. Oblique section of the trochospiral test. Calcitic wall. Genus range: Albian to Turonian. The maximum diversification and spreading of this genus and morphologically related foraminifera occurred in the Cenomanian. The sample is from Cenomanian inner platform deposits in the subsurface of Ras al Khaimah, United Arabian Emirates. Left figure: Exterior view of the trochospiral plano-convex test (after Loeblich and Tappan 1964). Cisalveolina fallax Reichel. Miliolina, Alveolinidae. Right: Oblique equatorial section of a megalospheric test. Left: Tangential section. Note the numerous chamberlets produced by septules. Genus range: Cenomanian. The sample is from the Cenomanian Mishrif Formation: Subsurface, Ras al Khaimah, United Arabian Emirates. Agglutinated foraminifera. Chrysalidina d’Orbigny (bottom, axial section); associated with orbitolinids (top). Note the large initial chamber (arrow) of the conical triserial test and the low and strongly overlapping chambers. Thick, finely agglutinated wall. Time range of Chrysalidina: Cenomanian-Turonian. Same locality as –> 2. Cuneolina d’Orbigny. Textulariina, Ataxophragmiacea. Oblique section. Common in inner shelf platform settings. Genus range: Valanginian to Santonian. Late Cretaceous: Subsurface, Ras al Khaimah, United Arabian Emirates. 5
Siderolites Lamarck. Rotaliina, Rotaliacea. Note the hyaline wall structure and the distinct spines (arrow) arising near the initial chamber in the plane of coiling. White grains are quartz. Common in upper slope environments. Maastrichtian: Subsurface, Saudi Arabia. 6 Lepidorbitoides Silvestri, embedded in an echinoderm packstone. Rotaliina, Orbitoidacea. Oblique section of the flattened, lenticular test. This is an example of poorly preserved tests that are rather common but that can be still determined at higher taxonomic levels, allowing a rough age assignment. Genus range: Campanian to Maastrichtian. Late Cretaceous: Northern Tunisia. The sample is a mosaic stone, used for the decoration of Punian houses in ancient Carthage. Microfacies and age indicates a source of the material within a distance of at least fifty kilometers. 7 Orbitoides d’Orbigny. Rotaliina, Orbitoidacea. Axial section. Taxonomic criteria of these larger foraminifera are the early ontogenetic stages reflected by the initial chamber and the subsequently formed chambers. Siderolites calcitrapoides. Late Maastrichtian Note the differentiation of a median equatorial layer of chambers (E) open marine shelf margin facies characterized by an assemblage of larger foraminifera, dasyand lateral chambers (L). Genus range: Latest Santonian to Maastrichclad algae and rudist debris. Salento Peninsula, tian. Orbitoidacean foraminifera are common in upper slope environPuglia, Southern Italy. Scale is 1 mm. ments. Maastrichtian: Wietersdorf, Carinthia, Austria. 8 Orbitoides. Vertical section. Diagnostic criteria are the arrangement and size of the equatorial chambers and the distribution of the radiating calcite pillars (P) within the lateral part of the test. The pillars give rise to granules on the outer surface. Same locality as –> 7. 9 Orbitoides. Axial section of the megalospheric iuvenarium. The arrow points to the diagnostically important trilocular embryonic chambers, surrounded by a thick perforated wall. Note the equatorial chambers with curved walls and the smaller lateral chambers. Same locality as –> 7. Axial section of Orbitoides displaying the ini- 10 Orbitoides. Vertical section. Note the embryonic chambers and the smaller lateral chambers. Same locality as –> 7. tial chamber and cyclically arranged equatorial chambers, surrounded by rows of lateral chambers. –> 5 and figure on this page: Parente 1997
Plate 72: Late Cretaceous Benthic Foraminifera
473
474
Cretaceous and Tertiary Planktonic Foraminifera
Plate 73 Late Cretaceous and Tertiary Planktonic Foraminifera Planktonic foraminifera are Protista that inhabit the upper zones of the pelagic realms of the oceans. Known since the Middle Jurassic (Triassic?), they are rock-building constituents of Cretaceous and Cenozoic deepmarine limestones, important biostratigraphic markers since the Early Cretaceous, and proxies for paleoceanographic, paleoclimatologic and paleobathymetric reconstructions. About 47 % of the modern sea floor is covered by Globigerina ooze. Planktonic foraminifera are characterized by trochospirally coiled tests with inflated, perforated chambers, sometimes with keeled outer margins (e.g. Globotruncanidae and Globorotaliidae). The wall consists of optically radial and bilamellar low-magnesian calcite. The observation of wall structure and surface ornamentation in thin sections is not practicable in all cases, but is currently refined by the analysis of oriented sections across isolated tests. Widely used in the correlation of Late Cretaceous open-marine carbonates are species of Globotruncana (–> 1, 2, 5, 8), in Cenozoic carbonates Globorotalia, Globigerina (–> 9) and Orbulina (–> 10). Globotruncana is differentiated according to the number of keels that can be recognized in axial sections (singlekeeled –> 8; double-keeled –> 7). 1
Axial and tangential sections of different planktonic foraminifera, including Whiteinella with trochospirally coiled tests and globular chambers. Late Cretaceous (Seewer limestone, Turonian): Glacial boulder, near Bern, Switzerland. 2 The occurrence of planktonic foraminifera (double-keeled Globotruncana – G, Heterohelix – H) associated with abundant ‘calcispheres’ (Pithonella, see Pl. 66/4, 5) indicates an open-marine deep-shelf depositional environment. Turonian (Seewer limestone): Seewen, Switzerland. 3 Planktonic foraminiferal packstone predominantly with Praeglobotruncana. Late Cretaceous: Siliana near Maktaris, Northern Tunisia. 4 Planktonic foraminiferal packstone with Morozovella McGowran. Genus range: Paleocene to Middle Eocene. Paleogene: Central Apennines, Italy. 5 Globotruncana Cushman (G, axial section) and Heterohelix Ehrenberg (H). High-trochospiral globotruncanids are common in Maastrichtian open-marine pelagic limestones. Genus range of Globotruncana: Santonian to Late Maastrichtian. Single-keeled (arrows) globotruncanids as in this sample are typical of a Campanian or Maastrichtian age. Heterohelix: Late Aptian to Paleocene. Maastrichtian: Delphi, Greece. 6 Heterohelix Ehrenberg. Note early planispiral and later biserial growth of the high trochospiral test. The maximum development and spreading of large heterohelicid foraminifera occurred in the Maastrichtian. Late Cretaceous: Northern Tunisia. The sample represents rock material used for Roman mosaics in Carthage. 7 Marginotruncata Hofker. Double-keeled form (arrows). Generic range: Turonian to Santonian. Spreading of doublekeeled planktic foraminifera started in the Turonian and Coniacian. Late Cretaceous: Northern Tunisia. 8 Globotruncana stuartiformis Dalbiez and Globotruncana ex gr. arca (Cushman) – top right. Axial sections of singlekeeled forms. Note the difference between the infillings of the tests and the dense pelagic carbonate mud. Arrows point tp the diagnostically important keels. Early Maastrichtian: Delphi, Greece. 9 Globigerinid wackestone with Globigerinoides. Coiled tests are cut in all planes. Note the globular chambers and the coarsely perforated calcitic walls. Genus range: Eocene to Holocene. Miocene: Southern Turkey. 10 Orbulina (O)-Globigerina (G) packstone. Spherical test. Wall calcareous, with numerous smaller pores interspaced among a few larger pores. Orbulina is common in Middle and Late Miocene deep shelf and basinal limestones. SMF 3-FOR. Miocene: Camposauro, southern Apennines, Italy. Typical planktonic foraminifera. After Ziegler (1983).
Exterior views and axial sections of Globotruncana. Arrows point to keels. After Bignot (1985).
Plate 73: Cretaceous and Tertiary Planktonic Foraminifera
475
476
Tertiary Foraminifera
Plate 74 Tertiary Nummulitid and Alveolinid Limestones Foraminifera are represented in thin sections of Tertiary limestones by benthic and planktonic forms (Pl. 74/9, 10). Of particular interest are benthic larger foraminifera, including the Nummulitacea (–> 2–5; Pl. 75/1, 2, 6-8; hyaline calcareous rotaliinids, common in shallow inner platform settings and sometimes associated with reefs) and the Alveolinacea (–> 1, 9; Pl. 75/3-5; porcelaneous calcareous miliolinids, common in shallow lagoonal platform settings). These groups have millimeter- to centimeter-sized tests with complex endoskeletal structures which must be studied in oriented thin sections. Axial sections show whorls, the character and dimensions of the chambers and the distribution of pillars. Equatorial sections display growth character of whorls, the number and shape of septa and character of initial chambers. Living and extinct taxa evolved in tropical or subtropical habitats. The taxonomic differentiation needs sections showing the size and shape of the test, the morphology of the initial and embryonal chambers, the character of the whorls, shape and number of septa as well as dimensional features. The lenticular and discoidal hyaline Nummulitidae occur in rock-building abundance. A famous example is the carbonate rock used in building the great Cheops pyramide in ancient Egypt. Nummulitids and alveolinids are index fossils for the biozonation of Early Tertiary shallow inner platform and shelf edge carbonates. Other stratigraphically significant groups are the rotaliinids include the Lepidocyclinidae (–> 7–8, Middle Eocene to Early Miocene) and Miogypsinidae (–> 6, Middle Oligocene to Early Miocene). 1
Alveolina d’Orbigny. The Alveolinidae, known from the Early Cretaceous up to today, have imperforate calcareous tests of porcelaneous appearance. Coiling is fusiform to ovate planispiral. Chambers are divided by septulae into numerous chamberlets arranged in one or more rows. Genus range: Late Paleocene to Late Eocene. Early Tertiary (Eocene): Slovenia. 2 Nummulites Lamarck. Nummulitidae are characterized by radial hyaline and perforate tests with septa. Coiling is biconvex planispiral. Chambers may be simple or differentiated into median and lateral layers. The sample is a foraminiferal packstone consisting predominantly of Nummulites. Dense packing, poor sorting and the minor amount of micrite indicate postmortem allochthonous deposition by high energy events. The thin section shows axial sections (A), equatorial sections (E) and random sections. Genus range: Paleocene to Holocene, abundant in Eocene. The maximum distribution of nummulitid facies is in the Middle Eocene. SMF 10. Early Tertiary (Eocene): Southern Bavaria, Germany. 3 Assilina d’Orbigny. Nummulitacea. Large evolute tests with rapidly enlarging whorls and numerous chambers per whorl. Genus range: Paleocene to Holocene, common in Eocene. Early Tertiary (Eocene): Gorica, Slovenia. 4 Discocyclina Gümbel. Nummulitacea. Tests are radial hyaline and perforate. A median layer of chambers is differentiated from lateral chambers. Radiating pillars give rise to granules on the outer surface. Genus range: Late Paleocene to Late Eocene. Example: Discocyclina (D, axial section) associated with Nummulites (N). Early Tertiary (Eocene): Southern Bavaria, Germany. 5 Operculina d’Orbigny. Nummulitacea. Planispiral flattened test. Wall calcareous, lamellar, finely perforate. Note imbrication structure indicating current transport. Genus range: Oligocene to Holocene. Miocene: Ekisce, southern Turkey. 6 Miogypsina Sacco. Rotaliina, Rotaliacea. Transverse sections. The Miogypsinidae are excellent index fossils for the Middle Oligocene to Early Miocene time interval (Abdelghany and Piller 1999). The genus is tolerant of terrigenous influx. It occurs both in lagoon and shallow subtidal environments of open platform. The sample is a foraminiferal packstone exhibiting an accumulation of foraminifera, bryozoan debris (bottom) and red algal fragments (black). Oligocene: Southern Turkey. 7 Lepidocyclina (Eulepidina) ephippioides Jones and Chapman. Rotaliina, Asteriginacea, Lepidocyclinidae. Vertical section of a macrosphere. Note central initial chambers and densely spaced median chambers. The shape of the median chambers is the diagnostic criterion. Lepidocyclinidae are common in lagoons and shallow subtidal environments. Genus range: Middle Eocene to Early Miocene. Oligocene: Argyrotopos, Epirus, western Greece. 8 Lepidocyclina (Eulepidina) eodilatata Douvillé. Vertical section of a microsphere (bottom). Same locality as –> 7. 9 Borelis melo (Fichtel and Moll). Tangential section. Miliolina, Alveolinacea. Small spherical test. Note numerous septula aligned from chamber to chamber. Genus range: Eocene to Holocene. Miocene: Kasaba, southern Turkey. 10 Encrusting foraminifera: Miniacina Galloway, arrows (Rotaliina, Acervulinidae) intergrown with coralline red algae (black). Rasberry-like clusters of globular chambers. Acervulinid foraminifers contribute to the formation of composite rhodoids and reef carbonates. Genus range: Eocene to Holocene. Paleocene: Slovakia. –> 5, 6, 9: Courtesy of M. Link (Erlangen); 7, 8: Bellas et al. 1995. Figure on this page from Bignot (1985)
Plate 74: Tertiary Foraminifera
477
478
Significance of Foraminifera
settings (Pl. 73). Although the taxonomic study of Cretaceous and Tertiary planktonic foraminifera is based on isolated specimens from disintegrated samples (Toumarkine and Luterbacher 1985), thin-section analysis is of great importance, both for the biozonation of basinal and slope carbonates and for the interpretation of the environmental constraints of oceanic water masses. Clues for determining planktonic foraminifera in thin sections can be found in many papers (Caron 1985; Nocchi et al. 1988; Sliter 1989; Wernli et al. 1997; Van Konijnenburg et al. 1998). Cenozoic: Shallow-marine carbonates formed by larger foraminifera are widely distributed in Early Tertiary sequences. Nummulitid buildups form shoals and banks in ramp and platform settings. Their biofabrics are proxies for hydrodynamic processes (Aigner 1985) and show potential trends in reservoir quality within shallow marine shelf sequences (Racey 1994). Accumulations of larger foraminifera are common in Ceno-
Plate 75
zoic shelf and ramp carbonates from North Africa to India, and the western Pacific. Productive nummulitid buildups are known from Tunisia and Libya. The taxonomic distributional patterns of larger and smaller foraminifera assist in recognizing paleoenvironments and settings. But because the taxonomic composition of Early Tertiary foraminiferal faunas differs considerably from living assemblages, paleoecologic interpretations based only on direct comparisons with recent distributional patterns are somewhat hazardous (Hohenegger 1994). Functional morphology produces better results, as shown by the subdivision of different parts of carbonate ramps using the distribution of alveolinids (Gietl 1998). The shape and thickness of the tests as well as distribution trends of modern larger foraminifera provide a basis for a model (Hallock and Glenn 1986) allowing further differentiation of the Standard Facies Zone model (Sect. 14.2). The model was tested for Early Miocene carbonates but can possibly be applied also to Early Tertiary carbonates.
Rock-Building Early Tertiary Benthic Foraminifera
This plate displays thin sections of typical nummulitid and alveolinid limestones. Allochthonous concentrations of rotaliinid and miliolinid larger foraminifera are of rock-building importance. These accumulations are caused by high-energy transport (–> 7, 8) or winnowing of finer material forming lag deposits (–> 5). Parautochthonous concentrations are the result of high production rates and in-place sedimentation. Note the different wall structures of agglutinated Textulariina (–> 9), hyaline calcareous Rotaliina (–> 1, 2) and porcelaneous calcareous Miliolina (–> 3-5). 1 Axial sections of Nummulites Lamarck (N) and Discocyclina (D). Poorly sorted foraminiferal packstone. Outer shelf environment. Early Tertiary (Eocene): Southern Bavaria, Germany. 2 Nummulites. Note the initial chamber (arrow) and the radial fibrous structure of the thick hyaline wall. Early Tertiary: Southern Bavaria, Germany. 3 Alveolina d’Orbigny. Axial section. Note the different wall structure (porcelaneous calcareous) as compared with nummulitid foraminifera (–> 1). The foraminifera are embedded within a calcareous sandstone. Eocene (Lutetian): Southern Bavaria. 4 Alveolina. Equatorial section. Early Tertiary: Divaca, Slovenia. 5 Alveolinid grainstone. The foraminifera association consists predominantly of Alveolina lepidula and miliolid and textulariid foraminifera. Near-coastal sand sheet deposited at the base of a transgressive system originating from reworking of beach wedges during incipient drowning. Early Tertiary (Paleocene, Ilerdian): West of Tremp, northeastern Spain. 6 Sandy glauconite-bearing nummulitid wackestone. Nummulitid tests occur within a fine-grained matrix consisting of angular quartz (white) and micrite clasts. Both large and small tests are well-preserved. Encrustations are absent. These features indicate short transport and rapid embedding within the calcareous sand. SMF 18-FOR. Eocene: Southern Bavaria, Germany. 7 Discocyclina (D) - Amphestigina (A) packstone. Predominantly axial sections. The absence of orientation of the flat discocyclinid tests indicates reworking. Larger black constituents are worn coralline algal thalli (CA). Early Tertiary: Southern Bavaria, Germany. 8 Sandy glauconite-bearing Assilina packstone. Note borings in some tests (arrows). Size ranges of tests give some information on the stratigraphic age. The large size in this sample points to a Middle Eocene age. Early Tertiary (Adelholzen beds, Middle Eocene): Agathazell, southern Bavaria, Germany. 9 Encrusting foraminifera: Haddonia heissigi Hagn (H) associated with miliolids (M) and textulariids (T). Haddonia and related agglutinated foraminifera are important members of the constructor and binder guilds of Early Tertiary reefs (Darga 1990). Genus range: Late Cretaceous to Eocene. Eocene: Coral patch reef Eisenrichterstein near Reichenhall, southern Bavaria, Germany. –> 6–8: Courtesy of H. Hagn (München)
Plate 75: Tertiary Foraminifera
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480
Significance of foraminifera in microfacies analysis Biozonations: Benthic foraminifera, particularly larger foraminifera, are reliable tools in the stratigraphic subdivision and correlation of Late Paleozoic, Mesozoic and Tertiary shelf carbonates. Zonations have been established both for low and high latitude deposits. Biozones of different latitudes have different taxonomic spectra. Planktonic foraminifera are important in the interregional correlation of Cretaceous and Cenozoic strata (Bolli et al. 1985). A useful summary of the timing of significant bioevents and changes in platform and open-marine pelagic Mesozoic and Cenozoic biofazies types characterized by the appearance, disappearance or abundance of thin-section foraminifera was given by Sartorio and Venturini (1988). Paleoenvironment: Abundance, diversity and composition of benthic foraminifera differ in modern subtropical, warm temperate and cool temperate shelf seas. Composition patterns of shelf sediments (e.g. ‘foramol facies’) as well as taxonomic differences in foraminiferal assemblages provide guidelines for distinguishing ancient shelf and ramp carbonates deposited in warm or cold waters and in low or high paleolatitudes (see Sect. 12.2). Frequency distribution, species association and diversity of larger foraminifera are strongly correlated with environmental factors changing along depth gradients. The resulting patterns can be successfully used in assessing paleodepths, if the coenocline concept is applied (Hohenegger 2000). Depositional setting and facies zones: The distributional patterns of benthic foraminiferal associations assist in recognizing and subdividing depositional settings and environments of ancient shelves, if depth transport from shelf to basinal environments is considered (Hohenegger and Yordanova 2001). Important papers are included in the list of Box 10.4. Fig. 10.28 gives a very generalized idea of the occurrence and abundance of common foraminiferal groups in Late Jurassic and Eocene carbonate environments. See Sartorio and Venturini (1988) for other Mesozoic and Tertiary distribution patterns of thin-section fossils. Carbonate production: Many tropical beaches are composed almost entirely of tests from shallow-water benthic foraminifera. Benthic foraminifera contribute to the formation of carbonate grains (e.g. oncoids, macroids), reefs and shelf carbonates. The proportion of modern benthic foraminifera is approximately 5% of the global carbonate reef budget (Langer et al. 1997). This is low as compared with the high calcium carbon-
Larger Foraminifera in Limestones
ate production of larger foraminifera in the Cretaceous and Tertiary (Hohenegger 1999). Reef-adapted encrusting foraminifera are known from Permian, Jurassic and Early Cretaceous reef limestones. Monospecific biostromes built by arenaceous foraminifera, and surrounded by bioclastic grain- and wackestones, occur in the Barremian of Haute Savoie, France. These up to one meter high buildups were formed in shallow water on a carbonate platform (Schulte et al. 1993). Other foraminiferal reefs have been reported from the Eocene (Solenomeris reefs: Plaziat and Perrin 1992) and the Miocene (Jell et al. 1965; Ghose 1977; Hallock 1985; Hallock and Glenn 1986). Larger foraminifera have been common constituents of shallow-marine shelf carbonates formed in warmwater environments since the Late Paleozoic. Extensive Cretaceous and Cenozoic shelf carbonates were produced by large, benthic foraminifera. Larger foraminifera (e.g. Amphistegina and Calcarina) contribute significantly to the formation of sand-sized carbonate sediments in reefs and other shallow, carbonate environments due to the high turnover rates of the populations (Hallock 1981). Microfacies implications of rock-building abundance of larger foraminifera: Rock-building concentrations of foraminifera (packstones, grainstones) are caused by • hydrodyamic processes: wave action hindering the deposition of fines between the tests or winnowing fines and smaller tests; current transport accumulating tests in protected or deeper environments; storms inducing deposition near beaches or in shoals on the shelf; turbidity currents leading to deposition on slopes or in basins, • high production rates, e.g. in sea grass or sea weed areas, associated with rapid deposition of epiphytic foraminifera forming washover deposits or in-place deposits (Davaud and Septfontaine 1995), and • bioturbation, e.g. by burrowing echinoids. In addition, the possibility of ‘diagenetic packstone’ or ‘condensed grainstone’ fabrics must be considered. Biofabric analysis of foraminiferal accumulations allows the hydrodynamic and sedimentologic history of foraminiferal limestones to be reconstructed (Aigner 1983, 1985; Fig. 5.3). Important criteria for the description of biofabrics are packing, sorting and size of foraminiferal tests, imbrication or bedding structures, and the amount of carbonate mud preserved in interparticle voids. Worn tests can indicate continued transport, as-
Distribution Patterns of Foraminfera
481
A
B
Fig. 10.28. Generalized patterns of the distribution of common foraminifera in different sedimentary environments. A: Late Jurassic. Protected muddy inner platform sediments yield many miliolid foraminifera. Agglutinated larger foraminifera, particularly lituolids with complex inner structures (e.g. Alveosepta, Pseudocyclammina) and Kurnubia were common in energetically different parts of the inner platform. Trocholinid foraminifera were abundant near or within platform-margin buildups (e.g. reefs formed by corals, ellipsactiniids and many encrusters). Upper slope environments are characterized by hyaline foraminifera, such as Conicospirillina or Protopeneroplis. Planktonic foraminifera appear in the late Middle Jurassic (this picture) and may be found also in Late Jurassic pelagic deep-water limestones, which are otherwise typified by radiolarians and other planktonic microfossils (pelagic crinoids and bivalves). B: Early and Middle Eocene. Shallow inner platform areas were inhibited by low-diversity small miliolinids as well as large alveolinids and discocyclinids. Thick-shelled nummulitids lived in shallow waters both in inner and outer platform or ramp settings. Reefs formed by corals and corallinacean red algae contain agglutinated encrusting foraminifera. Agglutinated foraminifera are also common in deeper platform environments and may occur together with planktonic foraminifera. Planktonic globigerinids and globorotaliids contributed to the formation of basinal carbonates.
Radiolaria in Thin Sections
482
sociated with reworking and redeposition (see Pl. 74/ 6, Pl. 118/2). Differences between the infill of worn foraminiferal sediment and the embedding matrix as well as rotated geopetals are criteria in assessing transport processes. Processes leading to the formation of Early Tertiary grain-, pack- and wackestones with nummulitid foraminifera can be derived from the abundance and proportions of large and small tests, their orientation and the amount of carbonate mud occurring together with the foraminifera. Large and small tests reflect sexual and asexual reproduction phases of the life cycle of the foraminifera. Assuming that within natural populations small foraminifera (A-form, with large initial chambers, megalospheric stage) dominate over large tests (B-form, with small initial chamber, microspheric phase), four assemblages can be differentiated that describe the autochthony or redeposition of the foraminiferal tests: 1 Relatively undisturbed assemblages: A/B ratio of the same species 10:1 or more; matrix-support common, wacke- and mudstones; back-bank environment, 2 Parautochthonous assemblages: A/B average 7:1, poor sorting, packstones; most fines are removed by in-situ winnowing caused by wave action; abundant in bank facies, 3 Residual assemblages: Layers enriched with B-forms but A-forms still commonly present; dominance of imbricated B-forms suggests selective removal of smaller A-forms and enriching of B-forms, creating a residual lag, 4 Allochthonous assemblages: Composed of almost only A-forms, indicating sorting and selective transport. Basics: Foraminifera Benthos ‘86, 3ème Symposium International sur les Foraminifères Benthiques (1988). – Revue de Paléobiologie, Volume spécial, 2, two volumes, 1005 pp., Genève. Bolli, H.M., Saunders, J.B., Perch-Nielsen, K. (1985): Plankton stratigraphy. – 1032 pp., Cambridge (University Press) Haq, B.U., Boersma, A. (1998): Introduction to marine micropaleontology. 2nd edition. – 376 pp., Amsterdam (Elsevier) Hemleben, C., Spindler, M., Anderson, O.R. (1989): Modern planktonic foraminifera. – 363 pp., New York (Springer) Jones, R.W. (1995): Micropaleontology in petroleum exploration. – 432 pp, New York (Clarendon Press) Lee, J.L., Anderson, O.R. (eds., 1991): Biology of foraminifera, 368 pp., Harcourt (Acad. Press) Loeblich, A.R.Jr., Tappan, H. (1964): Protista 2. Sarcodina chiefly ‘Thecamoebians’ and Foraminiferida. – Treatise on Invertebrate Paleontology, C. – 2 volumes, 900 pp., Lawrence (Geological Society of America) Loeblich, A.R.Jr., Tappan, H. (1988): Foraminiferal genera
and their classification. – Vols. 1 and 2, 970 pp., 847 Pls., New York (Van Nostrand Reinhold) Murray, J.W. (1991): Ecology and paleontology of benthic foraminifera. – 397 pp., London (Longman) Neumann, M. (1967): Manuel de micropaléontologie des foraminifères (systématique-stratigraphie). – 297 pp., Paris (Gauthier-Villars) Further reading: K111. Relevant papers dealing with foraminifera studied in thin sections of Devonian to Tertiary limestones are listed in Box 10.4.
10.2.2.2 Radiolaria Radiolarians are single-celled marine planktonic protozoans that secrete an amorphous (opaline) silica skeleton composed of a number of architectural elements (radial spicules, internal bars, external spines) that are joined together to form very regular and symmetrical lattices. The shell has a size of usually less than 2 mm and commonly between 100 and 250 µm across. The unicellular body is divided by a membrane into an internal (intracapsular) part (central capsule containing the endoplasm and the nucleus) and the surrounding external (extracapsular) part containing the peripheral cytoplasma. Radiolarians may be solitary or colonial. Morphology and classification: The shell morphology is highly variable but typically corresponds to a hollow sphere or vase containing one or more bars across the interior. Some forms consist only of simple tetraradiate or multiradiate branching needles. Radiolaria are a subclass of the class Actinopoda that belong to the subphylum Sarcodina and the phylum Protista. Modern radiolarians are classified using features of both the skeleton and the soft parts. Separate classifications have been established for Paleozoic, Mesozoic and Cenozoic radiolarians. Recent radiolarians comprise two major groups: (1) the polycystines, with solid skeletal elements of opaline silica, and (2) the phaeodarians with hollow siliceous skeletal elements connected by organic material, which are rapidly dissolved in seawater, and consequently rarely preserved in the sediments. The polycystines are subdivided into (1) the basically spherical-shaped Spumellaria, and (2) the basically conical-shaped Nassellaria (Fig. 10.29). Some major groups of extinct radiolarians differ substantially from Spumellaria and Nassellaria and are therefore considered as separate taxonomic units. Spumellarians are characterized by a pronounced radial symmetry and single or multiple concentric shells connected by radial bars. The skeletons have various shapes ranging from spherical to ellipsoidal to discoidal. Nassellarian
Diagenesis of Radiolarians
483
Fig. 10.29. Morphology of spumellarian (A, B) and nassellarian (C-F) radiolarians. A - Archaeodictyomitra sixi Yang (Tithonian: Longing Gap, Antarctic Peninsula; height 200 µm), B - Mirifusus minor Baumgartner (Berriasian: Jebel Buwaydah, Oman Mountains; height 500 µm), C - Higumastra sp. (same locality as A; maximum diameter 500 µm), D - Triactoma kellumi Pessagno and Yang (Tithonian: Southern Alps, Italy; diameter of cortical shell 200 µm), E - Acastea diaphorogona (Foreman) (same locality as B; diameter of cortical shell 130 µm), F - Leugeo hexacubica (Baumgartner) (Callovian: Southern Alps, Italy; pyritic specimen; diameter 175 µm). Courtesy of W. Kiessling (Berlin).
shapes derive from geometrically simple ringlike spicules from which latticed covers are added to form chambers. Methods: Radiolaria are usually studied by separation techniques, but also in thin sections. The fossils are investigated under the light microscope and in SEM. Cathodoluminescence microscopy enhances thin-section criteria. Radiolaria in thin sections: Radiolarians are common to abundant in thin sections of fine-grained openmarine pelagic limestones and radiolarites. Sections through spherical and ellipsoidal tests of spumellarians result in circular figures of different sizes (Pl. 76/1). Circular figures can also be cross sections of nassellarians. Differentiating these two groups in thin sections is in most cases not possible apart from rare exceptions (Pl. 76/4). Calcitized radiolarian tests exhibit some similarity with calcispheres (Sect. 10.2.1.9), but can be differentiated by a toothed appearance of the peripheral zone of the shell (Pl. 76/2). Unlike from radiolarian tests, the sections of calcispheres commonly show defined calcareous walls. Sections of calpionellids are considerably smaller than those of radiolarians (Pl. 76/8). Biology and ecology: Radiolarians are a part of the zooplankton and capture food from the water column. They consume various zooplankton and phytoplankton as well as bacteria and organic detritus. Some polycystine radiolarians host photosymbiotic dinoflagellate symbionts within their cells. Modern radiolarians are widely distributed in the oceans and occur in the photic surface zone of the ocean as well as in depths of hundreds of meters down to water depths of 4000 m and more. About the same number of modern radiolarian
species occur in tropical and subtropical surface waters and in deeper waters (> 200 m). Fewer species live in high latitudes. Homogeneity in radiolarian assemblages parallel to the latitudes, and heterogeneity across the latitudes indicate that water temperature influences the distributional patterns. Variations in the composition of assemblages depend on the difference of physical, chemical and biological parameters within water masses, e.g. oceanic currents. The composition of radiolarian assemblages and their distributional patterns, therefore, are excellent proxies for paleoclimatic and paleooceanographic interpretations of Meso- and Cenozoic oceans. Distribution: First records are known from the Cambrian. By the Silurian deep-water forms are believed to have evolved. All Early Paleozoic radiolarians are spumellarians or belong to extinct subgroups. Mesozoic and Cenozoic assemblages are characterized by various groups of spumellarians and nassellarians. A rapid diversification took place at the end of the Jurassic. Warm-water faunas have existed since the Cambrian. Cold-water assemblages appear in the Oligocene. Diagenesis and preservation: The fossil radiolarian record is strongly controlled by diagenetic factors. Dissolution of radiolarian tests takes place in the water column, and on and within the upper centimeters of the sea bottom. It depends on the degree of silica undersaturation of pore waters, the intensity of bioturbation and the accumulation rate. High sedimentation rates favor good preservation. The transformation of siliceous muds into siliceous rocks, connected with the maturation of the sediments corresponds to a sequence starting with opal A, followed by opal-CT (porcelanite) and ending with chalcedony or cryptocrystalline quartz (chert). Mesozoic
484
radiolarians consist exclusively of chalcedony or microquartz. The mineralogical composition seems to depend on the time involved in the long-term transformation processes, and is influenced by the lithology of the radiolarian-bearing sediments. Burial transformation of radiolarians in limestones is said to be much faster than in argillaceous sediments. In cherts radiolarian tests are commonly better preserved than in carbonate rocks, but exhibit a wide spectrum in regard to the preservation of morphological structures. This spectrum can be used to estimate recrystallization, corrosion and metasomatic alterations (Kiessling 1996). The preservation of radiolarian tests in limestones is generally poor as compared with radiolarians in cherts or radiolarites, and in concretions. The siliceous radiolarian tests are often replaced by calcite, pyrite or other ore minerals. Calcitization is common in limestones, but also occurs in radiolarites. It leads to poor preservation, which
Radiolarians and Sedimentation
is evident both in SEM studies of extracted specimens and in thin sections. The replacement of the siliceous skeletons starts as an early diagenetic process but continues under shallow burial conditions (Faupl and Beran 1983). Pyritization is an early diagenetic process that takes place in reducing diagenetic environments. It affects siliceous skeletal grains in different intensities, siliceous sponge spicules more frequently than radiolarians, probably because of their higher organic content, which in turn controls microbial sulphate reduction and the accumulation of iron sulfides at organics. Microfacies of radiolarian-bearing limestones Radiolarians occur predominantly in micritic carbonates and in marls. They are found in lime mudstones, wackestones and packstones. Common Jurassic microfacies types are (Kiessling 1996): • Pelagic and hemipelagic marly mudstones with radiolarians and sponge spicules.
Plate 76 Radiolarians and Calpionellids The plate displays Late Jurassic planktonic organisms represented by radiolarians and calpionellids, a group of the ciliate Protista. 1 Pelagic radiolaria packstone. Note the differences in size and preservation of radiolarians. Calcitized spumellarian radiolarians are recorded by circular sections of different size. Sponge spicules (arrows) exhibit parallel orientation, indicating the influence of bottom currents. The dark micritic matrix yields abundant coccolithophorids, as proved by SEM studies. Using the abundance of radiolarians and sponge spicules and the radiolaria/spicule ratio within a standard field of the thin section as a measure for the depositional setting (Kiessling 1996), sedimentation in the distal parts of a deep basin influenced by bottom currents appears to be likely. The sample comes from the Oberalm limestone, a representative of widely distributed pelagic carbonates, better known as Aptychus limestone in the Alpine region, and Biancone limestone in the Mediterranean area. These limestones are commonly associated with radiolarites. The limestone beds are often separated by marls and contain some cherts and intercalations of allochthonous limestones consisting of material derived from the nearby carbonate platforms. Radiolarians, sponge spicules, thin shell fragments (filaments) and tiny bioclasts form wackestone and packstone fabrics. SMF 3-RAD. Late Jurassic (Tithonian): Hallein, Salzburg, Austria. 2 Spumellarian radiolaria. The original opaline silica of the spherical test has been replaced by calcite, obscuring structural details. The pores on the surface of the test were partially infilled by carbonate mud, giving the margin a toothed appearance. This feature differentiates circular sections of radiolarian casts from sections of calcispheres or globular foraminifera. Same locality as –> 1. 3 Acanthocircus sp., a member of associations which occur in relatively shallow depths. Late Jurassic: Kraiburg, Bavaria, Germany. The sample is a white limestone, used in Roman mosaics of southern Bavaria, Germany. 4 Radiolaria packstone. Limestone concretion in mudstones. Concretions are protected with regard to aggressive pore waters and compaction and can, therefore, yield well-preserved microfossils. The densely packed radiolarians are multicyrtoid nassellarians. The sample comes from an oxygen-deficient basin in a high paleolatitudinal position. Late Jurassic (Nordenskjöld Formation): Longing Gap, Graham Land, Antarctic Peninsula. 5 Tritrabs ewingi Pessagno. The radiolarian skeleton is replaced by sparry calcite but has still preserved its morphology. Same locality as –> 3. 6 Podobursa sp. The nassellarian skeleton is preserved as microcrystalline calcite. Same locality as –> 3. 7 Calpionellid wackestone. Arrows point to the few longitudinal sections showing the criteria necessary for the taxonomic determination of calpionellids (Calpionella sp.). Well-bedded micritic limestones with abundant remains of extinct ciliates are common in Late Jurassic and Early Cretaceous open-marine basinal settings. SMF 3-CALP. Late Jurassic (Tithonian): Seebühel, Switzerland. 8 Calpionellid-radiolaria wackestone. C: Transversal and longitudinal sections of calpionellids; R: Recrystallized radiolaria. Note the difference in size and preservation between radiolarians and calpionellids. Locality as –> 7. –> 3, 5, 6: E. Flügel and Ch. Flügel 1997, 4: Kiessling 1996
Plate 76: Radiolarians and Calpionellids
485
486
• Hemipelagic radiolaria wackestones with common to abundant, partly or totally calcitized radiolarians and sponge spicules as well as benthic foraminifera. Samples contain up to 20% particles. This microfacies is common in carbonates formed in slope position. • Pelagic radiolaria packstones with high amounts of radiolarian tests consisting of microcrystalline calcite, and spicules (Pl. 76/1). Peloids and tiny micritic intraclasts are common. Samples contain up to 50% particles. This microfacies is common in current-influenced basinal settings. • Hemipelagic filament-radiolarian wacke- and packstones with equal amounts of radiolarians and thin bivalve shells, spicules, and remains of pelagic crinoids. Samples contain up to 60% particles. The microfacies occur in basinal settings, influenced by bottom currents. • Microbioclastic/lithoclastic packstones with radiolarians, spicules, filaments, benthic foraminifera and echinoderms, abundant very small lithoclasts, and with up to 70% particles. These microfacies are common in basinal settings; they may be a part of distal turbidites. These or similar microfacies types are also known from Triassic basinal limestones. In Cretaceous limestones, the tests of radiolarians are often overshadowed by abundant planktonic foraminiferal tests. Significance of radiolarians Radiolarian assemblages are important biostratigraphic, paleoecological and paleoenvironmental tools and contribute greatly to the formation of marine sediments. Biozonations: Whole-body radiolarians, extracted from limestones and cherts, have proved to be very useful for biostratigraphic subdivisions, particularly of Mesozoic and Cenozoic open-marine sedimentary rocks (Casey 1993; Jud 1994; Baumgartner et al. 1995). Contribution to sedimentation: Radiolarians accumulate on the sea bottom as a consequence of sinking from the overlying water after the death of the organisms. Only a fraction of the living population is preserved at the sea bottom because silica-deficient seawater dissolves the skeletons during sinking and at the sea water-sediment interface before burial. Radiolarian tests deposited on the sea bottom may be affected by contourite bottom currents. Many radiolarian tests occur in distal turbidites. Radiolarian-rich sediments are formed when the input of non-siliceous components is scarce, so that radiolarians are concentrated in the bottom sediment. Radiolarian skeletons are a major source of silica. Radiolarian accumulation contributes
Significance of Radiolarians
to the formation of cherts and radiolarites (characterized by at least 30% of radiolarian material). Both lithologies are commonly associated with fine-grained sediments, including biogenic pelagic carbonates, marls and terrigenous mudstones. Modern radiolarians are abundant in deep-sea sediments, particularly in Pacific equatorial regions where productivity is high in the overlying water column. The accumulation of radiolarian mud in high latitude areas is due to high productivity rates, deposition at great depths below the CCD, and low dissolution of radiolarian tests. Clues to depositional settings: Radiolarians are common to abundant in thin sections of Mesozoic limestones formed on open-marine slopes and in basins of differing bathymetry. Limestones and cherts with abundant radiolarians are generally thought to represent deep-marine deposits. Paleozoic and Mesozoic radiolarians sediments include, however, a wide range of facies, which vary strongly in composition and texture, fossil content, bedding style, sedimentary structure and rock color (O’Dogherty et al. 2001). Many radiolarian-rich rocks originated in deep-water upper slope or basinal positions, but reports describing the association of shallow-water facies and radiolarian limestones (Blendinger 1985) as well as the occurrence of radiolarians in lagoons should be considered in facies interpretations (discussion in Racki and Cordey 2000). Discrimination of proximal versus distal basin facies: Radiolarian abundance, the nassellarian/spumellarian ratio and the frequency of sponge spicules in the siliceous microfauna are used for characterizing bathyal and abyssal sediments (Kiessling 1996). Nassellarians increase in relative abundance in deeper-water sediments, whereas the relative abundance of sponge spicules is inversely correlated with the bathymetry or distance from the shelf. Spicules only make significant contributions to siliceous microfaunas in slope and proximal basin environments. Thin-section analysis exhibits comparable values regarding the abundance of radiolarians and the proportions of radiolarians and sponge spicules, as does the analysis of isolated extracted radiolarians. Abundance can be better evaluated from thin sections than from extracted faunas. Basics: Radiolaria Blendinger, W. (1985): Radiolarian limestones interfingering with loferites (Triassic, Dolomites, Italy). – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1985, 193-202 Casey, R.E. (1993): Radiolaria. – In: Lipps, J.H. (ed.): Prokaryotes and protists. – 249-284, Boston (Blackwell) Faupl, P., Beran, A. (1983): Diagenetische Veränderungen an
Calpionellid Biozonation
487
Fig. 10.30. Calpionellid limestone. The pelagic limestone displays numerous cross sections and some longitudinal sections. Only a small number of the latter shows the diagnostic criteria necessary for taxonomic assignments. The thin section yields three species: Calpionella alpina Lorenz (C; see Pl. 77/6, 7), Crassicollaria parvula Remane (CR; Pl. 77/4), and Remaniella sp. (R). The assemblage indicates the Calpionella zone, more specifically the elliptica subzone (D2 interval) of the global standard calpionellid biozonation (see Fig. 10.32). Early Cretaceous (Berriasian): Alpe Ra Stua near Cortina d’Ampezzo, northern Italy. Courtesy of J. Blau (Giessen). Radiolarien- und Schwammspicula-führenden Gesteinen der Strubbergschichten (Jura, Nördliche Kalkalpen, Österreich). – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1983, 129-140 Kiessling, W. (1996): Facies characterization of Mid-Mesozoic deep-water sediments by quantitative analysis of siliceous microfaunas. – Facies, 35, 237-274 Kling, S.A. (1978): Radiolaria. – In: Haq, B.U., Boersma, A. (eds.): Introduction to marine micropaleontology. – 203-244, Elsevier (New York) O’Dogherty, L., Martin-Algarra, A., Gursky, H.J., Aguado, R. (2001): The Middle Jurassic radiolarites and pelagic limestones of the Nieves Unit (Rondaide Complex, Betic Cordilleres): basin starvation in a rifted marginal slope of the western Tethys. – International Journal of Earth Sciences, 90, 831-846 Racki, G., Cordey, F. (2000): Radiolarian palaeoecology and radiolarites: is the present the key to the past? – Earth-Science Review, 52, 83-120 Further reading: K112
10.2.2.3 Calpionellids Calpionellids are extinct, bell-shaped very small calcitic microfossils (length about 50–200 µm) that occur abundantly in fine-grained, pelagic limestones of Late Jurassic and Early Cretaceous Tethyan oceans. The group suffered a dramatic decline within the Early Cretaceous age. The planktonic fossils are considered as a group of the subphylum Ciliophora and the class Ciliata. Ciliates are distinguished from other protozoans in having numerous tiny flagella (cilia) serving both for locomotion and food gathering. Calpionellids are regarded as a group of unknown systematic position (Remane 1978), or, more commonly, they are attributed to the Tintinnida with which they have strong morphological similarities and similar sizes, but differ in the material used for building the tests. Modern tintinnids are an important component of the microscopic zooplankton living predominantly in the photic zone of the oceans, but also occur with few taxa in estuaries and lakes. The bell-shaped cells
Fig. 10.31. Longitudinal sections and reconstructed morphology of calpionellids. Diagnostic criteria are the morphology of the collar, the type of the aboral pole, and the shape and dimensions of the test. After Brasier (1980).
488
Calpionellids
Plate 77 Calpionellids Calpionellids are very small bell-shaped, amphora-like or cylindrical microfossils restricted to Late Jurassic to lower Early Cretaceous pelagic limestones. The plate displays longitudinal sections of taxa used in the standard biochronology of the Jurassic-Cretaceous time interval (see Fig. 10.32). 1
Chitinoidella boneti Doben. The test has a microgranular calcitic wall and appears dark in transmitted light due to an ultrastructure that differs from almost all other calpionellids (collected in the family Calpionellidae). Middle Tithonian: Pontaux (Drôme), France. 2 Crassicollaria intermedia (Durand-Delga). Arrows point to the diagnostic criteria of the genus (thickening at the base of the collar caused by oral ring). Late Tithonian, interval A2: Serres-Montclus, France. 3 Crassicollaria brevis (Remane). Late Tithonian, interval A3: Same locality as –> 1. 4 Crassicolaria parvula Remane. Early Berriasian, interval B: Chevallon, France. 5 Praetintinnopsella andrusovi Borza. Tithonian (Chitinoidella zone): Pointaix (Isère), France. 6 Calpionella alpina Lorenz. Black arrows point to the short and erect collar characteristic of Calpionella. White arrow indicates the bell-shaped aboral part of the test. Lowermost Berriasian, lower part of interval B: Saint-Penteraix, France. 7 Calpionella alpina Lorenz. Large specimen, characteristic of time interval A2/A3. Late Tithonian, interval A2: Same locality as –> 2. 8 Calpionella elliptica Cadisch. Berriasian, interval C: Same locality as –> 2. 9 Remaniella ferasini (Catalano). Early Middle Berriasian, Calpionella zone, interval B: Ra Stua section, Belluno, Italy. 10 Remaniella catalanoi Pop. Early Middle Berriasian, Calpionella zone uppermost interval B or lowermost interval C: Same locality as –> 9. 11 Remaniella duranddelgai Pop. Early Middle Berriasian, Calpionella zone, interval B: Same locality as –> 9. 12 Remaniella colomi Pop. Early/Middle Berriasian. Calpionella zone, interval C: Same locality as –> 9. 13 Remaniella colomi Pop. Early/Middle Berriasian. Remaniella is characterized by a collar consisting of three elements (arrows). Calpionella zone, interval C: Same locality as –> 9. 14 Remaniella cadischiana (Colom). Middle Berriasian, Calpionella zone, interval C. Same locality as –> 9. 15 Calpionellopsis simplex (Colom). The arrow points to the simple ringlike collar surrounding the oral opening. Late Berriasian, Calpionellopsis zone, interval D: Majorca, Spain. 16 Calpionellopsis oblonga (Cadisch). Late Berriasian, interval D3: Berrias, France. 17 Sturiella dolomitica Grün and Blau. Late Berriasian, interval D1, Calpionellopsis zone, simplex subzone: Locality as –> 9. 18 Sturiella oblonga Borza. Late Berriasian. Calpionellopsis zone, interval D2; oblonga subzone: Same locality as –> 9. 19 Borzaites atavus (Grün and Blau). The arrow points to the ringlike collar. Holotype. Late Berriasian. Calpionellopsis zone, interval D2, oblonga subzone: Same locality as –> 9. 20 Praecalpionellites hillebrandti Grün and Blau. Praecalpionellites is characterized by a two-fold collar with both elements separated from the lorica (arrow). Late Berriasian. Calpionellopsis zone, oblonga subzone, interval D3. Same locality as –> 9. 21 Cone-in-cone stacking: Outer form is Praecalpionellites murgeanui (Pop), smaller inner form is ?Lorenziella sp. Late Berriasian. Calpionellopsis zone, oblonga subzone, interval D3: Same locality as –> 9. 22 Praecalpionellites siriniaensis Pop. Early Valanginian, Calpionellites zone, interval E: Same locality as –> 9. 23 Praecalpionellites filipescui (Pop). The nannomicrite consists of disintegrated elements of Nannoconus (inset bottom right). Nannoconid assemblages are used to establish biozones. Late Berriasian. Calpionellopsis zone, oblonga subzone, interval D3: Locality as –> 9. 24 Praecalpionellites dadayi (Knauer). Late Berriasian. Calpionellopsis zone, oblonga subzone, interval D3: Locality as –> 9. 25 Calpionellites darderi (Colom). The arrow points to a diagnostic feature of Calpionellites, the arrangement of parts of the collar parallel to the lateral walls. Early Valanginian, Calpionellites zone, interval E: Fanes area, Italy. 26 Calpionellites darderi (Colom). Early Valanginian, Calpionellites zone, interval E: Same locality as –> 25. 27 Calpionellites major (Colom). Early Valanginian, Calpionellites zone, interval E: Same locality as –> 25. 28 Calpionellites coronatus Trejo. Early Valanginian, Calpionellites zone, interval E: Same locality as –> 25. 29 Tintinnopsella carpathica (Murgeanu and Filipescu). Early Valanginian, Calpionellites zone, interval E: Same locality as –> 9. 30 Tintinnopsella carpathica (Murgeanu and Filipescu). Black arrows point to the flaring collar, the white arrow to the caudal appendix. Berriasian, interval C: Berrias, France. 31 Tintinnopsella longa (Colom). Early Valanginian, Calpionellites zone, interval E: Same locality as –> 25. 32 Tintinnopsella longa (Colom). Berriasian, interval D2: Berrias, France. 33 Lorenziella plicata Remane. Berriasian, upper part of interval D2: Same locality as –> 32. 34 Lorenziella hungarica Knauer and Nagy. Berriasian, interval D3: Same locality as –> 32. 35 Lorenziella hungarica Knauer and Nagy. Early Valanginian, Calpionellites zone, interval E: Same locality as –> 9. 36 Lorenziella ruggierii (Catalano). Early Valanginian, Calpionellites zone, interval E: Same locality as –> 9. –> 1–8, 15, 16, 30–34: Courtesy of J. Remane (Neuchâtel); 9–14, 17–29, 35, 36: Courtesy of J. Blau (Giessen)
Plate 77: Calpionellids
489
490
Significance of Calpionellids Grün and Blau 1997 Tintinnopsella
Pl. 77/27
darderi
Pl. 77/25-26
dadayi
Pl. 77/24
murgeanui
Pl. 77/21
filipescui
Pl. 77/23
oblonga
Pl. 77/16
simplex
Pl. 77/15
cadischiana
Pl. 77/14
elliptica
Pl. 77/8
alpina
Pl. 77/6-7
catalanoi
Pl. 77/10
intermedia
Pl. 77/2
Calpionellites 2
1
C
B
3 A
Jurassic Tithonian
major
2 1
Calpionellopsis
Berriasian
D
gr. hungarica Pl. 77/34-35
Calpionella
3
gr. carpathica Pl. 77/29-30
Crassicollaria
E
Chitinoidella
Cretaceous
Valanginian
Remane 1971
Significance of calpionellids Biozonation: Calpionellids are excellent index fossils enabling the biozonation of Late Jurassic (Middle Tithonian) to Early Cretaceous (Valanginian) Tethyan pelagic carbonates to be established. The current standard calpionellid biozonation is based on studies by Remane (1964,) Le Hégarat and Remane (1978), Remane et al. (1986), and Pop (1994). A revised zonal and subzonal subdivision was proposed by Grün and Blau (1997; Fig. 10.32). Calpionellid stratigraphy is generally combined with biozonations based on calcareous nannoplankton (nannoconids, coccolithophorids) and radiolarians (Weidich and Schairer 1990; Ondrejickova et al. 1993).
remanei andrusovi
Pl. 77/5
bermudezi boneti
appears hyaline except in stratigraphically old taxa which exhibit microgranular, brownish opaque walls (Chitinoidella; Pl. 77/1). The fossils are studied in thin sections or peels. Taxonomic differentiation needs sections exhibiting the shape of the test (U- and C-shaped figures), the type of the collar bordering the oral apertura as well the shape of the aboral region. These criteria are shown by longitudinal sections (Pl. 77). Cross sections are circular and statistically more frequent (Fig. 10.30). Geometric variations due to different cutting planes and the possibility of tectonical deformation (Remane 1964) must be thoroughly considered. The length and width of the test are further criteria used in discriminating species.
Pl. 77/1
dobeni
Sediment production: Calpionellids were rockbuilding constituents of Late Jurassic and Early Cretaceous fine-grained basinal carbonates in a tropical belt extending from Mexico to eastern Indonesia. The fossils occur in high abundance typically in very fine-grained micritic limestones or marly limestones that are rich in other pelagic microfossils (coccolithophorids, nannoconids; Pl. 77/23), and radiolarians). Nannoconids are calcitic conical nannofossils particularly common in Tethyan pelagic and hemipelagic limestones ranging from the Late Jurassic (Tithonian) to the Late Cretaceous (Campanian). They usually co-occur with coccoliths, calpionellids and radiolarians, but almost purely nannoconid limestones are also known.
Fig. 10.32. Calpionellid biozonation after Grün and Blau (1997). The zones and subzones reflecting faunal successions and evolution events are used in worldwide correlations. Many of the species used to name the subzones are shown on Plate 77. Note that the generic assignment of these species must not correspond to the genera characterizing the major calpionellid zones.
are enclosed within an organic test (lorica) exhibiting a large oral opening (apertura) and a differentiated aboral pole. Morphology and thin-section criteria: Calpionellids have a bell-shaped, amphora-like or cylindrical, calcareous test (lorica) with axial symmetry, and a large oral opening (aperture) surrounded by a collar of variable shape and construction (Fig. 10.30, Fig. 10.31). The collar is usually formed by prolongation of the lateral walls, but can also correspond to independent elements. The aboral pole is rounded, pointed or characterized by a caudal appendix. In thin sections the wall
Other tintinnid-like thin-section fossils Morphologically similar microfossils characterized by bell-shaped figures and thought to be tintinnids have been described from the Paleozoic (Silurian: Hermes
Benthic Sessile Organisms
1966, Chennaux 1967; Devonian: Cuvillier and Sacal 1963; Viséan: Cuvillier and Barreyre 1964), elsewhere in the Mesozoic (Eichler 1965; Borza 1972), and from Tertiary limestones (Tappan and Loeblich 1968). Because of their considerably larger size and because of different environmental settings, only a few of these fossils are acknowledged to be true tintinnids. Basics: Calpionellids and other tintinnids Boorova, D., Lobitzer, H., Skupien, P., Vasicek, Z. (1999): Biostratigraphy and facies of Upper Jurassic-Lower Cretaceous pelagic sediments (Oberalm-, Schrambach- and Rossfeld Formation) in the Northern Calcareous Alps, south of Salzburg. – Abhandlungen der Geologischen Bundesanstalt Wien, 56, 273-318 Grün, B., Blau, J. (1996): Phylogenie, Systematik und Biostratigraphie der Calpionellidae Bonet, 1956: Neue Daten aus dem Rosso Ammonitico superiore und dem Biancone (Oberjura/Unterkreide): Tithon-Valangin von Ra Stua (Prov. Belluno, Italien). – Revue de Paléontologie, 15, 571-595 Grün, B., Blau, J. (1997): New aspects of calpionellid biochronology: proposal for revised calpionellid zonal and subzonal division. – Revue de Paléobiologie, 16, 197-214 Imbrie, J., Newell, N.D. (eds., 1964): Approaches in paleontology. – 432 pp., New York (Wiley) Rehakova, D., Michalik, J. (1993): Observations of ultrastructure of the Upper Jurassic and Lower Cretaceous calpionellid tests. – Geologica Carpathica, 44, 75-79 Remane, J. (1964): Untersuchungen zur Systematik und Stratigraphie der Calpionellen in den Jura-Kreide-Grenzschichten des Vocontischen Troges. – Palaeontographica A, 123, 1-127 Remane, J. (1985): Calpionellids. – In: Bolli, H.M., Saunders, J.B., Perch-Nielsen, K. (eds.): Plankton stratigraphy. – 555572, Cambridge (University Press) Tappan, H. (1993): Tintinnids. – In: Lipps, J.H. (ed.): Fossil prokaryotes and protists. – 285-303, Oxford (Blackwell) Further reading: K113
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10.2.3 Benthic Sessile Organisms The first two major groups treated in the discussion of thin-section fossils were cyanobacteria and calcareous algae, and foraminifera and other protozoans (Sect. 10.2.1 and Sect. 10.2.2). The third major group comprises sponges, corals and bryozoans as well as some hydrozoans. These groups have a common life-style: They are benthic sessile organisms, that contribute greatly to the production of marine carbonate rocks, particularly by forming reefs and accumulating skeletal grains in shallow-marine and deep-marine depositional settings. The information provided by sponges, corals and bryozoans greatly enhances our understanding of paleoenvironmental controls and facies belts.
10.2.3.1 Sponges The paleoenvironmental interpretation of microfacies data is often confronted with the need to differentiate and evaluate sponge fossils. Limestones with abundant sponges occur all over the Phanerozoic and in nearly all depositional settings from intertidal to deep marine. Sponges (Porifera) are benthic sessile metazoans with a differentiated inhalant and exhalant aquiferous system with external pores, in which flagellated cells (choanocytes) pump a water current through the body. The water is expelled by canals and the larger exhalant osculum. The water flowing through sponges provides food and oxygen, and also functions as waste removal. The firmness of the sponge body is provided by spongin fibres and a skeleton of silica or calcium carbonate.
Fig. 10.33. Introduction to sponges. A: Niphates digitalis from the reef at 24 m depth in Chan Kanaab, Cozumel, Mexico. Height of the pink sponge approximately 25 cm. B: Needles of Agelas cf. conifera: Slope in 51 m depths in Colombia, Cozumel, Mexico. C: Leucetta. Triactine and tetractine needles of a calcareous sponges. Cozumel, Mexico. The sponges shown in A and B are Demospongiae characterized by siliceous spicules. C shows a sponge belonging to the class Calcarea. The spicules consist of calcite. Courtesy of H. Lehnert (Oberottmarshausen).
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Sponge Ecology
Box 10.5. Describing sponges in microfacies thin sections involves a bewildering number of valid and outdated terms for sponge groups, some of which are systematic units, whereas others are combined by rather loose definitions. The definitions listed in the table are based mainly on the articles included in the comprehensive treatise edited by Hooper and Van Soest (2002). Archaeocyatha: Extinct sponges characterized by calcareous cups and a non-spiculate skeleton consisting of Mg-calcite. Now regarded as a group (class?) of Demospongiae with uncertain affinities. Calcarea: Porifera have a skeleton composed entirely of calcium carbonate. Characterized by free, rarely linked di-, tri-, tetra- and/or polyactine spicules to which a basal skeleton can be added. Comprise less than 5% of modern sponges. Calcareous sponges: The non-systematic term is used in a rather loose manner for ‘calcisponges’ exhibiting calcite spicules or a calcareous skeleton. Chaetetida: Formerly regarded as a group of tabulate corals. Now considered as a name for a constructional grade occurring in different sponge groups. In microfacies studies the name is commonly used in an informal way referring to fossils composed of densely packed, thin calcareous tubes with horizontal partitions. Coralline sponges: A polyphyletic group with affinities both to Demospongiae and Calcarea. Demospongiae: Porifera with siliceous spicules and/or fibrous spongin skeleton, or occasionally without a skeleton. Spicules are monaxon (either monactine or diactine) or tetraxon (tetractine), never triaxon. Comprise about 85% of all living sponges. Minor groups have developed a hypercalcified basal skeleton consisting of aragonite or Mg-calcite in addition to other elements, or exhibit a solid aragonite structure lacking free spicules. Hexactinellida: Sponges defined by siliceous triaxon hexactine spicules or derivations from that basic form. Comprise about 7% of the described recent sponges. Important members of deep-marine communities. Inozoa: Originally a name for coralline sponges without segmentation. Now applied to non-segmented calcareous sponges with spicules. Note that the term ‘inozoid’ is used in a non-systematic context referring to nonsegmented calcareous sponges, but also as the name for a systematic group.
The skeleton consists of discrete, articulated or fused spicules (Fig. 10.33) and/or a hypercalcified mineralized basal skeleton. Major groups: The subdivision of sponges into siliceous and calcareous sponges is anything but correct, but in the context of facies analyses still in use. In fact, there are three distinct classes (Hexactinellida, Demospongiae, and Calcarea, all appearing in the Late Proterozoic), and the extinct Cambrian class Archaeocyatha having suspected affinities wih Demospongiae. Several predominantly fossil groups including the Sphinctozoa, Stromatoporoidea and Chaetetida are in-
Ischyrospongia: A name introduced for a group which should include stromatoporoids, sclerosponges and pharetronids. The term should be abandoned but is still in use in papers dealing with microfacies of Paleozoic limestones. Lithistids: A polyphyletic group of demosponges characterized by articulated spicules (desmas) forming a rigid skeleton. Pharetronids: A term including Sphinctozoa and Inozoa. Also used for calcareous sponges exhibiting spicules. Sometimes considered as a subgroup of the Calcarea. The term is now only rarely used. Sclerosponges: A polyphyletic group of demosponges characterized by a calcareous basal skeleton with embedded siliceous spicules and different constructional grades resembling those of stromatoporoids and chaetetids. Siliceous sponges: Commonly used as an informal term for sponges with a skeleton composed of siliceous spicules. Include demosponges as well as hexactinellids. Sphinctozoa: Originally proposed as a term for chambered sponges with a rigid calcareous skeleton. Now defined by an external rigid skeleton and chambered organization. The name should not be used in a systematic context; it includes sponges that are polyphyletic in origin. Chambered construction has developed independently several times in different sponge groups and occurs in demosponges, Calcarea and hexactinellids. Most ‘sphinctozoans’ are attributed to the Demospongiae. Stromatoporoidea: Until the 70s of the twentieth century commonly regarded as hydrozoans. Giving credit to the affinities between the stromatoporoid skeletal structure and that of some recent sclerosponges, the group is now placed into the Porifera, and regarded as being related to Demospongiae, or as a separate class of sponges. Thalamida: Synonymous with Sphinctozoa.
terpreted as grades of construction and not phylogenetic clades, although the assignment of these groups and their lower taxa to the established classes is largely unresolved. Ecological constraints on modern sponges: Sponges are benthic, almost exclusively marine organisms. They occur at all latitudes, and from the intertidal to the deep sea, down to more than 5000 meters. A small group of freshwater sponges originated in the Cretaceous. The distribution of modern sponges in shallow and deepmarine environments is controlled by nutrients, substrate, sediment input, and water energy:
Sponge Ecology
• Nutrients: Sponges are suspension feeders. Nutrients are planktonic bacteria, organic particles and dissolved organic matter. Deeper-living sponges are exclusively heterotroph, whereas demosponges living in well-lighted environments may be autotroph due to the association with photosynthetic cyanobacteria. • Substrate: Most demosponges and hexactinellids need hard substrates (rock surface, shells), but hexactinellids with rhizomes also occur on soft substrates in the deep sea. • Sediment input: Deposition of too much sediment will effect sponge distribution negatively. • Water energy: Sponges occur predominantly within quiet-water environments; they are rare in high-energy environments (crustose growth forms). Sponges in reefs: In the past, modern sponges were not considered to be major reef-builders (e.g. bioconstructions formed in shallow waters of the Bahama
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Bank by demosponges: Wiedenmayer 1978), but rather more important as cryptic constituents of coral reefs. However, the recent exciting discovery of impressive reefs off British Columbia, formed by hexactinellids in waters 180 to 250 meters deep, has changed this view (Krautter et al. 2002). These reefs are particularly interesting because they may provide a living comparison to extensive reefs created by hexactinellids throughout the Jurassic. Sponges with a soft body occur in cryptic habitats of modern and ancient reefs. These sponges have a low preservation potential, unless the dispersal of the spicules after the decay of the dead sponge can be prevented by an early taphonomic calcification of the sponge tissue. As shown by the preservation of nonrigid sponges in Late Jurassic oyster patch reefs, this calcification may sometimes be induced by heterotrophic, symbiotic bacteria inhabiting the soft parts of the sponges (Delecat et al. 2001).
Fig. 10.34. Siliceous sponges were important in Late Jurassic reefs. The figure shows a marly sponge reef limestone with dish-shaped siliceous sponges carrying thrombolitic microbial crusts at the tops (see Chap. 9.1). This microfacies originated at the base as well as on top of small bioherms formed above sedimentary continuities in relatively deep environments (Brachert 1992). Note fracturing caused by differential compaction. These gray and yellowish sponge limestones are widely used as decorative building stones in central Europe. Late Jurassic (Kimmeridgian): Ludwag near Bayreuth, northern Frankenalb, Bavaria, Germany.
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Preservation of non-rigid soft sponges has been recorded from Silurian, Devonian and Jurassic mud mounds (Bourque and Gignac 1983; Flajs and Hüssner 1993; Reitner and Keupp 1991). Sponge spicules The fossil record of many sponge groups, particularly those of siliceous sponges, consists mainly of spicules which are studied in standard micropaleontological samples, but can sometimes also be recognized in thin sections of limestones (Keupp et al. 1989; see Pl. 78/2, 3). The composition and general architecture of spicules are used to classify sponges into classes. The form of the spicules and the manner in which they are combined to construct the skeletal structure are the basis for further systematic subdivisions. A complex terminology has been developed to describe the morphological variations of mega- and microscleres. Megascleres (length > 100 µm to several milli-
Sponge Spicules
meters) are the large principal frame-building elements. Microscleres (typically 10 to 100 µm long) commonly occur freely within the soft parts of the sponge. Major architectural spicule types differentiated according to the number of axes are monaxons (with a single axis), triaxones (basically consisting of six rays and three axes extending at right angles from a central point; restricted to hexactinellids), tetraxones (four rays radiating from a central point), and polyaxons (several rays approximately of equal length extending from a central point). Desmas are irregularly shaped spicules forming the skeletal network of lithistid sponges. Further morphological differentiations use the number of points or rays on the spicule; resulting terms are often characterized by the root ‘actine’. Sponge spicules may show a central canal (Pl. 112/1). Cross sections of calcitized spicules can be confused with radiolarians. Drawings of common spicule types which help to typify spicules in thin sections and a list of relevant papers can be found in Rigby (1985).
Plate 78 Siliceous Sponges ‘Siliceous sponges’ is a non-taxonomic term referring to sponges characterized by a skeleton consisting of siliceous (opaline) spicules. These sponges comprise taxa belonging to two polyphyletic systematic units with spicular skeletons, the Hexactinellida and the Demospongiae. Both groups include important reef builders (–> 1, 4, 6), include organisms living on soft and firm bottoms, and contribute to the formation of deep-shelf and basinal sediments (–> 2). Spicules can be differentiated according to the number of axes and rays and the shape of the scleres (–> 3). Most figures in the plate display hexactinellid sponges whose siliceous skeletons have been partly replaced by calcite. 1 Sponge reef limestone exhibiting hexactinellid sponges with a rigid skeleton. Hexactinellida (or ‘glass sponges’) are characterized by triaxon spicules consisting of six rays intersecting at right angles. Arrows point to the canal system. The sponge is strongly bored (B). Late Jurassic (Treuchtlingen limestone, Kimmeridgian): Bonhof, southern Bavaria, Germany. 2 Spiculite composed of predominantly monaxone megascleres. Causes for concentrations of spicules are in-place deposition of spicules derived from the disintegration of soft-bodied demosponges, or an accumulation of spicules of decaying soft sponges within organic mats. Some current transport is indicated by parallel alignment of the spicules. Early Jurassic: Roman mosaic stone found in Kraiburg, southern Bavaria, Germany. 3 Different types of sponge spicules. Megasclere spicules are classified according to the number of axis and to the number of points or rays on the spicule. C – Monaxon spicule with stalked rings (cricorhabd), M – Monaxon spicule pointed at one end (style), TE – Tetraxon spicule exhibiting one long and three short rays (triaenes), TR –triaxone spicule with three rays in one plane (triactines). The siliceous spicules have been dissolved; the molds are filled with granular calcite cement. The good preservation of the spicules points to slow dissolution and rapid cementation. Late Triassic: North of Isfahan, Iran. 4 Growth form. Tabular platy hexactinellid siliceous sponges covering fine-grained micritic matrix (coverstone). The arrow points to the canal system within the wall. The breccioid texture of the matrix is due to burrowing followed by dissolution. Same locality as –> 1. 5 Constructional grade. Some hexactinellid sponges exhibit internal segmentation similar to sphinctozoid demosponges (see Pl. 79). The sample is selectively silicified (white areas). Late Triassic: North of Isfahan, Iran. 6 Preservation of hexactinellid skeleton. A characteristic feature of sponge structures in thin sections is the fine network apparently consisting of discontinuous horizontal and vertical elements. Note the voids within the sponge skeleton caused by dissolution of the skeleton prior to the lithification of the lime mud between the spicules. Same locality as –> 1. –> 3, 5: Courtesy of B. Senowbari-Daryan (Erlangen)
Plate 78: Siliceous Sponges
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Spiculites (Pl. 78/2) are sediments consisting predominantly of accumulations of sponge spicules. Accumulations are widely distributed in modern cold-water environments. Spiculite mats consisting of an organic network of hexactinellid spicules and bryozoan fragments are known from Arctic seamounts which are covered by sponge-bryozoan bioconstructions (Henrich et al. 1992) but they also occur in other, e.g. near-shore, environments (Lane 1981). Fossil spiculites are common in deep-marine settings, both in basinal and slope position (Wiedenmayer 1980; Jach 2001; Pl. 112/1, Pl. 139/1). Calcareous and muddy sediments with a high percentage of spicules are usually interpreted as deep or cold-water deposits. However, today many siliceous demosponges live in warm shallow waters, suggesting that ancient spiculites could have originated in shallow-marine shelf and near-coast environments, too (Gammon and James 2001). Accumulations of spicules derived from siliceous sponges and spicule-moldic porosity may contribute to the formation of reservoir rocks in cherts (Saller 2001). The following text summarizes the criteria and significance of sponge groups commonly found in thin sections of limestones. Glossaries of the morphological terms used in describing sponges and extensive bibliographies can be found in Hooper and Van Soest (2002). Siliceous sponges These fossils are characterized by spicula composed of opal. They are polyphyletic and comprise two major
Siliceous Sponges
taxonomic units (Hexactinellida and Demospongiae with spicular skeletons). More than 85 % of all species of recent Porifera are ‘soft siliceous sponges’ with spicular supporting skeletons that are loosely embedded within the soft body. The skeleton disintegrates after the decay of the body. These sponges are already known in Liassic deposits and are important constituents of Jurassic and Cretaceous mud mounds. Hexactinellids are exclusively marine animals, today mostly known between 200 and 2000 meters, often living on soft substrates. These sponges are abundant and diverse at shallower depths in polar regions. Large mats of their spicules provide a hard substrate for other organisms. The second group of siliceous sponges is characterized by a rigid skeleton consisting of interconnected spicula (‘Lithistida’ and dictyonal hexactinellids). These sponges were important reef builders (Fig. 10.34). Siliceous sponges and ancient reefs: Associations between mound-building, benthic microbial communities and siliceous sponges are common in reef mounds (Early Cambrian, Early-Middle Ordovician, Late Silurian, Late Devonian, Late Mississippian, Late Permian, Late Jurassic (Brunton and Dixon 1994). Thromboliteor stromatolite-forming calcimicrobes acted as constructors. Demosponges and hexactinellid sponges acted as bafflers, binders and even as constructors of reefs commonly formed below wave base. Hexactinellids contributed to the formation of regionally restricted reefs in the Late Triassic (Wendt et al. 1989) and Early and Middle Jurassic, and wide ranging reefs in the Late Jurassic (see Leinfelder et al. 2002 for a review). The formation of these reefs was controlled
Fig. 10.35. The community in this reef cave is represented by spherical demospongid coralline sponges (Astrosclera, the variously sized balls) and associated with dendroid hydrozoans (Stylaster, right), a colonial sphinctozoid sponge (Vaceletia, top right corner) and various other encrusting sponges (center bottom). The bright orange-colored aragonitic basal skeleton of Astrosclera encloses siliceous spicules. Reef caves are different in light, nutrients and sediment input as compared with free reef surfaces. The specifically adapted fauna of these cryptic habitats yields some ‘living fossils’ e.g. coralline sponges whose ‘ultraconservative’ constructional style is reminiscent of Paleozoic and Mesozoic reef-building organisms. Great Barrier Reef (Osprey reef), western Australia. Courtesy of G. Wörheide (Göttingen).
Coralline Calcareous Sponges
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Fig. 10.36. Coralline sponges. The thin-section photograph shows a characteristic reef assemblage consisting of nonsegmented (inozoan, I) sponges, chambered sponges (sphinctozoans, S), and chaetetid coralline sponges (C). Note the competition for space, a common feature in reefs built by coralline sponges. Late Triassic (Zlambach beds, Rhaetian): Northern Calcareous Alps, Austria.
by an interplay of oscillating sea-level and fluctuating oxygen-minimum zones. The former favored the expansion of microbes, the latter in conjunction with successive highstands supported the growth of sponges in deeper water on foreslopes and drowned platforms (Reitner and Neuweiler 1995). Siliceous sponge diagenesis: In many limestones the skeleton of siliceous sponges is preserved as calcite pseudomorphs. As shown for Jurassic hexactinellids, sponges from different depositional environments reflecting different hydrodynamic regimes are preserved differently (Brachert 1991). Preservation is best in lowenergy basinal carbonates and bedded limestones formed between sponge mounds due to slow silica dissolution and rapid micrite cementation. As water turbulence (and water temperature) increases, dissolution accelerates with respect to micrite cementation. This leads to a complete calcitization of spicules in shallow and higher-energy areas, or the dissolution of the skel-
eton, which may then be recorded only as voids containing some relicts of sponge structures. Stagnant basins are suitable environments for the dissolution of spicules, which in turn may be responsible for the chert formation in limestones. The diagenesis of siliceous sponge spicules follows different paths (Gaillard 1983; Lang and Steiger 1984; Brachert et al. 1987). These include the infilling of the axial canals of the spicules with micritic sediment, iron and manganese oxides and cement, followed by the dissolution of spicules and infilling of the molds by HighMg calcite cement (Pl. 78/3) or by fine-grained sediment obscuring the former skeleton. Fossil siliceous sponges are frequently enclosed by or associated with micritic or peloidal ‘crusts’. Calcification of decaying tissue of siliceous sponges leads to the formation of various automicrite types that may mirror the shape of the sponge body (Warnke 1995). Major processes in sponge diagenesis and in the formation of sponge buildups are (a) early lithification and Text continued on p. 504
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Plate 79
Sphinctozoan and Inozoan Sponges
Calcareous Sphinctozoan and Inozoan Sponges
Many coralline sponges are characterized by chambered, segmented skeletons (–> 4, 8), others are non-segmented (–> 7). The former group has commonly been called Sphinctozoa, the latter Inozoa. These names are now no longer used in a systematic context, but are still applied to separate morphological groups of polyphyletic origin. Sphinctozoans are differentiated according to external and internal morphological features. External features include the shape (cylindrical, club-shaped, sheet-like or irregularly formed), the segmentation of the rigid skeleton, the pattern of chamber arrangement, and the inhalant and exhalant canal system of the outer wall. Internal features seen in thin sections are internal segmentation (–> 2, 4), chamber shapes (–> 2, 4, 6, 8) and internal fillings of the chambers (–> 1, 2, 3, 5) separating those parts of the skeleton which have been left behind by the soft parts during growth, internal walls separating the chambers (–> 4, 8) and the central exhalant system (siphon, spongocoel). The latter corresponds to single or multiple openings appearing as defined central canal or canal bundles (–> 3, 8).The schematic drawings on this page give an idea of the complex internal morphology of a sphinctozoan sponge (A), and demonstrate the possible relationship between the calcareous skeleton (black) and the soft parts (stippled) in a reconstruction (B). A - modified from Senowbari-Daryan (1990), B - after Reitner (1992). Most photographs display sphinctozoans (except –> 7, which shows a section of an inozoan sponge). Sphinctozoan sponges are abundant in Permian and Triassic reefs where they formed rigid organic frameworks (–> 1, 5), sometimes together with microbial crusts (–> 4). 1 Typical reef framework formed by encrusting sphinctozoans. Cross section of Jablonskya andrusovi (Jablonsky) (J) and Uvanella irregularis Ott (U) growing on another sphinctozoan sponge. Both sponges consist of microgranular High-Mg calcite skeleton (black in transmitted light). Note the vesicular filling skeleton of Jablonskya. Late Triassic (Carnian): Dolomites, Southern Alps, Italy. 2 Morphological criteria of sphinctozoans. Oblique longitudinal section of Solenolmia manon (Münster) characterized by segments with a coarsely reticulate filling skeleton. Late Triassic (Carnian): Molignon, Dolomites, Southern Alps, Italy. 3 Sphinctozoan morphology: Vesicocaulis reticuliformis Jablonsky with central canal bundles (arrow) connecting the catenulate segments and the reticular filling skeleton. Longitudinal section. Late Triassic (Carnian): Huda Juzna, Slovenia. See schematic drawing on this page. 4 Sphinctozoan morphology: Longitudinal section of Colospongia dubia (Laube). The moniliform stem consists of stalked barrel-shaped segments with differentiated porate walls (PW) and vesiculae (V) within the chambers. Note the different appearance of the originally aragonitic skeleton as compared with originally Mg-calcitic skeletons (–> 1). Late Triassic (Carnian) patch reef: Huda Juzna, Slovenia. 5 Sphinctozoan morphology: Several densely spaced specimens of Solenolmia manon (Münster). The segments contain a coarsely reticular filling skeleton. Internal walls of the segments are perforated. Solenolmia is a common sponge in Middle Triassic and Carnian reefs of the Alpine-Mediterranean region. Late Triassic (Carnian): Huda Juzna, Slovenia. 6 Sphinctozoan morphology: Alpinothalamia slovenica (Senowbari-Daryan). Numerous glomerate cyst-like chambers are arranged around the axial spongocoel. The wall consists of High-Mg calcite. Late Triassic (Pantokrator limestone, Carnian): Hydra Island, Greece. 7 Inozoan sponge morphology: Oblique section of Sestrostomella robusta Zittel. Similar to sphinctozoan sponges, inozoans are characterized by the absence of segmentation. The sponge body is transected by axial canal bundles of the exhalant system. Sestrostomella is known from Triassic and Late Jurassic-Early Cretaceous limestones. Late Triassic (Carnian): Dolomites, Southern Alps, Italy. 8 Sphinctozoan and inozoan sponges: Longitudinal section of the sphinctozoan Discosiphonella lercarensis (SenowbariDaryan and Di Stefano). The central cavity is surrounded by ringlike rows of hollow chambers. The internal walls are perforated. The sphinctozoid sponge was competing for space with the inozoid sponge Maeandrostia (arrows). These sponges occur in shelf habitats, not in reefs. Early Permian: Lercara, Sosio Valley, Western Sicily, Italy. –> 2–8 and Figs. on this page: Courtesy of B. Senowbari-Daryan (Erlangen)
Plate 79: Sphinctozoan and Inozoan Sponges
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Paleozoic Stromatoporoids
Plate 80 Paleozoic Stromatoporoids Stromatoporoids are common in thin sections of Ordovician to Late Devonian limestones as well as Jurassic and Cretaceous shelf and reef carbonates (see Pl. 81/3, 5). For a long time regarded as hydrozoans, stromatoporoids are now recognized as sponges. Major points in favor of including the Stromatoporata within the sponges are the similarity of the astrorhizal canal system and the gross skeleton patterns to modern sclerosponges. The basal calcareous skeleton of the Stromatoporata consists of horizontally continuous elements (laminae, –> 8; cysts, –> 5) and vertical elements (pillars, –> 3, 7, 8; columns, –> 4). Most of the Paleozoic stromatoporoids appear black in transmitted light, some appear hyaline (–> 1) pointing to differences in the original skeletal mineralogy (High-Mg calcite and calcite: Rush and Chavetz 1991; Yoo and Lee 1993; aragonite: Smosna 1984; Stearn and Mah 1987; Stearn 1989; Tobin and Walker 1998). Much of the hydrocarbon production from Silurian and Devonian carbonates is from stromatoporoid reef and forereef limestones. Growth forms, taphofacies as well as distributional patterns of genera in reef complexes are useful criteria in recognizing environmental conditions (water energy, substrate) and ecologic zonations. The plate displays Devonian stromatoporoids with laminar-encrusting (–> 3), massive-nodular (–> 7, 8) or cylindrical-dendroid (–> 1, 2) growth forms. 1 Stachyodes. Order Syringostromatida. Oblique section. Note the dendroid growth form and the dense packing of skeletal elements. Genus range: Eifelian to Frasnian. The genus is common on reef flats and near-reef platform environments. Late Devonian (Frasnian): Canning Basin, western Australia. 2 Amphipora. Transversal (TS) and longitudinal sections (LS). These small, cylindrical, branching organisms lived in calm shallow-marine waters of lagoonal and back-reef environments and were frequently swept into other environments by storms. Black, bituminous Amphipora limestones characterize a widely distributed organic-rich facies type of Devonian reef and shelf carbonates. Diagnostic criStructure and microstructure of stromatoporoids. Cross secteria: Dendroid skeleton, commonly with axial canal, surtion of Stachyodes (–> 1). The cylindrical branch exhibits rounded by a network formed by pillars and connecting an axial canal. The structure is characterized by a network elements. Frequently peripheral vesicles bounded by a calcarof radiating pillars and voids. The latter correspond to ireous membran. Generic range: Late Silurian to Late Devonregular radial canals opening at the margin. Dark microian. Late Devonian (Frasnian): Canning Basin, Australia. laminae parallel to growth surfaces form concentric rings. 3 Intergrowth of stromatoporoids and corals: A rugose coral The originally microreticulate microstructures of the thick (RC) is encrusted by a consortium consisting of stromatoskeletal elements appear striated, intensified by recrystalliporoids (S, Clathrocoilona), alveolitid tabulate corals (TC) zation. Late Devonian (Frasnian): Canning Basin, western and porostromate cyanobacteria (PC, Rothpletzella, see Pl. Australia. 53/8). Overgrowth of rugose corals by stromatoporoids is a common feature in Devonian and Silurian reef communities. Middle Devonian (Givetian): Letmathe, Sauerland, Germany. 4 Stromatoporoid microstructure: Longitudinal section of Parallelopora characterized by thick superposed but discontinuous vertical skeletal elements (coenostele) with a cellular microstructure. No distinct laminae but upward arched dissepiments. Generic range: Wenlockian to Middle Devonian. Eifelian: Meseta, Morocco. 5 Labechiid stromatoporoids: Longitudinal section. The last Devonian stromatoporoids have a similar construction to stromatoporoids appearing during the initial development of the group in the Ordovician. Predominant skeletal elements are cysts. Vertical elements are reduced to spines and short pillars. Some labechiids had aragonitic skeletons. Cyst-dominated skeletons are common in Ordovician and Silurian carbonate banks and reefs. Generic range of Pseudolabechia: Late Devonian (Famennian): Dubic-Debnik, south of Cracow, southern Poland. 6 Irregular meshwork: Oblique section of Hammatostroma albertense Stearn. Criteria: Persistent laminae; incomplete pillars knotted into irregular tangled structures within the interlaminar space. Generic range: Givetian to Frasnian. Late Devonian (Frasnian): Subsurface, Red Water Reef Complex, Alberta, Canada. 7 Regular meshwork: Vertical (VS) and tangential sections (TS) of Actinostroma. Circular figures correspond to sections through dome-shaped parts (mamelons) of the fossil. Note geopetal fabric in bivalve boring (arrow). Generic range: Lochkovian to Famennian. Late Devonian (Famennian): Same locality as –>5. 8 Growth pattern: Longitudinal section of Actinostroma papillosum (Bargatzky). Variations in laminae spacing indicate growth disturbances (GD). Growth units (latilaminae) offer valuable proxies for paleoclimatic variations. Reef margin. Late Devonian (Frasnian): Canning Basin, western Australia.
Plate 80: Paleozoic Stromatoporoids
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Chaetetids
Plate 81 Calcareous Chaetetid Sponges and Other Still Disputed Taxa The plate shows a selection of Triassic and Jurassic fossils characterized by uniform or differentiated calcareous tubes and sometimes summarized under the name ‘tabulozoans’. This term should be forgotten. With the discovery of living, very similar animals (sclerosponges and related forms), it was learned that most of these fossils are sponges although the former systematic assignment of some taxa to hydrozoans (–> 3, 5) has not yet been completely rejected. Figs. (–>1-4 and 6-8) display some common chaetetid coralline sponges. The skeleton of chaetetids is formed in two steps. Skeleton formation starts with the precipitation of a spicular primary skeleton, followed by the formation of a secondary skeleton consisting of calcareous walls and horizontal partitions (see schematic drawings on this page). Important diagnostic criteria used in taxonomic differentiations are the morphology and mode of the formation of the tubes, the primary mineralogy and microstructure of the walls, the type of the horizontal partitions (tabulae; –> 8) and the mode of the formation of new tubes (–> 7, 8, and schematic drawings on this page). Chaetetids range from Ordovician to Miocene and peaked in the early Late Carboniferous, where they formed shallow-marine biostromal reefs. These reefs were constructed more or less only by chaetetids, in contrast to Late Jurassic reefs where chaetetids played only a small part in the association of stromatoporoid sponges and corals. Chaetetids in Early Cretaceous rudist reefs were usually dwellers.
1 Different types of coralline calcareous sponges. The reef limestone contains an assemblage of spongiomorphids (S), sphinctozoan sponges (SS), inozoan sponges (IS) and skeletons composed of calcareous tubes (formerly summarized as tabulozoans, T). The latter exhibit the chaetetid constructional type known from sclerosponges. Note the different preservation of spongiomorphids and inozoans as compared with sphinctozoans and chaetetids, pointing to differences in the original skeleton mineralogy. Many Mesozoic chaetetids had an aragonitic skeleton (Wendt 1990; Mastrandrea and Russo 1995). Spongiomorphids have been regarded as corals or hydrozoans but an assignment to sponges is now under discussion. Late Triassic (Dachstein reef limestone, Norian): Hoher Göll, Bavarian Alps, Germany. 2 Preservation of chaetetid skeletons: The infill of micrite into the peripheral part of the skeleton simulates the existence of a morphological differentiation. Blastochaetetes sp. Same locality as –> 1. 3 Astrostylopsis (A), a sponge with Mg-calcite skeleton formerly regarded as hydrozoan, encrusted by a chaetetid (C; Bauneia sp.) with preserved aragonitic skeleton. Note the difference in color in transmitted light between Mg-calcite and aragonite. Late Jurassic: Jebel Zhaguan, Tunisia. 4 Oblique longitudinal section of a chaetetid. Same locality as –> 1. 5 Ellipsactinia (E), sometimes regarded as a hydrozoan, and Astrostylopsis (A) assigned to coralline sponges. The constructional pattern of Ellipsactinia corresponds to the stromatoporoid type. Ellipsactiniids contributed significantly to the formation of Late Jurassic and Early Cretaceous platform-margin reefs. Late Jurassic: Central Apennines, Italy. 6 Ptychochaetetes globosus Koechlin exhibiting distinct growth zonation. Late Jurassic (Tithonian): Krahstein near Mitterndorf, Styria, Austria. 7 Ptychochaetetes. Longitudinal section. The mode of the formation of new tubes (here basal budding) is used as a generic criterion. Late Jurassic (Kimmeridgian): Vöksen/Deister, Lower Saxony, Germany. 8 Neuropora, a chaetetid sclerosponge formerly attributed to bryozoans. Longitudinal section exhibiting densely spaced cross partitions. New tubes are formed by intramural budding. Same locality as –> 6. 9 Spongiomorpha ramosa Frech showing a fine reticulate pattern consisting of crenulated vertical elements. The assignment of spongiomorphids to hydrozoans is now questioned and sponge affinities preferred. Spongiomorphids were common in deeper subtidal parts of Triassic to Early Cretaceous reefs and platforms. Note that the term ‘spongiomorphid’ is also loosely used for sponge-like fossils! Late Triassic (Rhaetian): Steinplatte, Tyrol, Austria. –> Figure on this page: Cremer (1993)
Plate 81: Chaetetid Sponges
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replacement of the former sponge tissue by in situ precipitated carbonate cement (automicrite, Sect. 4.1.2) and (b) early or late dissolution and replacement of the siliceous skeleton by calcite. Dissolution depends on pressure, temperature and the type of modification. Skeleton opal is relatively stable under normal marine conditions (pH about 8). Dissolution, therefore, requires changes in pH values and an increase in carbonate alkalinities. An increase in pH values may be caused by (1) bacterial degradation of organic matter including ammonification, or desulfurication triggered by sulfur bacteria, or (2) removal of CO2 due to changes in physical parameters. The importance of these processes in pH changes is still a matter of discussion (Fritz 1958; Reitner et al. 1995; Rehfeld 1996). Distribution of siliceous sponges: During the Paleozoic siliceous sponges flourished in quiet-water environments, contributing to the formation of mud mounds. In the Mesozoic siliceous sponges are abundant in the Jurassic and Cretaceous. Liassic sponges are known
Coralline Calcareous Sponges
from soft bottom deeper-water environments and are abundant in Late Jurassic ramp and basinal settings where they formed extensive and widely distributed reefs. During the Late Cretaceous a major turnover in demosponges took place, as demonstrated by the highly diverse sponge faunas of the Chalk facies in northwestern Europe. The adaption of siliceous sponges to deepwater environment was continued during the Tertiary. Coralline calcareous sponges Morphology and classification: Coralline sponges are defined as sponges with a secondary rigid calcareous skeleton in addition to a spicular skeleton. Rigid calcareous skeletons occur in different sponge groups including sphinctozoan and inozoan sponges, stromatoporoids and chaetetids as well as archaeocyaths. The great debate on the assignment of these groups to the major sponge classes was strongly triggered by studying modern sclerosponges that exhibit morphological criteria occurring in different fossil groups.
Plate 82 Archaeocyatha Archaeocyaths represent an extinct group of benthic, sessile organisms that lived in Cambrian equatorial openshelf and reef environments. The cup-shaped, vase-like or cylindrical calcareous skeletons are characterized by a large central cavity surrounded by a concentric porous double wall (–>1, 2) or by a porous single wall (–> 3). The space between the inner and the outer wall (intervallum) is subdivided by radial-longitudinal and/or radialtransverse partitions (septa, flat tabulae, curved cysts). The arrangement of these partitions, their pore systems and the constructional pattern are of taxonomic importance (–> 1, 2). Archaeocyaths are often associated with calcimicrobes (–> 4, 5) and form the substrate for microbes and submarine carbonate cements, thus contributing to the formation of Early Cambrian reef mounds and reefs. 1 Double-walled archaeocyaths: Various sections of double-walled archaeocyaths. Cross- and oblique sections (CS, OS) show the central cavity bounded by porous inner and outer walls. The intervallum is subdivided by septa. The septa are finely porous, as shown in a tangential section (TS). These structures exhibit the characteristic archaeocyathid constructional type of the skeleton. Other constructional types used in taxonomic differentiations are indicated by black and white arrows (vesicular and dictyonid network respectively). The lower part of the figure below the wavy line shows a parallel section to the upper part taken within a distance of about 2 mm. Note the apparently different filling of the central cavity. Early Cambrian: Western Antarctica. 2 Archaeocyath morphology: Cross section (CS) of double-walled archaeocyaths and oblique section (OS) of archaeocyaths corresponding to another constructional type characterized by septa arranged in a network form. The white arrow points to tubular cyanobacteria, the black arrows to juvenile archaeocyathids. SMF 6. Early Cambrian: Sardinia, Italy. 3 Single-walled archaeocyath: Oblique cross section. The tubelike structures have been described as Proaulopora (P) and were referred to green and red algae and cyanobacteria as well. This genus indicates an Early Cambrian age. Early Cambrian: Same locality as –> 2. 4 Associated calcimicrobes: Many archaeocyaths occur in association with calcimicrobes that use the archaeocyaths as a substrate. The figure shows Epiphyton. The disintegration of the erect-growing micritic branches of Epiphyton and other calcimicrobes provides coarse silt to very fine sand-sized fragments that lack characteristic microstructures and therefore correspond to peloids. Cyanobacterial peloids contributed greatly to the production of carbonates deposited adjacent to Early Paleozoic reef mounds (Coniglio and James 1985). Same locality as –> 2. 5 Associated calcimicrobes: Widely distributed Cambrian calcimicrobes are Renalcis (R) and Epiphyton (E). Renalcis is characterized by clusters of thick-walled botryoidal chambers, with the larger ones enveloping smaller ones. Epiphyton exhibits micritic, narrow dendritic filaments. Both taxa are known from Early Cambrian to Late Devonian. Same locality as –> 2. –> 2–5: Courtesy of M. Boni (Napoli) and T. Bechstädt (Heidelberg)
Plate 82: Archaeocyatha
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Diving in tropical coral reefs resulted in the discovery of puzzling sponges restricted to reef caves and deep forereef areas with low light and strict oligotrophic conditions in the West Atlantic and Indo-Pacific region (Fig. 10.35). These sponge, called sclerosponges by Hartman and Goreau (1972), are characterized by a basal calcareous skeleton by which the animals are attached to the substrate (Fig. 10.35). The living tissue is restricted to a thin veneer upon or within the uppermost few millimeters of the skeleton and contains, usually but not always, tiny siliceous spicules. The taxa were first regarded as a separate class of the Demospongiae, but subsequent to the recognition of the polyphyletic character of these sponges, they are now attributed to different families of coralline sponges. The skeleton of these sponges consists of aragonite or Mg-calcite. It is organized into tubes and calyces (ceratoporellid type), tabulated tubes (chaetetid type), a more or less regular network of the primary skeleton with horizontal and vertical elements (stromatoporoid type), or a chambered skeleton (sphinctozoid or thalamid type). Comparing the morphological criteria of sclerosponges and other ‘living fossils’ with the criteria of fossil groups of until then unassigned systematic position, such as the stromatoporoids and chaetetids, has now led to the assignment of these groups as well as sphinctozoans and inozoans to the demosponges. Some authors oppose this systematic treatment and regard these groups primarily as different grades of the morphological organization of skeletons occurring independently in different sponge classes (Reitner 1989; Wood 1991; Reitner and Wörheide 2002). Overview of major groups of coralline sponges: The following text and the plates summarize the criteria, distribution and significance of sphinctozoan and inozoan sponges (Pl. 79), stromatoporoids (Pl. 80) and chaetetids (Pl. 81) as well as archaeocyaths (Pl. 82). Sphinctozoans and inozoans Diagnostic morphological criteria of chambered sponges (sphinctozoans) and non-chambered calcareous sponges (inozoans) are shown in Plate 79. Sphinctozoans appear early in the Paleozoic, diversified and dominated during the Permian and Triassic, and became nearly extinct after the Cretaceous (Senowbari-Daryan 1990; review and extensive bibliography in SenowbariDaryan and Garcia-Bellido 2002). Sphinctozoans and inozoans as well as other coralline sponges contributed significantly to the formation of Late Paleozoic (e.g. Permian Reef Complex in Texas, Pl. 145/1) and Middle and Late Triassic reefs (e.g. Middle Triassic Wetterstein
Stromatoporoids
reefs and Late Triassic Dachstein reefs in the Northern Calcareous Alps; Fig. 10.36, Pl. 116/2). The first inozoans are known from the Carboniferous. Permian inozoan sponges were common both in non-reef and reef environments (Rigby et al. 1989, Fan et al. 1991, Rigby and Senowbari-Daryan 1996). The group reached a maximum in the Late Triassic (Senowbari-Daryan et al. 1997) and again in the Middle and Late Jurassic (Müller 1984). The last inozoans with a rigid non-spicular skeleton were described from the latest Cretaceous. Stromatoporoids Stromatoporoids are extinct fossils known from the Early Ordovician to the Late Devonian and again in the Mesozoic. The relations between Paleozoic and Mesozoic taxa are still open to discussion. Both Paleozoic and Mesozoic stromatoporoids have been placed into the Hydrozoa but are now recognized as Porifera. The fossils are represented by calcified skeletons of domical, bulbous, branching or columnar form, and ranging from a few centimeters to meters in diameter. Internally the skeleton exhibits elements more or less parallel and perpendicular to the surface. Distinct aquiferous canal systems (astrorhizae) occur at the surface and internally. The canals converge towards a center and are radially arranged. No traces of spicules have been found in Early and Mid-Paleozoic stromatoporoids. In this feature Paleozoic stromatoporoids differ from Mesozoic stromatoporoids, some of which contain spicules. These taxa, including many former ‘hydrozoans’ (see Pl. 81), are placed into the Demospongiae. Many authors regard the Paleozoic stromatoporoids as a separate class of the Porifera with close relationships to other poriferans secreting calcareous skeletons. Morphology and classification: A glossary of the morphological terms applied to the Paleozoic Stromatoporoidea and a comprehensive overview of the current state of classification, including a critical list of the genera was provided by Stearn et al. (1999). Criteria used in the taxonomic discrimination include the type of skeletal elements, their microstructure and features of the exhalant water system represented by astrorhizal canals and central tubes. Some microstructural types are identical with those of other coralline sponges and are therefore regarded as primary skeletal structures. Chaetetid coralline sponges and ‘tabulozoans’ Thin sections of Late Paleozoic and Mesozoic reef limestones may exhibit sections of fossils composed of calcareous tubes, sometimes with tabula-like horizontal partitions. In the context of microfacies studies
Archaeocyaths
these fossils have often been called ‘tabulozoans’, using a non-systematic term proposed by Kühn (1942). Plate 81 displays some of these taxa that were formerly assigned to various fossil groups, including hydrozoans, tabulate corals and sponges. Most taxa, particularly the chaetetids, are now recognized as sponges related to calcified Demospongiae. The chaetetids compose a small group of organisms that was presumed to be tabulate corals, but are now attributed to sponges as a result of the discovery of a recent representative. The rigid skeleton corresponds to a cluster of closely packed nonporous calcareous tubes with horizontal floors (tabulae). Spicules have been described from Mesozoic chaetetids. Fossil chaetetids appear in the Ordovician. They contributed to the development of Paleozoic (particularly Late Carboniferous) and Mesozoic reefs, both overgrowing and serving as substrate for other reefbuilding organisms. Chaetetid coralline sponges were important reefbuilders both in the Paleozoic and Mesozoic. In Middle Carboniferous platform and platform-margin settings they formed small mounds and broad low banks where they were often associated with beresellid and ungdarellid algae (see Pl. 56/6, 7 and Pl. 60/2, 5). In the Triassic, Late Jurassic and Early Cretaceous chaetetids lived in open-marine platform environments and in reefs. Archaeocyaths This is an extinct group of Cambrian fossils now regarded as Demospongiae with uncertain affinities (Pl. 82/1-3). The group appears near the base of the Cambrian (Early Tommotian) and had a worldwide distribution in the Early Cambrian (Atdabanian and Botomian). Some relicts persisted in Middle and Late Cambrian. Compositional differences of faunas from Siberia, Europe, Northern America, North Africa and Australia indicate distinct provincialism. Morphology and classification: The skeleton consisting of microgranular Mg-calcite corresponds to inverted cones or cups exhibiting a distinct central cavity surrounded by porous walls formed by one or two elements . The ‘intervallum’ between the outer wall and the highly variable inner wall contains a variety of radially or irregularly arranged partitions (blade-like septa; irregular curved plates or subradial pseudosepta). These elements can be connected by bars. The lower end (tip) of the cone was buried in soft sediment or, more commonly, rooted or encrusted onto hard substrates. The diameter of the cups changes with the ontogenetic stages and varies between a few millimeter
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and about 600 mm. The average diameter of adult forms is 5 to 25 mm. The height can be more than 300 mm. The classification is based on ontogenic development stages, intervallum structures and pore systems. Generic determination needs oriented transverse, longitudinal and tangential thin sections. Transverse sections can be confused with corals and dasyclad green algae, but most corals lack a distinct double wall. Transverse sections of dasyclads are smaller, and the simple calcareous walls of the algae are penetrated by small pores. Ecology and facies: Archaeocyaths were exclusively marine filter-feeders living on fairly energetic ramps and platforms. Solitary and colonial archaeocyaths were important constituents of the first metazoan reef communities but also occur in subtidal bottom-level associations. Biohermal and biostromal reefs are believed to have been formed in shallow water depths (about 20–50 m). Depth estimates are based on the abundant association of archaeocyaths and various calcimicrobes (Renalcis, Epiphyton; sometimes quantitatively more important than archaeocyaths), current-produced sedimentary structures, and the abundant fragmentation and resedimentation of archaeocyathan skeletons. Associated biota include trilobites, hyalithids, various shells and fossils with supposed cnidarian affinities. Many limestones with archaeocyaths are boundstones or bioclastic and peloidal floatstones and wackestones.
Basics: Sponges (publications which assist in taxonomic differentiation are marked by an asterisk) Overviews Broadhead, T.W. (ed., 1983): Sponges and spongiomorphs. Notes for a short course. – 220 pp., Knoxville (University of Tennessee) Hartman, W.D., Wendt, J., Wiedenmayer, F. (1980): Living and fossil sponges. – Sedimenta, 8, 274 pp. Hooper, J.N.A., Van Soest, R.W.M. (eds., 2002): Systema Porifera. A guide to the classifications of sponges. 2 volumes. – 1708 pp., New York (Kluwer Academic/Plenum) Reitner, J., Keupp, H. (eds., 1991): Fossil and recent sponges. – 595 pp., Berlin (Springer) Siliceous sponges Brachert, T. (1991): Environmental control on fossilization and siliceous sponge assemblages: a proposal. – In: Reitner, J., Keupp, H. (eds.): Fossil and recent sponges. – 543-553, Berlin (Springer) *Krautter, M. (1997): Aspekte zur Paläökologie postpaläozoischer Kieselschwämme. – Profil, 11, 199-324 Krautter, M. (2002): Fossil Hexactinellida: an overview. – In: Hooper, J.N.A., Van Soest, R.W.M. (eds.): Systema Porifera. A guide to the classifications of sponges, 2. – 1211-1222, New York (Kluwer Academic/Plenum)
Hydrozoans
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*Mostler, H. (1976): Poriferenspiculae aus der alpinen Trias. – Geologisch-Paläontologische Mitteilungen Innsbruck, 6, 1-42 Mostler, H. (1990): Hexactinellide Poriferen aus pelagischen Kieselkalken (Unterer Lias, Nördliche Kalkalpen). – Geologisch-Paläontologische Mitteilungen Innsbruck, 17, 143178 Pisera, A. (2002): Fossil ‘lithistids’. – In: Hooper, J.N.A., Van Soest, R.W.M. (eds.): Systema Porifera. A guide to the classifications of sponges, 1. – 388-402, New York (Kluwer Academic/Plenum) *Rigby, J.K. (1995): Sponges as microfossils. – In: Blome, C.D. (ed.): Siliceous microfossils. – Short Course in Paleontology, 8, 1-17 (The Paleontological Society) Warnke, K. (1995): Calcification processes of siliceous sponges in Viséan limestones (Counties Sligo and Leitrim, Northwestern Ireland). – Facies, 33, 215-228 Coralline sponges, stromatoporoids and chaetetids Hartman, W.D., Goreau, T.F. (1970): Jamaican coralline sponges: their morphology, ecology and fossil relatives. – Symposium of the Zoological Society London, 25, 205-243 Reitner, J. (1992): ‘Coralline Spongien’. Der Versuch einer phylogenetisch-taxonomischen Analyse. – Berliner Geowissenschaftliche Abhandlungen, E, 1, 1-352 Reitner, J., Wörheide, G. (2002): Non-lithistid fossil Demospongiae. Origin of their paleobiodiversity and highlights in the history of preservation. – In: Hooper, J.N.A., Van Soest, R.W.M. (eds.): Systema Porifera. A guide to the classifications of sponges, 1. – 52-68, New York (Kluwer Academic/Plenum) *Senowbari-Daryan, B. (1990): Die systematische Stellung der thalamiden Schwämme und ihre Bedeutung in der Erdgeschichte. – Münchner Geowissenschaftliche Abhandlungen, A, 21, 1-326 Senowbari-Daryan, B., Garcia-Bellido, D.C. (2002): Fossil ‘Sphinctozoa’: chambered sponges (polyphyletic). – In: Hooper, J.N.A., Van Soest, R.W.M. (eds.): Systema Porifera. A guide to the classifications of sponges, 2. – 15111533, New York (Kluwer Academic/Plenum ) *Stearn, C.W., Webby, B.D., Nestor, H., Stock, C.W. (1999): Revisited classification and terminology of Palaeozoic stromatoporoids. – Acta Palaeontologica Polonica, 44, 1-70 *Turnsek, D., Buser, S., Ogorelec, B. (1981): An Upper Jurassic reef complex from Slovenia. – In: Toomey, D.F. (ed.): European fossil reef models. – Society of Economic Paleontology and Mineralogy, Special Publications, 30, 361-369 *Turnsek, D., Masse, J.P. (1973): The lower Cretaceous Hydrozoa and Chaetetidae from Provence (south-eastern France). – Slovenska Akademija Znanosti in Umetnosti, Razprave, 16, 217-244 Wendt, J. (1984): Skeletal and spicular mineralogy microstructure and diagenesis of coralline calcareous sponges. – Palaeontographica Americana, 54, 326-336 Wörheide, G. (1998): The reef cave dwelling ultraconservative coralline demosponge Astrosclera willeyana Lister 1900 from the Indo-Pacific. – Facies, 38, 1-88 Wood, R.A. (1987): Biology and revised system of some late Mesozoic stromatoporoids. – Special Papers in Palaeontology, 37, 1-89 Wood, R.A. (1991): Non-spicular mineralization in calci-
fied demosponges. – In: Reitner, J., Keupp, H. (eds.): Fossil and recent sponges. – 322-340, Berlin (Springer) Archaeocyaths *Debrenne, F., Rozanov, A., Zhuravlev, A. (1990): Regular archaeocyaths. Morphology, systematic, biostratigraphy, paleogeography, biological affinities. – Cahiers de Paléontologie, 218 pp. *Debrenne, F., Zhuravlev, A. (1992): Irregular archaeocyaths. Morphology, ontogeny, systematics, biostratigraphy, palaeoecology. – Cahiers de Paléontologie, 212 pp. *Debrenne, F., Zhuravlev, A.Yu., Kruse, P.D. (2002): Class Archaeocyatha Bornemann. – In: Hooper, J.N.A., Van Soest, R.W.M. (eds.): Systema Porifera. A guide to the classifications of sponges, 2. – 388-402, New York (Kluwer Academic/Plenum) Rowland, S.M., Gangloff, R.A. (1988): Structure and palaeoecology of Lower Cambrian reefs. – Palaios, 3, 111-135 Zhuravlev, A., Wood, R. (1995): Lower Cambrian reefal cryptic communities. – Palaeontology, 38, 443-470 Further reading: Sponges K114, siliceous sponges K115, sphinctozoan and inozoan sponges K116, stromatoporoids K117, chaetetid sponges K118, archaeocyaths K214
10.2.3.2 Hydrozoans If Paleozoic stromatoporoids and most Mesozoic fossils previously described as ‘hydrozoans’ or ‘stromatoporoids’ now are to be assigned to the sponges, the fossil record of true calcareous hydrozoans found in thin sections of limestones is meager and restricted to Cretaceous and Tertiary Milleporidae, Stylasteridae (Pl. 147/5), and some Hydractiniidae. Still one hydrozoan candidate is Heterastridium, a puzzling fossil known only from the Late Triassic and occurring abundantly in open-marine limestones (Cuif 1971; Fig. 10.37).
Fig. 10.37. Heterastridium Reuss. The globular or nodular fossil ranges in size between a few centimeters to tens of centimeters. It consists of radially arranged tubes of different morphology, some of which are reflected by pustules on the surface. Mass occurrences of Heterastridium have been described from Norian basinal limestones all over the Tethyan ocean. The genus was considered to be a hydrozoan - but who knows? Late Triassic (Nayband Formation, Norian): Kuh-e-Nayband, central Iran. Courtesy of B. SenowbariDaryan (Erlangen).
Major Coral Groups
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10.2.3.3 Corals
members of the exclusively marine, usually sedentary Anthozoa. The number and symmetry of vertically arranged infoldings of the gut wall in the body cavity of the polyps and the type and composition of the skeleton serve to divide the class into various groups, such as the soft corals, sea fans (Octocorallia), sea anemones, and the stony or true corals. Octocorals occur from tropical to polar regions and from tidal zones to deep sea. The skeleton of octocorals consists of small calcitic sclerites (0.10 to 1 mm) that are common in recent sediments but very rare in the fossil record. True corals, commonly collected in the subclass Zoantharia,
Corals are common constituents of limestones. Paleozoic rugose corals and tabulate corals are widely distributed in shelf and ramp environments and contribute to the formation of reefs. Mesozoic and Cenozoic scleractinian corals are important reef builders in tropical and subtropical warm-water zones but also occur in temperate and cold-water environments. Major groups: Corals are Cnidaria comprising the hydrozoans, scyphozoans and anthozoans. Corals are
Box 10.6. Major coral groups. Technical terms used are: axial structure - structural element occurring in the center of corallites; corallum - the skeleton of a coral colony; corallite - the skeleton of an individual coral polyp; coenosteum colonial skeletal deposits formed between corallites; dissepiments - blister-like calcareous elements between septa; edge zone - a fold in the body wall extending over the edge of the corallum; epitheca - outer skeletal sheet laid down by the edge zone; microarchitecture - the spatial arrangement of microstructural elements; microstructure - basic elements of septa and walls; protosepta - initial septa formed at first; septa - radial and vertical element extending from the wall toward the center of corallites; tabulae - horizontal elements within the whole or a part of the corallite; trabecula fibrous microstructural element corresponding to vertically stacked cones and related to skeleton formation; wall peripheral calcareous element bounding the corallites. Rugosa - Pl. 83/3, 8, 10 Criteria: Calcitic skeleton. Solitary or colonial. Cupshaped, conical or cylindrical solitary corals ranging in diameter from a few millimeters to several centimeters. Centimeter- to meter-sized massive or dendroid colonies with variously integrated corallites. Major and minor septa. Long major septa inserted serially in four positions; shorter minor septa inserted between major septa. Septa usually with trabecular microstructure. Dissepiments and tabulae always present. Axial structures present or absent, often of septal origin. Walls formed by laterally fused septa dissepiments of additional thickened stereoplasm. Significance: Many rugose corals serve as good indicators of paleoenvironment and age. Distribution: Ordovician to Late Permian. Tabulata - Pl. 41/1; Pl. 83/1, 4, 7, 9, 11; Pl. 135/1, 2, 4; Pl. 143/7, 8 Criteria: Calcitic skeleton. Exclusively colonial. Sheetlike to dome-shaped, or erect-branched, or chain-like forms. Colonies generally small, but colonies with a diameter of several meters are also known. Slender tubelike corallites ranging in size between 0.5 to about 5 mm. Pores or tubes between corallites. Septa absent or short, commonly expressed as rows of spines or low ridges. Tabulae common and numerous. Dissepiments present or absent. Significance: Important reef-building organisms in the Silurian and Devonian, often together with stromatoporoid sponges. Distribution: Early Ordovician to Late Permian. Common in Middle Ordovician, Silurian, and Devonian. Less common in the Carboniferous and Permian.
Heliolitida - Pl. 83/5, 6 Criteria: Calcitic skeleton. Exclusively colonial. Massive, dome-shaped or laminar forms. Slender corallites are separated by extensive coenosteum consisting of smaller tubes. Septa generally present, commonly 12 in number, either spinose or laminar. Tabulae common. No evidence of communication between corallites. Similar to tabulate corals in having corallites with tabulae but different in having corallites embedded apparently within a colonial tissue developed between the polyps. Significance: Sometimes reef builders. Distribution: Late Ordovician to Middle Devonian. Heterocorallia - Pl. 83/2 Criteria: Calcitic skeleton. Elongate solitary corallites with four axially connected protosepta. Subsequent septa in each quadrant remain attached to earlier septa, dividing successively to form characteristic Y-shaped patterns. Tabulae present. Significance: Predominantly restricted to the Viséan, useful as index fossils. Distribution: Uppermost Devonian and Early Carboniferous. Scleractinia - Pl. 84/1-6, Pl. 115/1, Pl. 116/3, Pl. 117/1; Pl. 134/2; Pl. 147/1, 2, 4 Criteria: Aragonitic skeleton. Solitary or colonial. Varying strongly in shape and colony form. Sixfold symmetry of septal insertion (hence the name hexacorals). Complex microarchitecture and microstructure of septa and walls. Significance: Prominent reef builders since the Late Triassic and even more since the Late Jurassic. High carbonate production in shallow-marine environments and periplatform areas by accumulation of coral debris. Distribution: Middle Triassic to Holocene.
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Paleozoic Corals
Plate 83 Paleozoic Coral Groups The plate displays examples of the two major Paleozoic coral groups and another small group (see Box 10.6 for definitions). Rugose corals (–> 3, 8, 10) are known from the Middle Ordovician to the Late Permian. Rugose corals were widespread and common by the Silurian and increased in number and diversity in the Devonian. Following a marked extinction event at the Frasnian-Famennian boundary and at the end of the Devonian, new families with new morphological features (e.g. complex axial structures, –> 10) arrived in the Early Carboniferous and persisted until the end of the Permian. Tabulate corals (–> 1, 4, 7, 9, 11) received their name from horizontal partitions within the corallites, known as tabulae. The group was common from the Early Ordovician to the Late Permian. Diversification took place at the end of the Ordovician. Tabulate corals played a significant role in the formation of Silurian and Devonian reefs. Many tabulate coral taxa disappeared in the Late Devonian. Late Paleozoic tabulates are rare. Heterocorallia (–> 2) restricted to the latest Devonian and Early Carboniferous differ from all other corals by a specific pattern of septa insertion. All these groups apparently had primarily calcitic skeletons. 1
Tabulate corals. Encrusting colony of Alveolites (A) intergrown with stromatoporoids (S). The colony consists of inclined corallites opening obliquely at the surface and appearing rounded polygonal in transverse section. Adjacent corallites are connected by mural pores. Matrix is a peloidal packstone. Reef facies. Late Devonian (Frasnian): Canning Basin, western Australia. 2 Heterocorals. The solitary corallites of this small group are characterized by a specific shape and arrangement of the septa with four axially joined initial septa and later septa forming a Y-shaped pattern. The thin wall consists of microgranular calcite. Three cross-sections of Hexaphyllia mirabilis Duncan. Heterocorals are common fossils of Viséan inner shelf environments but are also known from the latest Devonian. Matrix is a peloidal micrite with benthic foraminifera (bottom right). Early Carboniferous (Marbella Formation, Late Viséan): Betic Cordillera, southern Spain. 3 Colonial rugose coral. Cross section of Argustrea quadrigemina Crickmay. The walls of adjacent corallites are united, forming a massive cerioid corallum (coral skeleton). The blister-like elements between the septa are dissepiments corresponding to small convex plates. Middle Devonian (Givetian): Ermberg, Rheinisches Schiefergebirge, Germany. 4 Tabulate corals. Oblique and longitudinal sections of Thamnopora sp., a widely distributed tabulate in Devonian lagoons or restricted back-reef environments. Note the conspicuously black matrix. This is a common feature of Devonian shelf carbonates, deposited in protected environments rich in organic matter. Early Devonian (Emsian): Cantabrian Mountains, Spain. 5 Heliolitid coral. Heteromorph colony of Pachycanaliculata barrandei (Penecke). The slender corallites are separated by a zone (coenosteum) composed of narrow polygonal tubes. Heliolitids have been included within the tabulate corals because of the existence of tabulae both in broad corallites and narrow tubes, but are now considered to represent a separate coral group because of the development of distinct equal septa. Middle Devonian (Eifelian): Graz, Austria. 6 Heliolitid coral. Longitudinal section. Corallites and the smaller tubes of the coenosteum differ in the size and frequency of the tabulae. Late Devonian (Frasnian): St.Barbe Quarry, Dinant Basin, Belgium. 7 Tabulate coral. Longitudinal section of Favosites styriacus Penecke. The corallites are subdivided by numerous, evenly spaced horizontal partitions (tabulae). Adjacent corallites communicate through mural pores (arrows) arranged in rows. Same locality as –> 5. 8 Colonial rugose coral. The cross section of Multimurinus shows a cerioid growth pattern characterized by corallites that are touching each other. Note the distinct axial structure (arrow), a common feature of Late Paleozoic Rugosa. Tethyan Middle and Late Permian bedded shelf sediments often contain intercalated banks consisting of these tens of centimetersized compound corals, which in places can form frame- and bafflestones (Weidlich and H. Flügel 1995). Middle Permian (Wordian): Wadi Aday, southwestern Muscat, Oman. 9 Tabulate corals. Favositid (F) and thick-walled encrusting auloporid (A) corals. Auloporid corals acted as bafflers, contributing to mound formation. C: Crinoid ossicle, T: Trilobite debris. Late Carboniferous (Kasimovian): Carnic Alps, Italy. 10 Rugose coral. The cylindrical corallites of Siphonodendron are not in contact and form a fasciculate structure. Note the peripheral zone formed by dissepiments and the distinct axial structure. The latter is a characteristic feature of Late Paleozoic rugose corals. Cross section. Early Carboniferous: Western Anatolia, Turkey. 11 Tabulate coral: Parastriatopora. The peripheral margins of the coral have been destroyed. This might indicate abrasion during transport or dissolution within the sediment. Early Devonian (Emsian): Erfoud, Anti-Atlas, Morocco.
Plate 83: Paleozoic Corals
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are defined as those anthozoans that secrete massive external calcium carbonate skeletons. These corals include the Rugosa, Tabulata and Scleractinia as well as some minor groups. For definitions of these groups see Box 10.6. Ecology: Major controls on the distribution of modern scleractinian corals include surface water temperature, light, symbiosis, and surface circulation (see Veron 1995 for a critical discussion). The symbiosis of corals and endosymbiotic algae has many physiological advantages for corals: enhancement of calcification, removal of metabolic waste, and concentration and recycling of nutrients. The term zooxantellae refers to endosymbiotic photosynthesizing dinoflagellates, red and green algae as well as cyanobacteria. The terms zooxanthellate and azooxanthellate applies to corals liv-
Corals
ing with or without benefit of algal symbiosis (Schuhmacher and Zibrowius 1985). Zooxanthellate corals are restricted both bathymetrically and geographically by their algal symbionts and are therefore concentrated in shallow tropical and subtropical waters where they build coral reefs. Azooxanthellate corals (Pl. 147/1, 2, 4) occur both in shallow and deep environments, but are more abundant at greater water depths of several hundreds of meters and more. Deep-water scleractinian corals with constructional frameworks (e.g. Lophelia, Pl. 4/ 3) form extended reef structures (Sect. 2.4.4.3). Azooxanthellate deep-water corals may have existed since the Early Mesozoic (Stanley and Cairns 1988). Skeletal mineralogy, microstructure and preservation: The originally aragonitic skeletons of scleractinian corals become unstable once removed from marine
Plate 84 Jurassic Scleractinian Corals Scleractinia or hexacorals (because of the sixfold symmetry of septal insertion) are colonial or solitary corals with an aragonitic skeleton. The group became diversified and widely distributed in reefs and shelf environments in the Late Triassic and continued to expand during the Jurassic and Cretaceous. They were important reef builders in the Late Jurassic and during the mid and Late Cretaceous. The diversity and importance as reefbuilders increased starting with the Eocene and reached a peak in the Miocene, when most of the modern reef corals originated. The plate displays growth types and morphological characteristics of Jurassic scleractinian colonial corals. In massive skeletons adjacent corallites may correspond to the astraeoid type (without walls separating individual corallites, septa in contact), maeandroid type (–> 3) or cerioid type (–> 6). The corallites of bushy, fasciculate colonies may be arranged parallel to each other (phaceloid, –> 1, 5) or branch irregularly in a tree-like pattern. Jurassic corals, constructing frameworks and baffling fine-grained sediment, contributed to the formation of reefs in high-energy shelf and platform-margin environments as well as on ramps. Environmental demands on Jurassic corals might have been considerably different from those of modern scleractinians (Leinfelder et al. 2002). 1 Growth type. Transversal section of Stylosmilia corallina Koby. The colony consists of separated non-integrated corallites which are more or less parallel (phaceloid growth type). Late Jurassic (Tithonian): Madonie Mountains, Sicily, Italy. 2 Reef corals: Transversal section of Calamophylliopsis flabellum (Michelin). The coral colony is encrusted (black arrows) and bored (B). The peloidal bindstone texture with spar-filled fenestrae (F) between the corallites indicates microbial binding of bioclastic material. Middle to Late Jurassic. Pierre Chatel, south of Grenoble, France. 3 Growth type. Oblique section of Myriophyllia rastellina (Michelin). The closely integrated chains of corallites of the maeandroid colony are arranged in linear series (black arrows) within the same walls. The exsert edges of the septa and walls form a ridge (white arrows). The structure is caused by intratentacular budding where new polypes and corallites are formed through division within the parent corallites. Same locality as –> 1. 4 Reworked corals. The thin section exhibits an allochthonous assemblage of broken and redeposited colonial corals, including Isastraea sp., I, Mesomorpha sp., M, and Thamnasteria gracilis (Münster), T. Note the absence of definite walls in Thamnasteria and Mesomorpha; the corallite centers are closely united by confluent septa. In thin sections coral debris can be recognized not only in fragments exhibiting skeletal elements but also in bioclasts smaller than 0.20 mm (Stanton and Flügel 1989). Jurassic exotic blocks floating in Cenozoic deposits: Northern margin of the Alps, Salzburg, Austria. 5 Reef corals: Transversal section of Thecosmilia longimana (Quenstedt). The bushy colony consists of laterally free corallites forming tufts of parallel corallites (phaceloid growth type). Note the biogenic encrustation (arrows) and the complete filling of the interspaces between the corallites by binding and encrusting organisms contributing to the formation of an organic framework. Late Jurassic: Same locality as –> 1. 6 Growth type. Transversal section of Latistraea foulassensis Beauvais characterized by closely appressed prismatic corallites whose walls are fused to each other (cerioid growth type). Late Jurassic: Untersberg, Salzburg, Austria. –> 1–6: Determination by D. Turnsek (Ljubljana)
Plate 84: Jurassic Scleractinian Corals
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Bryozoans
Fig. 10.38. Preservation of rugose corals compared with that of scleractinian corals. A: Late Permian rugose coral from Oman. All the diagnostic skeletal elements (septa, dissepiments, axial structure, external wall) are well preserved. B: Late Jurassic scleractinian coral from the Northern Calcareous Alps. Diagnostically important skeletal elements are only partly preserved. The thickness and shape of the septa are moderately modified by diagenesis, but the center and the periphery of the corallites are strongly affected. Scale is 2 mm.
environments and are susceptible to dissolution and neomorphism under burial or subaerial conditions. Dissolved skeletal elements may be preserved as spar-filled or sediment-filled molds. Identification of scleractinian corals requires recognizing the microstructures of calices and septa, but these fine structures are often destroyed by diagenesis. Sometimes weak neomorphism of coral skeletons can lead to the preservation of the internal microstructures. Common diagenetic modifications of scleractinian coral skeletons include the thickening of skeletal elements and the formation of pseudoskeletal elements. Rugose corals apparently secreted a calcitic skeleton and are therefore not as strongly affected by diagenetic modifications as scleractinians (Pl. 29/2; Fig. 10.38). Coral structures in thin sections: Plate 83 and 84 show some of the characteristic features of rugose and tabulate corals, heterocorals and scleractinians seen in thin sections. Basics: Corals Hill, D. (1981): Rugosa and Tabulata. – Treatise on Invertebrate Paleontology, Coelenterata, supplement 1. – 762 pp., Lawrence (Geological Society of America) Oliver, W.A., Coates, A.G. (1987): Phylum Cnidaria. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 140-193, Palo Alto (Blackwell) Schuhmacher, H., Zibrowius, H. (1985): What is hermatypic? A redefinition of ecological groups in corals and other organisms. – Coral Reefs, 4, 1-9 Scrutton, C. (1999): Palaeozoic corals: their evolution and palaeoecology. – Geology Today, 15, 184-193 Sorauf, J.E. (1993): The coral skeleton: analogy and comparisons. Scleractinia, Rugosa and Tabulata. – Courier For-
schungsinstitut Senckenberg, 164, 63-70 Veron, J.E.N. (1995): Corals in space and time. The biogeography and evolution of the Scleractinia. – 321 pp., Sidney (UNSW Press) Wells, J.W. (1956): Scleractinia. - Treatise on Invertebrate Paleontology, F, Coelenterata. – 328-440, Lawrence (Geological Society of America) Further reading: K119, K120, K121
10.2.3.4 Bryozoans Bryozoan carbonates reflect specific depositional settings and paleoclimatic conditions. Bryozoans are predominantly marine, active suspension feeders forming millimeter- to a several centimeter-sized colonies occurring from tidal level down to bathyal depths. They dominate in shelf seas, are abundant in depths between the intertidal zone and about 80 m and live both in low and high latitudes. Morphology: The colony (zoarium) consists of interconnected individuals (zooids) usually living within calcareous tube- or boxlike chambers (zooecia) that may or may not exhibit a number of transverse plates. An opening (apertura) is developed at one end of the zooecia. Zooecia and apertures exhibit strong polymorphism reflecting different vital functions of zooids. The colonies have wide ranges in form and size. They may be massive, encrusting, stick-like or correspond to a delicate branching network. The taxonomic differentiation of bryozoans requires thin-section studies. The skeleton of modern marine bryozoans consists of Low-Mg calcite or intermediate-Mg calcite (Cyclo-
Bryozoans
stomata) or aragonite (Cheilostomata). Differences in skeletal mineralogy may control diagenetic pathways of bryozoan cold-water carbonates (Bone and James 1993). Classification: Modern Bryozoa include the marine classes Stenolaemata and Gymnolaemata and the freshwater class Phylactolaemata, the first two with skeletons, the last without a skeleton. The Stenolaemata are characterized by elongated tubular or conical zooids and tube-like zooecia. The class covers the wealth of bryozoans found in limestones of Ordovician through the mid-Cretaceous age. By the Late Cretaceous the stenolaemates were exceeded in diversity by the gymnolaemates but continue to flourish in marine communities. The Stenolaemata include the orders Trepostomata (Early Ordovician to Triassic), Cystoporata (Early Ordovician to Triassic), Cryptostomata (Early Ordovician to Late Permian), Fenestrata (Early Ordovician to Triassic), and Tubuliporata (Early Ordovician to Holocene). The Gymnolaemata are the morphologically most varied class and exhibit box-shaped zooecia. It includes the orders Ctenostomata (uncalcified, Ordovician to Holocene) and Cheilostomata (calcified, Late Jurassic to Holocene). The greater part of the gymnolaemate diversity has evolved since the mid-Cretaceous. Ecology: Most colonies are attached to hard material such as rock, shells or sediment grains. Intertidal and upper subtidal encrusters are common on the undersides of hard substrates. Other habitats of encrusting bryozoans are long-lived stable hard substrates (reefs, rock walls). Bryozoans adhere also to flexible soft substrates, e.g. sea weeds and algal fronds. The majority of bryozoan habitats is found in areas of low sedimentation. Paleoenvironmental significance of bryozoans: Bryozoans are excellent paleoenvironmental and paleoclimatic indicators particularly for temperate and coldwater carbonates. The presence or absence of bryozoans, diversity and abundance, zooid morphology, colonial plasticity and growth form as well as reef-like structures provide information on the physical controls on the habitat. Paleoenvironmental parameters, such as as temperature, salinity, water energy, character of the substrate, and sedimentation rate may be indicated by bryozoans (Smith 1995). Modern cyclostome and cheilostome bryozoans show distinct zonation of growth form assemblages and special growth-form morphologies depending on the environment (Stach 1936; Nelson et al. 1988; James and
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Bone 1991; James et al. 1993). These differences are used in interpreting the water energy, substrate and temperature conditions of ancient bryozoan-bearing carbonates. Differences in colony growth forms are used as paleobathymetrical proxies (for shallow and deep shelf environments) and indicators of substrate stability (Hageman et al. 1998). This approach is successful in the study of Cenozoic bryozoan carbonates (Herbig 1991) and sometimes also Mesozoic carbonates, but has limitations in the study of Paleozoic bryozoans (Kelly and Horrowitz 1987), see also Pl. 85/3. Carbonate production: Recent bryozoans are common in sediments of the continental shelves, but also occur in reefal environments (Cuffey 1972). Bryozoan bioclasts are abundant constituents of modern and ancient cool-water shelf carbonates contributing to the formation of carbonate particles of all grain sizes. Grain sizes of bryozoan carbonates are strongly controlled by skeletal architecture, as shown by the bryozoan sands, muds and gravels on the southern Australia Lacepede Shelf. Skeletal architecture is related to the strength and disintegration potential of bryozoan skeletons (Cheetham and Thomsen 1981). A significant number of modern bryozoans live attached to sessile, benthic invertebrate hosts, which are not or only slightly calcified and leave no or little body fossil record. Bryozoan sediment production from epizoans on ephemeral substrates may have been important in the Tertiary (Hageman et al. 2000). Bryozoans can contribute to Cenozoic shelf deposits with a quantity of up to 95% (e.g. off New Zealand). Huge areas of the Pacific deep shelf bottom between western Australia, New Zealand and the Chatham Islands extending about 6000 km are dominated by bryozoan sediments. This ‘bryomol facies’ is a characteristic feature of non-tropical carbonates (see Sect. 12.2). Bryozoan-rich microfacies types are common in platform and ramp carbonates throughout the geological past, particularly in Paleozoic shelf limestones and may contribute to the formation of reefs (e.g. Permian Zechstein reefs; Pl. 85/11; Early Tertiary Danian reefs, Pl. 147/2, 3). Distribution: Bryozoans are known since the Ordovician. Five of six known Paleozoic orders became extinct in the Triassic, but the Cyclostomata persisted from the Ordovician to the present, dominating many faunas in the Mesozoic. The Cheilostomata are the other important modern marine group. They arose in the Late Jurassic and became dominant in the Late Cretaceous and Cenozoic.
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Bryozoans in thin sections (Pl. 85): Bryozoan sections can be distinguished from other fossils consisting of calcareous tubes (e.g. chaetetid sponges) by the presence of polymorphic tubes of various widths, and by a characteristic fanlike pattern in longitudinal and oblique sections, resulting from a different arrangement of the zooecia in the inner axial zone of the colony and near the surface (Pl. 85/2, 7). Paleozoic trepostome bryozoans may display morphologically different tubes in addition to the zooecia, differing in the diameter and the development of horizontal elements. Small acanthopores have a laminated cone-in-cone structure pointing outward within the laminated wall. Mesopores of Early Paleozoic trepostome bryozoan contain numerous transverse plates, in contrast to the mesopores of Late Paleozoic taxa. Many Paleozoic trepostome and cryptostome tubular bryozoans exhibit zooecia walls that are thin in the inner zone of the colony and thick in
Bryozoans
the outer zone. Paleozoic cyclostome bryozoans are characterized by simple tests with walls of constant thickness and generally without internal structures. Post-Paleozoic cyclostomes exhibit a thick finely laminated wall structure in the outer zone of the colony and minutely porous walls. Paleozoic cystoporate bryozoans have a vesicular to cystose framework between the zooecia and simple tubes which are commonly modified by a hood (lunarium) at their outer ends. This structure is reflected internally in tangential sections by a crescent-shaped notch in the wall and a thickening of the wall tissue behind the notch (Pl. 85/4). Bryozoan wall microstructure may be laminated, granular or fibrous. The laminated structure occurs in many bryozoan groups and is one of the distinctive features of bryozoan debris (Pl. 85/4). Details of the microstructural types can be found in Horowitz and Potter (1971).
Plate 85 Bryozoans The plate displays thin sections of bryozoans belonging to different systematic units. Characteristic thin-section features of bryozoans are (a) colonies consisting of variously sized tubes (–> 1), (b) different arrangement of tubes in the inner and outer part of the colony (–> 3) resulting in a fan-like pattern (–> 7), (c) window-frame network structures (–> 10), and (d) finely laminated wall structures (–> 4, 10). See also Pl. 4/5; Pl. 9/5, 6; Pl. 51/ 6; Pl. 106/1-3; Pl. 107/1, 2; Pl. 144/7, and Pl. 147/2, 3 for further photographs. 1
Trepostome bryozoans (Reptonoditrypa cautica Schäfer and Fois), characterized by polymorph tubes (zooecia). Oblique sections. Sections of trepostome bryozoans are characterized by robust massive or dendroid colonies and small tubes between larger ones. Middle Triassic (Anisian): Olang Dolomites, Southern Alps, Italy. 2 Cystoporid bryozoan are characterized by encrusting or massive colonies consisting of large zooecia (Z) embedded within a framework of smaller tubes (cystopores (C). Oblique section. Early Carboniferous: Nötsch, Austria. 3 Bifoliate cyclostomid bryozoan. Note the differentiation of central and peripheral parts of the colonies. The vinculariid erect growth form with dichotomously branched zooecia indicates a protected environment. Late Tertiary (Miocene): Subsurface, Suez region, Egypt. 4 Broken branch of a cystoporid bryozoan. The cross section of zooecia (Z) showing the typical crescent-shaped notch in the wall is infilled with micrite and embedded in a thickened wall exhibiting a lamellar structure (arrow). Mississippian: Turtle Mountains, Frank Slide, southeastern Alberta, Canadian Rocky Mountains. 5 Cryptostome bryozoan. Cryptostome bryozoans are characterized by delicate dendroid or bifoliate colonies, axial zooecia bundles and in cross sections short elliptical zooecia in the peripheral zones. Oblique section. Ordovician (Caradoc, Macrourus limestone): Öland Island, Sweden. 6 Fenestrate bryozoan exhibiting cross sections of zooecia and lamellar wall structure. Middle Permian: Straza, Slovenia. 7 Cryptostome bryozoan. Transverse (T) and longitudinal (L) section. Note radial distribution and the curvature of zooecia toward the outer margin and the thickened wall structure at the exterior margin (arrow). Cold-water carbonate. Permian (Berriedale limestone): Southeastern Tasmania, Australia. 8 Fenestrate bryozoan with serially arranged zooecia on the surface of a broken frond. Mississippian (Tournaisian): Mouse Mountains near Calgary, Alberta, Canada. 9 Cystoporid bryozoan (Fistulipora). Oblique section of a broken colony. Note zooecia (Z) separated by vesicular tissue (cystopores, C). Early Permian: Powow Canyon, Hueco Mountains, Texas, U.S.A. 10 Fenestrate bryozoan. Outer part of a fan-shaped colony. Zooecia occurring on the surface of the colony (inner part of the fan) can not be seen in this section. Fenestrate bryozoans are characterized by erect rigid fronds connected by cross bars forming a characteristic network with open meshes (fenestrulae, F). Note lamellar wall structure. Most members of this group preferred quiet sheltered environments. Mississippian: Turtle Mountains, Frank Slide, Alberta, Canada. 11 Bryozoan debris consisting of fragments of fenestrate bryozoans. Reef facies. Late Permian (Zechstein): Bartolsfelde, Lower Saxony, Germany.
Plate 85: Bryozoans
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Basics: Bryozoans Boardman, R.S., Cheetham, A.H. (1987): Phylum Bryozoa. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 497-549, Palo Alto (Blackwell) Bone, Y., James, N. (1993): Bryozoans as carbonate sediment producers on the cool-water Lacepede shelf, southern Australia. – Sedimentary Geology, 86, 247-271 Brood, K. (1978): Bryozoa. – In: Haq, B.U., Boersma, A. (eds.): Introduction to marine micropaleontology. – 189201, New York (Elsevier) Conti, S., Serpagli, E. (1988): Bimineralic (calcareous and phosphatic) skeleton in Late Ordovician bryozoa from Sardinia: geological implications. – Bolletino della Societa Paleontologica Italiana, 27, 129-162 Cuffey, R.J. (1972): The role of bryozoans in modern coral reefs. – Geologische Rundschau, 61, 542-550 Gordon, D.P., Smith, A.M., Grant-Mackie, J.A. (eds., 1996): Bryozoans in space and time. – 442 pp., Wellington (National Institute of Water and Atmospheric Research) Hageman, S.J., Bock, P.E., Bone, Y., McGowran, B. (1998): Bryozoan growth habits: classification and analysis. – Journal of Paleontology, 72, 418-436 Hageman, S.J., James, N.P., Bone, Y. (2000): Cool-water carbonate production from epizoic bryozoans on ephemeral substrates. – Palaios, 15, 33-48 McKinney, F.K., Jaklin, A. (2001): Sediment accumulation in a shallow-water meadow carpeted by a small erect bryozoa. – Sedimentary Geology, 145, 397-410 Robison, R.A. (ed., 1983): Bryozoa (revised): introduction, order Cystoporata, order Cryptoporata. – Treatise on Invertebrate Paleontology, 644 pp., Lawrence (Geological Society of America and University of Kansas)
Shell Microstructures
Scholz, J. (2000): Eine Feldtheorie der Bryozoen, Mikrobenmatten und Sedimentoberflächen. – Abhandlungen der Senckenbergischen naturforschenden Gesellschaft, 552, 1-193 Smith, A.M. (1995): Paleoenvironmental interpretation using bryozoans: a review. – In: Bosence, D.W., Allison, P.E. (eds.): Marine paleoenvironmental analysis from fossils. – Geological Society of London, Special Publication, 83, 231-243 Stach, L.W. (1936): Correlation of zooarial form with habitat. – Journal of Geology, 44, 60-65 Taylor, P.D. (1998): Bryozoan carbonates through time and space. – Geology, 26, 459-462 Further reading: K122
10.2.4 Shells The fourth part of the discussion of thin-section fossils deals with shell-bearing organisms characterized by bivalved or univalved shells (brachiopods, mollusks, serpulids, ostracods) or by multi-element/multi-plate shells (trilobites, crustaceans, echinoderms). Most of these fossils are too large to be included in a standard microfacies thin section that will most commonly only exhibit parts or fragments of the skeletons. An identification of these bioclasts as to higher systematic levels is possible when original shell microstructures are still preserved or unique diagnostic criteria are visible. The
Box 10.7. Some microstructure types occurring in shells of brachiopods, mollusks, serpulids and ostracods. Definitions based on Majewske (1969) and Carter (1980). Complex crossed-lamellar: Characterized by a disorderly arrangement of the elements of the crossed-lamellar microstructure. Found in bivalves and gastropods. Crossed: Consisting of elongate structural subunits organized into adjacent, interdigitating aggregations referred to as first-order lamellae. Includes crossed-lamellar and complex crossed-lamellar microstructures. Common in bivalves and gastropods. Crossed-lamellar: Consisting of numerous branching and wedging lamellae oriented perpendicular or slightly inclined to the plane of the shell layer. Each lamella consists of small crystals inclined at a constant angle and inclined in opposite directions in adjacent lamellae. Most perfectly developed in aragonitic shells but also occurring in calcitic shells. May constitute all microstructural layers in mollusk shells or may alternate with other microstructures. Common in gastropods and bivalves. Foliated: Composed of numerous very thin calcite lamellae which may be uniformly oriented over long distances or randomly oriented over short distances. Common in brachiopods, some bivalves and gastropods, and some serpulids.
Homogeneous: Characterized by orderly arrangement of uniformly small crystals. Appears bright in transmitted light and will progressively extinguish under crossed nicols if the stage is rotated. Common in ostracods and aragonitic mollusk shells. Laminar: Characterized by sheets of strongly flattened structural units developed uniformly parallel or oblique to the surface. Includes aragonitic nacreous microstructures and calcitic foliated microstructures. Brachiopods, bivalves, gastropods. Nacreous: Refers to an orderly arrangement of numerous thin aragonitic laminae, each separated by thin films of organic material and uniform orientation (‘mother-ofpearl’). Common in bivalves, gastropods and cephalopods. Prismatic: Consisting of closely packed, generally parallel prisms of calcite or aragonite, generally oriented with their length perpendicular or slightly inclined to the plane of shell layers. Prism columns are mostly longer than wide. Common in bivalves. Spherulitic: Consisting of densely packed spherical to subspherical aggregates of elongate structural subunits which radiate in all directions from a central nucleation site. Rare in aragonitic bivalve shells.
Brachiopods
study of microstructures in thin sections requires magnifications of at least x30 and the use of polarized light. Staining techniques (Adams and MacKenzie 1998) and cathodoluminescence microscopy are strongly recommended. In-depth evaluations of microstructural features are based on SEM data. A comparison of these data with thin-section observations is sometimes difficult, partially because of different definitions of microstructures seen in SEM and in thin sections. The terms listed in Box 10.7 are those commonly applied to microstructures exhibited in thin sections.
10.2.4.1 Brachiopods Brachiopods are bilaterally symmetrical sessile marine organisms with an external shell consisting of two dissimilar but equilateral valves. Brachiopods are major contributors to the bioclastic content of shallow-marine limestones, particularly in the Paleozoic. Morphology and classification: Traditionally brachiopods have been split up into two major groups, the Inarticulata and the Articulata. The Inarticulata dominated in the Cambrian and Ordovician; they received their name from the fact that the two valves do not have an articulating hinge and are united by muscles. Their shells are composed of alternating layers of calcium phosphate and chitin. All other brachiopods included within the Articulata have articulating hinges and calcitic shells. The main groups of the Articulata are the Strophomenida (Ordovician to Middle Jurassic; including the Late Paleozoic productids and brachiopods that are attached with one cone-shaped valve at the substrate, while the other valve acts as a lid; Pl. 86/2), Orthida
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(Middle Cambrian to Permian), Pentamerida (Late Cambrian to Devonian), Rhynchonellida (Ordovician to Holocene, abundant in the Mesozoic), Spiriferida (Ordovician to Early Jurassic, extremely diverse during the Devonian), and the Terebratulida (Carboniferous to Holocene), the most common brachiopods today. The body of brachiopods is enclosed by a valve except for the pedicle which serves to attach the animal to the substrate. The larger valve on the ventral side of the body is the pedicle valve, the smaller valve on the dorsal side the brachial valve, taking its name from the brachia. These are arm-like projections of the lophophore, a feeding organ that also carries them. Shell structures: All articulate brachiopods have Low-Mg calcite shells. The shell grows continually throughout the life of the animal; new material is laid down around the larger rounded edge of the shell. The original wall structure is commonly well preserved. The calcitic shell is overlain by a thin chitinous layer (periostracum), never preserved in fossil shells. The calcitic shell consists of two layers - (1) a thin outer primary cryptocrystalline layer of fairly constant thickness, and (2) an inner, thicker layer that thickens with growth and from which all internal structures are formed. The primary layer is difficult to detect, but it may be recognizable in shells of different brachiopod groups. The boundary between primary and secondary layers is commonly sharp. The characteristic feature of brachiopod shells is a foliated lamellar structure of the secondary layer appearing as thin parallel laths of calcite which are variably fused and amalgamated, commonly having a wavy banded appearance (Fig. 10.39). The laminae are oriented roughly parallel to the outer surface of the valve.
Fig. 10.39. Laminar microstructure of strophomenid brachiopod shells characterized by wavy lamellae that are buckled adjacent to the pseudopunctae. The latter feature distinguishes the shell from bivalve shells, which may exhibit laminar microstructures as well. Middle Devonian (Givetian): Letmathe, Sauerland, Germany. Scale is 1 mm.
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Shells consisting of calcite prisms are a common feature of rhynchonellid brachiopods. The basic types of brachiopod shell microstructures were summarized by Williams (1990). In punctate brachiopods the shells are perforated by tiny open tubes extending almost to the outer surface (e.g. terebratulids). In pseudopunctate brachiopods (most strophomenids, including productids and chonetids) the shell has solid rods of calcite reminiscent of true punctation that end within the fibrous layer; the fibers contributing to the formation of pseudopunctae are buckled into a series of superimposed cones. In impunctate brachiopods the shells are imperforate and dense all the way through (pentamerids, most rhynchonellids and spiriferids). The first punctated (endopunctated) shells occur in the Ordovician. Pseudopunctate and impunctate shells are already known from the Cambrian. Ecology: Brachiopods are active filter feeders commonly attached to hard and firm substrates. Modern brachiopods live in all parts of the sea as sessile organisms attached to the substrate by a pedicle. Today bra-
Brachiopods
chiopods are also found in temperate and cool waters of normal salinity. Many articulate species are deepwater specialists and occur in depths down to several hundreds of meters and more. Inarticulate forms, e.g. Lingula, live in relatively shallow coastal waters. Brachiopod shells as environmental recorders: Brachiopod shells are widely used as paleoenvironmental indicators in the reconstruction of ancient seawater temperatures (Veizer et al. 1986; Curry and Fallick 2003). Because Low-Mg calcite is less susceptible to diagenetic alteration, it is more likely that brachiopod shells have retained the isotopic composition in equilibrium with the environment in which the organisms lived. Possible diagenetic alterations are commonly determined by cathodoluminescence (Barbin and Gaspard 1995). Distribution: Brachiopods appeared in the Cambrian, were abundant in the Paleozoic and most diversified in the Devonian. There was a distinct drop in the number of taxa following the Permian-Triassic extinction event, but brachiopods continued to be part of the shelf fau-
Plate 86 Brachiopods The plate displays thin sections of articulate brachiopods characterized by two valves that are bilaterally symmetrical but dissimilar in size and shape (–> 1, 3). Internal structures are represented by a calcareous framework (brachidium) that supports the food-gathering system of the soft body. This framework takes several forms, e.g. calcareous ribbons coiled in helical spirals (–> 1), or elements supporting the brachidium (–> 3). Characteristic microstructures distinguish brachiopod shells from bivalve shells (–> 1, 5). 1 Microstructure and internal elements: Cross sections of articulate spiriferid brachiopods. Note the well-preserved internal brachialia (B), the median septum (MS) and the low-angle finely layered microstructure of the calcitic impunctate shell (arrow) characteristic of brachiopods. SMF 8. Middle Triassic (Anisian): Olang Dolomites, Southern Alps, Italy. 2 Richthofenia. Richthofeniacean brachiopods (order Strophomenida) differ from other brachiopods by a cone-shaped pedicle valve fixed by the apex to the seafloor and anchored by spines. The brachial valve acts as a lid-like structure. These brachiopods, together with sponges and microbes, contributed to the formation of Permian reefs in tropical and temperate paleolatitudes. The group is restricted to the Permian. Oblique section exhibiting the open framework of the pedicle valve. Late Permian: Chios Island, Greece. 3 Internal elements. Cross section of a smooth brachiopod shell cut parallel to the hinge line exhibiting the larger pedicle valve (right) and the smaller brachial valve. Note the broken interior double plates (arrows) corresponding to supporting elements of the brachialia. Bioclastic wackestone. Early Devonian (Emsian): Erfoud, southern Anti-Atlas, Morocco. 4 Brachiopod coquina. Most shells are cut longitudinally. Note microbial encrustations (arrow). Rock-building accumulations of brachiopod shells are common on shelves and in slope settings. Flanking beds of a microbial-stromatoporoid buildup. Late Devonian (Frasnian): Kielce, Holy Cross Mountains, Poland. 5 Microstructure. Tangential section of a punctate articulate brachiopod exhibiting very small tubules (punctae, P) within the shell. Early Jurassic (Hierlatz limestone): Maurach near Innsbruck, Tyrol, Austria. 6 Brachiopod spines: Some pseudopunctate brachiopods yield spines that are found as displaced elements in the sediment. Diagnostic criteria are the hollow central canal surrounded by concentric inner and thick radial-fibrous outer layers. Transverse section. These spines are common in Carboniferous and Permian shelf carbonates. Early Carboniferous: Holy Cross Mountains, Poland.
Plate 86: Brachiopods
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nas in the Mesozoic. From about the mid-Jurassic onwards, the articulate brachiopods retreated to cryptic tropical reef niches, or were confined to colder and deeper waters. Only three groups (rhynchonellids and terebratulids and the inarticulate Lingulata) crossed the Cretaceous-Tertiary boundary and are still present today. Brachiopod limestones: Accumulations of brachiopod shells are abundant in Paleozoic limestones (Pl. 102/1, Pl. 132/10) and common in Mesozoic limestones. Brachiopod limestones originated in different depositional environments (shelves, ramps, slopes) and became more frequent in the Silurian and Devonian. Shell beds composed of brachiopods can not be compared with bivalve shell beds because of differences in skeletal mineralogy (Low-Mg calcite vs. aragonite), a stronger tooth-socket hinge (holding the valves together after death), lack of ligament forcing the valves apart like bivalves, and low vulnerability to predators and boring organisms (Copper 1997). Some brachiopod shell beds are composed of in-situ assemblages of life and dead shells, others by transported assemblages deposited by currents within depressions of the sedimentary surface or accumulated on the surface as ‘dunes’ (Hagdorn and Mundlos 1982). The accumulation of shells related to a transport of brachiopods from shallow to deep-water environments is demonstrated by the Jurassic Hierlatz facies of the Calcareous Alps and the Carpathians. This facies is one of the most characteristic Early Jurassic facies in the western Tethys. It comprises typically variegated predominantly brachiopod limestones and crinoid-brachiopod limestones that overlie drowned Late Triassic carbonate platforms or occur as fillings of neptunian dikes. Associated biota are bivalves, gastropods and small ammonites. Brachiopods and crinoids lived originally in moderate water depths on the top of seamounts, from where they were transported as a bypass margin deposit to the slopes and foot of the seamounts (Vörös 1986; Uchman and Tchoumatchenco 1994). Basics: Brachiopods Cambell, J.D., Lee, D.E., MacKinnon, D.I. (eds., 1990): Brachiopods through time. – Proceedings of the second International Brachiopods Congress, University of Otago, Dunedin, New Zealand, 580 pp., Rotterdam (Balkema) Copper, P. (1997): Articulate brachiopod shell beds: Silurian examples from Anticosti, eastern Canada. – Geobios, 29, 133-148 Copper, P., Jin, J.S. (eds., 1996): Brachiopods. – 373 pp., Rotterdam (Balkema) Broadhead, T.W. (ed., 1981): Lophophorates. Notes for a short course. – 251 pp., Cincinnati (Paleontological Society) Popov, E.L,.Bassett, M.G., Holmer, L.E., Laurie, J. (1993):
Bivalves
Phylogenetic analysis of higher taxa of brachiopods. – Lethaia, 26, 1-5 Kaesler, R.L. (ed., 1997): Brachiopoda (revised), part 1. – Treatise on Invertebrate Paleontology, 60 pp., Lawrence (Geological Society of America) Kaesler, R.L. (ed., 2000): Brachiopoda (revised), part 2 and 3. – Treatise on Invertebrate Paleontology, 960 pp., Lawrence (Geological Society of America) Rowell, A.J., Grant, R.E. (1987): Brachiopoda. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 445-496, Palo Alto (Blackwell) Vörös, A. (1986): Brachiopod palaeoecology on a Tethyan Jurassic seamount (Pliensbachian, Bakony Mountains, Hungary). – Palaeogeography, Palaeoclimatology, Palaeoecology, 57,241-271 Williams, A. (1990): Biomineralization in the lophophorates. – In: Carter, J.G. (ed.): Skeletal biomineralization: patterns, processes and evolutionary trends. – Volume I, 6783,volume II (atlas), pls. 141-156, New York (Van Nostrand) Williams, A., Rowell, A.J. (1965a): Brachiopod anatomy. – Treatise on Invertebrate Paleontology, part H, volume I, 6-57, Lawrence (Geological Society of America) Williams, A., Rowell, A.J. (1965b): Brachiopod morphology. – Treatise on Invertebrate Paleontology, part H, volume I, 57-138, Lawrence (Geological Society of America) Further reading: K123
10.2.4.2 Bivalves Bivalves are the most important mollusks contributing to the bioclast content of limestones, both in marine environments (since the Cambrian, and abundant in the Mesozoic and Cenozoic) and in freshwater and brackish environments (since the Carboniferous). Morphology and classification: Bivalves (also called pelecypods or lamellibranchs) have a layered shell consisting of two valves. The valves may be either equal or unequal in shape and size. Most bivalves articulate along a hinge line by means of teeth and sockets. The plan of symmetry passes between the two valves parallel to the hinge line. Each valve contains a thin organic covering (periostracum) and multiple layers of carbonate. The carbonate shell layers consist entirely of calcite, entirely of aragonite or of a mixture of aragonite and calcite in alternating layers. Shell structure: Common calcitic microstructures are foliated, prismatic, crossed-laminar and homogeneous. Microstructures can sometimes be used to attribute shell fragments seen in thin sections to major taxonomic units. Excellent summaries of microstructural types can be found in Majewske (1974) and Carter (1980, 1990). Majewske provides a key which should enable families with related microstructures to be identified.
Shell Beds
Ecology: Bivalves live predominantly marine as endobenthic suspension feeders on soft bottoms and as epibenthos on the sediment. Some bivalves are borers. Bivalve limestones Many limestones yield high numbers of bivalve shells contributing to the formation of bioclastic wackestones, packstones, floatstones, and rudstones. Bivalve limestones originate in a variety of marine depositional settings, from the beach across the shelf and the shelf margin to the slope. To interpret these limestones basic concepts developed for the discussion of shell beds must be taken into account. Shell beds: Biofabric and taphonomic signatures of shell accumulations provide information on sedimentary processes, such as fair weather and storm waves, storm flows, and long-term currents as well as rates of sedimentation. Fürsich and Oschmann (1993) distinguished nine types of skeletal concentrations. Essential criteria are sedimentary structures (e.g. cross bedding, planar lamination), erosive base, grading, intraformational pebbles, bioturbation, matrix, packing of skeletal elements (bioclast-support or matrix-support),
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disarticulation, plan-view and side-view orientation, fragmentation, abrasion, bioerosion and encrustation, sorting, species diversity. Many of these criteria can only be seen in large thin sections if shell accumulations should be genetically interpreted. Shell concentrations (lumachelles, coquinas) can be formed by various processes (Kidwell 1991). Event concentrations originate from short-term events, e.g. storm deposits (Pl. 102/3, 5), composite concentrations are caused by multiple reworking and transport), condensed concentrations are formed by multiple reworking and transport of long duration associated with low sedimentation rates, and lag concentrations as a result of multiple reworking associated with the winnowing of finegrained sediment and corrosion. Shell beds should be studied in detail because they may be used as markers for recognizing the development of sequences and cyclothems (Carter et al. 1991). Transgressive system tracts and nearshore highstand system tracts contain either worn and transported bivalve shells, or in-situ faunas diagnostic of intertidal and shallow subtidal conditions. In contrast, the maximum flooding surface is characterized by an offshore fauna.
Fig. 10.40. Megalodon limestones. The fine-grained bioclastic wackestone contains abundant molds of bivalves, most of which are double-valved and still in life position. The megalodontids lived on shallow muddy substrates in lagoonal lowenergy environments. The well-bedded Megalodon limestones are a part of cyclic platform sequences in Oman as well as in other parts of the Triassic Tethys recorded by alternations of bivalve limestones and microbial/algal carbonates. Late Triassic (Mishrif Formation, Norian): Wadi Ala, Jabal Kawr, Oman Mountains. Coin for scale. Courtesy of M. Bernecker (Erlangen).
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Limestones with abundant bivalves representing allochthonous shell concentrations or autochthonous occurrences of bivalves (Pl. 87/4, Pl. 119/1) are key elements of characteristic facies types which sometimes can be found in worldwide distribution as demonstrated by the following examples: Triassic Megalodon limestones: The Late Triassic Dachstein limestone of the Northern Calcareous Alps as well as time-equivalent platform carbonates all over the Tethys (Fig. 10.40) are characterized by cyclic platform carbonates containing thick-bedded limestones with abundant shells of pachydont megalodontids. These equal-sized bivalves may become several tens of centimeters in size. They occur commonly in life
Lithiotis Limestones
position within a fine-grained sediment corresponding to wackestone and packstone with bioclastic debris and in places with some corals and dasyclad green algae. The Megalodon limestones represent the subtidal part of the Lofer cycles (see Pl. 140/4) . The depositional setting of Late Triassic megalodontid limestones is confined to the standard facies zone 8 (see Sect. 14.3). Liassic Lithiotis limestones: A specific limestone type characterized by the mass occurrence of Lithiotis, a thick-shelled heterodont bivalve several centimeters in size, and related genera, developed during the Early Liassic (Sinemurian, abundant in Pliensbachian and Toarcian) in carbonate platforms of Tethys and Panthalassa forming a low-latitude belt (Nauss and Smith
Plate 87 Bivalves This plate displays thin sections of Mesozoic and Cenozoic bivalve shells. Shell shape (-> 1) and microstructures (–> 5-7, 9) are used to differentiate bivalve shells from other mollusk shells and brachiopods. 1 Complete bivalve. Radial section through joined calcitic valves of an equivalved megalodontid bivalve. D: Dorsal, V: Ventral. The interior of the shell is filled partly with sediment, partly with recrystallized calcite cement. Late Triassic: Northern Alps, Austria. 2 Growth lines. Radial section through a recent aragonitic Donax shell. Note the widely and densely spaced growth lines (GL) and the crystals of the prismatic microstructure (white arrow). R: Rib area. Recent: North Sea. 3 ‘Filaments’. The fossils are larval or juvenile shells of pelagic bivalves. Filaments are common constituents of Mesozoic deep-marine basinal limestones, e.g. in the Ammonitico Rosso facies of the Mediterranean region. Note dense packing, bioclast-support and common bioclast/bioclast contacts. More than 70% of bioclasts lie within two adjacent size classes. High degree of close packing reflects episodes of gregarious deposition of pelagic shells on the sea bottom within a very short time interval. Early Jurassic: Korfu Island, Greece. 4 Bivalve reef. Densely spaced calcitic valves of fossil ‘oysters’ (Placunopsis ostracina) forming reef-like structures by attachment of the right valve to hard substrate. Example of a ‘primary biogenic shell concentration’ (Fürsich and Oschmann 1993) resulting from a gregarious habit and characterized by epifaunal bivalves growing one shell above the other. Note growth lamellae (GL) characterizing the ‘foliated microstructure’ of the shells, and sediment-filled borings (B). The high frequency of life-positioned oysters in the densely packed well sorted fabric together with the lenticular shape of the deposit indicates an accretionary buildup. Compare this fabric with the example shown in –> 3 and note the almost complete absence of carbonate mud due to failure in sediment supply. Middle Triassic (Muschelkalk): Berlichingen, southwestern Germany. 5 Microstructure. Bivalve shell fragment showing a finely prismatic calcite microstructure. Small circular black structures are microborings. Late Jurassic: Untersberg, Salzburg, Austria. 6 Microstructure. Fossil ‘oysters’. Note the microstructural differentiation of the calcitic valve into a prismatic outer layer and a foliated layer. Late Tertiary (Miocene): Ekisce, southern Turkey. 7 Microstructure. The Ostrea shell exhibits alternating foliated (F) and vesicular (V) layers characteristic of pycnodont bivalves. Late Tertiary (Miocene): Mesarevlev, southern Turkey. 8 Inoceramid bivalves: Longitudinal and cross sections through the outer layers (ostracum) of calcitic Inoceramus shells exhibiting the characteristic prismatic microstructure. Note the honeycomb pattern in plan view (center top) and the difference in prism size in the outer and inner shell layer. Individual prisms act as single crystals displaying unit extinction between crossed nicols. Inoceramid bivalves are known from the Jurassic to the Late Cretaceous. They were common during the Late Cretaceous and became extinct in the Maastrichtian well before the Cretaceous-Tertiary boundary. The fossils are important Cretaceous index fossils (McLeod and Ward 1990), provide interesting paleoclimatic proxies (Dhondt and Dieni 1992) and are used as diagenetic markers (Elorza and Garcia-Garmilla 1998). Disarticulated calcitic prisms of the ostracum are important constituents of Cretaceous open-marine limestones. Note growth lines (top left). Small sparfilled structures in the micritic matrix are calcispheres (see Pl. 66/7). Late Cretaceous (Cenomanian): Teutoburger Wald, Germany. 9 Microstructure. Bivalve wackestone with angular fragments of pycnodont shells showing vesicular and foliated microstructures. Middle Cretaceous (Aptian): Subsurface, Ras al Khaimah, United Arabian Emirates.
Plate 87: Bivalves
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Limestones with Filaments
1988; Scheibner and Reijmer 1999; Monaco 2000). The bivalves preferred different subenvironments and lived in intertidal and subtidal low energy, marginal marine as well as in high energy shelf-margin settings comprising the standard facies zones 4, 5, 7 and 8. Shell concentrations were controlled by physical and biological agents such as shell transport by currents and waves, winnowing fine-grained material leaving behind shell coquinas, bioturbation resulting in shell concentration, and in-situ growth contributing to the formation of lowdiversity bivalve reefs near platform margins and on upper slopes (Fig. 10.41). Fig. 10.41. Lithiotis limestone contributing to the formation of mounds near a platform-upper slope margin. The dark finegrained sediment corresponds to bioclastic pack- and wackestones. Early Jurassic: Jebel Bou Dahar, High Atlas Mountains, Morocco. Coin for scale. From Scheibner and Reijmer (1999).
Middle and Late Jurassic limestones with ‘filaments’: The name refers to fossils first recognized in thin sections of Jurassic limestones and later also in rock-building quantities in other Mesozoic open-marine limestones, e.g. the Late Triassic Hallstatt lime-
Plate 88 Cretaceous Rudist Bivalves Rudists are sessile gregarious bivalves characterized by a lower attached and an upper opercular valve of different sizes. The group appeared in the Late Jurassic and disappeared at the end of the Cretaceous. Rudists were common throughout the Cretaceous. Diversity dropped during the Early Aptian and the latest Cenomanian and increased significantly in the Early Maastrichtian. Major subgroups are differentiated by shell morphology, the number and position of teeth, accessory cavities, canals (–> 2), pores and pillars. Thin sections exhibit characteristic microstructural patterns. The shells consist of an inner and middle aragonitic layer (nearly always recrystallized and replaced by blocky calcite, –>5) and an outer Low-Mg calcite layer with compact (–> 1, 6) and/or cellular microstructures (–> 1, 4, 5). Caprinidae (–> 2; Aptian to Maastrichtian, common in Cenomanian) are characterized by characteristic canals within the shell. The shells of Radiolitidae (–> 1, 4, 5, 6; Late Aptian to Maastrichtian, common from the Late Albian onwards, abundant from Santonian to Maastrichtian) exhibit a reticulate pattern caused by the calcitic cellular prismatic outer layer, and a compact inner layer. Both groups were abundant in outer and mid-shelf environments. The Hippuritidae (–> 3; Turonian to Maastrichtian), forming densely packed biostromes, are characterized by pillars within the shell wall. 1 Radiolitidae. Lower valves of Lapeirausia sp. Oblique cross-section. Note pseudopillars (P) and thick outer layers showing the diagnostic alternation of the thick cellular boxwork ostracum and thinner compact microstructures. The interspaces between the shells are infilled with silt-sized prismatic calcite crystals (CC) resulting from the disintegration and spalling of parts of the shells. Radiolitids are common in outer to mid-shelf environments. SMF 8. Late Cretaceous (Campanian): Greece. 2 Caprinidae. Upper valve of Sphaerocaprina sp. Transverse section. Note the inner shell layer exhibiting a system of morphologically differentiated polygonal and pyriform canals (C). The interior cavities of the shell are filled with bioclasts, coated grains and rounded intraclasts. Caprinids are common in outer to mid-shelf deposits. Late Cretaceous (Cenomanian): Northern Greece. 3 Hippuritidae. Transverse section of the cylindrical lower valve of Hippurites colliciatus Woodward. Hippuritid rudists are characterized by a thick outer compact calcite layer (CL), a thin inner aragonite layer (AL) and two pillars (P) projecting towards the general cavity. Late Cretaceous (Early Campanian): Pontides, northern Turkey. 4 Radiolitidae. Radial section. Note the folded cellular network made of horizontal laminae and vertical elements. Late Cretaceous: Lake Belton, Texas, U.S.A. 5 Radiolitidae. Sauvagesia sp. The primarily aragonitic inner layer is replaced by blocky calcite. The thick, finely cellular calcitic outer layer consists of prismatic cells. Late Cretaceous (Cenomanian): Greece. 6 Radiolitidae. The outer layer of the upper valve consists of compact calcite. Note the degradation of the shell by boring (arrow). These Cretaceous rudist limestones in the Italian and Slovenian part of the Trieste area are famous decorative building stones. Late Cretaceous (Santonian): Duino near Trieste, Italy. –> 1–3, 5: Courtesy of T. Steuber (Bochum)
Plate 88: Cretaceous Rudist Bivalves
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stones. The thin, curved elements exhibit a single-layered or multilayered calcitic prismatic microstructure. These accumulations within distinctly defined beds were originally regarded as algal filaments and therefore termed ‘filaments’. The fossils have been now proved to be protoconchs or juvenile valves of thinshelled bivalves (Peyre 1959; Conti and Monari 1992). Most authors ascribe Jurassic filaments to juvenile shells of posidoniid bivalves (Bositra, Posidonia) and Triassic findings to halobiid bivalves. Interpretations of the life habits of Bositra and related taxa are still controversial and include nektoplanktonic or pseudoplanktonic as well as benthic modes of life. ‘Limestone with filaments’ is common and widely distributed in the Jurassic (ranging from Toarcian to Kimmeridgian; abundant in the Middle Jurassic and early Late Jurassic) of the Tethys (Pl. 87/3; Pl. 113/2; Pl. 137/1). The depositional settings of filament-rich limestones correspond to the standard facies zones 1, 2 and 3. Cretaceous rudist limestones: Rudists are sessile gregarious bivalves known from Late Jurassic (e.g. Diceras limestones) until the end of the Cretaceous, characterized by a lower attached and an upper opercular valve of different sizes. Three morphotypes are distinguished based on life orientation (Skelton and Gili 1991): conical or cylindrical elevators (vertical upward growth; growing alone, or in close associations or in
Fig. 10.42. Monospecific association of rudists (Vaccinites vesicularis Woodward) contributing to the formation of spectacular Late Cretaceous reefs that may represent outstanding reservoir rocks. Campanian: Saiwan, Central Oman. Hammer for scale. From Schumann (1995).
Rudists
dense clusters with mutual cementation), clingers (growing alone, in close associations or in dense clusters with mutual cementation. Attached to other rudists, corals or hardgrounds), and recumbents (horseshoe-like forms with the long axis of both valves lying parallel to the sediment-water interface). There is currently a lively discussion on whether or not rudists formed rigid reef frameworks or were bafflers forming bafflestones. The depositional settings of rudist limestones usually correspond to the standard facies zones 4, 7 and 8. Rudists are common in Cretaceous platform, ramp and reef carbonates exhibiting a biofacies response to the windward/leeward position and shallowing-upwards cycles. Their habitats stretched from protected and open-marine parts of platforms to shelf margins where they produced huge amounts of bioclastic rudist sands. Rudists formed small bouquets, larger clusters or extended thickets, and contributed significantly to the formation of extended reef belts (e.g. Oman; Fig. 10.42). Rudist density reflects differences in sedimentation rates. Elevator rudists were efficient producers of bioclastic sands that accumulated near Late Cretaceous carbonate platform margins. Boundstones and associated bioclastic sand piles form major oil and gas reservoirs along shelf margins (Mexico, Dubai, Abu Dhabi, Oman). Rudist shells have a high intraskeletal porosity. Additional fracturing increases the reservoir potential. Leaching of the aragonitic parts or of the whole shell during freshwater diagenesis can cause moldic porosity (Cestari and Sartorio 1995). Rudist limestones are widely used as decorative building stones all over the world. Basics: Bivalves Carter, J.C. (1980): Guide to bivalve microstructures. – In: Rhoads, D.C., Lutz, R.A. (eds.): Skeletal growth of aquatic organisms. – 645-673, New York (Plenum Press) Carter, J.C. (1990): Evolutionary significance of shell microstructure in the Palaeotaxodonta, Pteromorpha and Isofilibranchia (Bivalvia: Mollusca). – In: Carter, J.C. (ed.): Skeletal biomineralization: patterns, processes and evolutionary trends. – Volume I, 136-199; volume II (atlas), pls. 1-121. New York (Van Nostrand) Cestari, R., Sartorio, D. (1995): Rudists and facies of the Periadriatic domain. – 207 pp., S. Donato Milanese (Agip) Conti, M.A., Monari, S. (1992): Thin-shelled bivalves from the Jurassic Rosso Ammonitico and Calcari a Posidonia formations from the Umbrian-Marechean Apennine (Central Italy). – Palaeopelagos, 2, 193-213 Elorza, G., Garcia-Garmilla, F. (1998): Paleoenvironmental implications and diagenesis of inoceramid shells (Bivalvia) in the mid-Maastrichtian beds of Sopelana, Zumaya and Bidard sections (coast of the Bay of Biscay,
Gastropods
Basque Country). – Palaeogeography, Palaeoclimatology, Palaeoecology, 141, 303-328 Fürsich, F.T., Oschmann, W. (1993): Shell beds as tools in basin analysis: the Jurassic of Kachchh, western India. – Journal of the Geological Society London, 150, 169-185 Gill, E., La Barbera, M. (1998): Hydrodynamic behaviour of long cylindrical rudist shells: ecological consequences. – Geobios, 22, 137-145 Hagdorn, H., Mundlos, R. (1982): Autochthonschille im Oberen Muschelkalk (Mitteltrias) Südwestdeutschlands. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 162, 332-351 Kidwell, S.M. (1991): The stratigraphy of shell concentrations. – In: Allison, P.A., Briggs, D.E.G. (eds.): Taphonomy: releasing the data locked in the fossil record. – Topics in Geobiology, 9, 211-290, New York (Plenum) Kidwell, S.M., Tothfus, D.A., Best, M.M.R. (2001): Sensitivity of taphonomic signatures to sample size, sieve size, damage scoring systems, and target taxa. – Palaios, 16, 26-52 Lee, C.W. (1983): Bivalve mounds and reefs of the Central High Atlas, Morocco. – Palaeogeography, Palaeoclimatology, Palaeoecology, 43, 154-168 McLeod, K.G., Ward, P.D. (1990): Extinction pattern of Inoceramus (Bivalvia) based on shell fragment biostratigraphy. – In: Sharpton, V.L., Ward, P.D. (eds.): Global catastrophs in earth history. An interdisciplinary conference on impacts, volcanism and mass mortality. – Geological Society of America, Special Paper, 247, 509-518 Monaco, P. (2000): Biological and physical agents of shell concentrations of Lithiotis facies enhanced by microstratigraphy and taphonomy, Early Jurassic, Trento Area (Northern Italy). – GeoResearch Forum, 6, 473-486 Moore, R.C. (ed., 1969): Mollusca 6 (Bivalvia). Volumes 1 and 2. – Treatise on Invertebrate Paleontology, part N, 952 pp., 613 pp., Lawrence (Geological Society of America) Schumann, D. (1995): Upper Cretaceous rudists and stromatoporoid associations of Central Oman (Arabian Peninsula). – Facies, 32, 180-202 Stenzel, H.B. (ed., 1971): Mollusca 6, Bivalvia (Oysters). Volume 3 – Treatise on Invertebrate Paleontology, part N. – 275 pp., Lawrence (Geological Society of America) Further reading: K124
10.2.4.3 Gastropods Gastropods are abundant constituents of limestones (Fig. 10.43). They are known throughout the Phanerozoic, but are most abundant in the Mesozoic and Cenozoic. Morphology: Most gastropods have an external univalved, usually unchambered shell that is usually coiled. Coiling may be either in a single plane or a helical spiral about an imaginary axis. The whorls are tightly wrapped about the preceding whorl to form a central pillar (columella, Pl. 89/5), or the whorls are closely wrapped to form a central cavity. The exterior of the shells may be smooth or ornamented with ribs, nodes and spines.
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Skeletal mineralogy and microstructure: The calcareous shell (ostracum) is covered by an organic layer (periostracum). Most gastropod shells consist completely of aragonite; a few families have a mixed mineralogy, with an outer layer of calcite and an inner layer of aragonite. Conchiolin forms a thin external cover over the shell (periostracum). The whorls are composed of one to several layers of calcium carbonate, which are distinguished by differences in mineral composition and microstructure. The most abundant microstructure in aragonitic layers is crossed-lamellar, characterized by layers of crystals oriented in different directions. Other common microstructures are the normal prismatic found in calcitic layers (fine calcite prisms oriented normal or inclined to the outer surface), complex crossed-lamellar, and nacreous (aragonite crystals arranged in thin sheets and separated by sheets of organic material; e.g. some Archaeogastropoda). Photographs of gastropod microstructures can be found in Majewske (1974). Small fragments of gastropods are distinguished from bivalves by the existence of multiple layers of well-developed crossed-lamellar microstructures, small composite prisms, and nacreous layers associated with layers exhibiting dissimilar microstructures. Preservation: Most gastropods are preserved as molds and casts (Pl. 89/8). Shells composed of aragonite are subject to two main fates: they are either dissolved to form complete or partial molds that may be subsequently filled with cement, or they may be transformed during recrystallization to Low-Mg calcite by the simultaneous volume per volume dissolution of aragonite and precipitation of calcite along intervening solution films. In recrystallized shells some trace of the original microstructures can be retained (Pl. 89/4, 7). Students may confuse small gastropods with foraminifera but the tests of the latter consisting of calcite are rarely preserved as casts. Classification: Three subclasses are differentiated based on the respirative system comprising the Prosobranchia, with more than half of all gastropods, the Archaeogastropoda, Mesogastropoda and Neogastropoda (Cambrian to Holocene), the highly diverse Opisthobranchia (Mississippian to Holocene) and the terrestrial Pulmonata (Pennsylvanian to Holocene). The Opisthobranchia and Pulmonata have aragonitic shells. Most calcite-bearing shells belong to the Prosobranchia. Ecology: Gastropods live in marine, freshwater and terrestrial environments. Most gastropods are mobile benthic organisms, but some have become adapted to sedentary life and to pelagic life (pteropods).
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Plate 89
Gastropods
Gastropods
The gastropod shell corresponds to a conical tube, closed at the pointed end and open at the wide end (apertura). Whorls are wrapped about the preceding whorls forming a central pillar (columella, –> 5) or an axial cavity. The exterior of the shells is smooth or variously ornamented (–> 5). Many gastropod shells are infilled with sediment. Infilling may occur (a) within the environment of the deposition of the shells (as shown by the sediment type within the shell and in the matrix; –> 1, 2, 5), or (b) prior to the deposition of the shells (different sediment in the shell and the matrix; –> 3, 4). Some small, encrusting tubelike fossils, formerly assigned to ‘vermetid gastropods’, are now included within a separate mollusk group - the Microconchida (–> 9, 10). 1 Gastropod wackestone. Longitudinal section (LS) parallel to the axis of coiling, transverse section through the upper whorls (TS), and oblique sections (OS). Black coating of the shells corresponds to the exceptionally well-preserved external organic cover (black arrow). Note different preservation of originally calcitic outer part and the primarily aragonitic and now recrystallized part (white arrow) of the mollusk shell. The sample is an example of a ‘loosely packed, bimodal sorted‚ biofabric’, following the classification of Kidwell and Holland (1991). The fabric is characterized by matrix-support, closely spaced coarse shells and large bioclasts associated with finer bioclasts. SMF 8. Early Triassic (Scythian, Werfen Formation): Gartnerkofel, Carnic Alps, Austria. 2 Preservation. High-spired gastropod. Note double-layered shell construction (white arrow) and displacement of the shell along microfractures (black arrow). Longitudinal section. Late Triassic: Northern Alps, Austria. 3 Preservation. Longitudinal section of a high-spired gastropod. Outlines of the shell are well preserved due to the enclosure within an oncoid. The facies corresponds to a peri-reef environment with some terrigenous influx. Note the angular quartz grains within the gastropod infilling. Late Tertiary (Sarmatian): Prut Valley, Simleu Basin, Romania. 4 Preservation. Oblique section. Note the still preserved growth lines (arrow). Silurian (Ludlowian): Hobuggen, Gotland Island, Sweden. 5 Shell exterior. Longitudinal section. The shell is coiled tightly about its axis, forming a solid central pillar (columella, CO). Note outline of ornamentation (arrow) consisting of strong spiral ribs. Jurassic: Northern margin of the Alps near Salzburg, Austria. 6 Actaeonellid gastropod. Cross section. Actaeonellids are common in brackish-influenced Cretaceous sediments. Early Cretaceous: Subsurface, southern Bavaria, Germany. 7 Low-spired shell. Note entire holostome apertura (A). Arrow points to relicts of prismatic microstructure. Note the geopetal structure. Late Jurassic: Northern margin of the Alps near Salzburg, Austria. 8 Diagnostic criteria. Typical appearance of a gastropod section perpendicular to the axis of coiling characterized by a ‘projection’ of the wall into the central cavity. Cross sections of high-spired gastropod shells have been used to differentiate Tertiary gastropod genera (Strausz 1959). The original wall structure was obliterated during inversion to calcite, but the diagnostic morphological criteria (spire) are still recognizable. The mold is filled by drusy calcite cement. The sediment is pelmicrite or peloid packstone, respectively. Lagoonal facies (see Pl. 117/2). Early Tertiary (Eocene): Monte Bolca,Verona, Italy. 9-10 ‘Microconchids’, formerly considered to be ‘vermetid gastropods’ or spirorbid worms, are tiny coiled tubes with differentiated walls. Erect, helicoidal, gregarious forms are frame builders in peritidal schizohaline lagoonal settings (Burchette and Riding 1977). Prostrate, discoidal forms are encrusters, commonly associated with porostromate algae, forming oncoids in high-energy settings. Microconchus is now assigned to a separate group of mollusks (Microconchida). The group is known from Silurian to the This figure shows a gastropod limestone composed of densely Jurassic (Dreesen and Jux 1995), but more commonly packed shells still exhibiting the original layered microstructure found in Devonian and Early Carboniferous limepreserved as aragonite. Late Triassic Zlambach beds, a near-reef stones. Late Jurassic: Argentina. facies, Austrian Northern Calcareous Alps. Scale is 5 mm.
Plate 89: Gastropods
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Basics: Gastropods
Fig. 10.43. Oolitic gastropod-bivalve grainstone. Gastropods are represented by Holopella gracilior Schauroth. The skeletal grains have thin oolitic coatings exhibiting a reddish color. This microfacies is widely distributed in Early Triassic platform carbonates and known from the Alpine-Mediterranean region (see Pl. 89/1) to Iran and China. Early Triassic (Scythian): Subsurface, Precaucasian foredeep, Russia. Scale is 1 mm.
Marine species are predominantly benthic herbivores and detritus feeders, some species are carnivore. Planktonic pteropods and heteropods known since the Early Tertiary contribute to the formation of aragonitic muds in the deep sea. Basics: Gastropods Bandel, K.J. (1990): Shell structure of Gastropoda excluding Archaeogastropoda. – In: Carter, J.G. (ed.): Skeletal biomineralization: patterns, processes and evolutionary trends. Volume I. – 117-134, New York (Van Nostrand) Dreesen, R., Jux, U. (1995): Microconchid buildups from Late Famennian peritidal-lagoonal settings (Evieux Formation, Ourthe Valley, Belgium). – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 198, 107121 Herman, Y. (1978): Pteropods. – In: Haq, B.U., Boersma, A. (eds.): Introduction to marine micropaleontology. – 151-159, New York (Elsevier) Moore, R.C., Pitrat, C.W. (eds., 1960): Mollusca 1 (Mollusca general features, Scaphopoda, Amphineura, Monoplacophora, Gastropoda general features, Archaeogastropoda, mainly Palaeozoic Caenogastropoda and Opisthobranchia). – Treatise on Invertebrate Paleontology, part I, 351 pp., Lawrence (Geological Society of America)
Strausz, L. (1959): Über Gastropoden-Gehäuseschnitte. – Acta Geologica Budapest, 6, 219-229 Weedon, M.J. (1990): Shell structure and affinity of vermiform ‘gastropods’. – Lethaia, 23, 297-309 Further reading: K124
Although this picture is only included to fill this page, it can tell a story about a sculptured gastropod. Please continue! Triassic: Carinthia, Austria.
Cephalopods
10.2.4.4 Cephalopods Cephalopods are mobile marine mollusks that have adapted to a swimming mode of life. Major groups are distinguished by their soft parts and shell morphology: Nautiloidea (Late Cambrian to Holocene) and some related Paleozoic groups, Ammonoidea (Early Devonian to Cretaceous) and Coleoidea (Early Devonian to Holocene) with the extinct belemnites (Carboniferous to Eocene but most abundant in the Jurassic and Cretaceous where the group is used as index fossils). Morphology: The skeleton of nautiloids and ammonoids consists basically of a straight, curved, spiral or coiled calcareous cone. The cone is chambered by curved, transverse partitions (septa) that unite the inner shell walls. The lines of junction of the septa and the shell walls (sutures) are simple, circular or undulated in nautiloids and related groups, but fluted, crenulated and complex in ammonoids. The first spherical or ovoid chamber is called protoconch, the remaining chambers of the shell form the phragmocone, and the large unchambered space between the last septum and the apertura is the body chamber. The chambers of the phragmocone communicated via a tube (sipho) passing through a perforation in the septa and positioned in the center or at the margin of the shell. The belemnites are characterized by an internal skeleton consisting of the chambered phragmocone, which was partly contained within a cavity at one end of a solid calcitic guard (rostrum). The latter is commonly the only preserved part. Shell structure, diagenesis and preservation: The shells of nautiloids and ammonoids were aragonitic. The wall is layered, exhibiting an outer prismatic layer and inner nacreous and prismatic layers. The shells are frequently recrystallized or leached and subsequently infilled by calcite cement or sediment. Aptychi of ammonoids and the rostra of belemnites consist of calcite (Pl. 90/4, 6). The diagenesis of cephalopod shells has been studied by Dullo and Bandel (1988): Preburial diagenesis is characterized by the rapid decomposition of the organic tissue and the settlement of endolithic organisms, which produces a highly porous structure or micritic envelopes. The decay of organic shell material (e.g. of the siphuncular tube) leads to accessibility for sediment infilling. Aragonitic cements grow syntaxially upon aragonitic microstructural parts of the shell. Burial diagenesis comprises the dissolution of aragonite and the loss of the mineralized shell, often long before the sediment is compacted. If the intercrystalline organic tissue is
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still preserved after dissolution of the aragonite shell crystals, compaction may result in the flattening of cephalopod shells (Pl. 90/3) or replacement of the aragonite biocrystals by phosphate, silica or pyrite. Finally, the organic shell may disappear, leaving a mold that can later be infilled by sediment of cements (Pl. 90/5). In a meteoric environment, the aragonitic shells of cephalopods are greatly affected by dissolution. Leached shells with High-Mg calcite micrite envelopes create molds; the molds are cemented by Low-Mg calcite crystals coarsening towards the center of the cavity or by blocky calcite. Cephalopods in thin sections: Fossil cephalopods have coiled or straight chambered external shells (ammonoids, nautiloids), or internal shells (belemnoids). Coiled forms are usually planispiral forming thin-section patterns comparable to those in planispirally coiled gastropods (Pl. 89/4). However, cephalopod shells are distinguished by having involute coiling (last spiral covers almost all previous ones). In thin sections ammonites are represented by random sections through fragments (Pl. 90) and sometimes abundant sections of juvenile ammonites. Due to the primary aragonitic mineralogy larger shells are usually recrystallized or completely dissolved. Juvenile ammonites i.e. embryonal or larval stages are more common, particularly in Devonian and Mesozoic open-marine wackestones. Sections of these very small fossils (diameter of the shell about 0.5 mm) are characterized by a large globular protoconch followed by one, rarely two whorls subdivided by thin septa. Sections lacking septa, can hardly be distinguished from those of trochospiral dwarf gastropods. The occurrences of abundant juvenile ammonites can indicate the rapid death of plankton. Cephalopod limestones Limestones characterized by abundant ammonoid and nautiloid shells are common in the Paleozoic and Mesozoic (Pl. 90/1, 2). The limestones originated on pelagic platforms, deep shelves or ridges of various origin (submarine volcanoes, drowned reefs, structural highs), platform slopes and basins. Estimates of depositional depths of Paleozoic and Mesozoic open-marine ammonoid limestones are controversial and range between some tens and several hundreds of meters to several thousands of meters. There is good evidence, that the depths of platforms on which Devonian cephalopod limestones in Europe and North Africa accumulated reached an order of several tens to about one hundred meters (Wendt and Aigner 1985). Commonly these carbonates were formed during times of reduced sedi-
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mentation rates and periods of non-deposition. Silurian cephalopod limestones (‘Orthoceras limestones’) accumulated on paleo-highs below wave base at depths of several tens to several hundreds of meters (Ferretti and Kriz 1996). Paleozoic cephalopod limestones are bioclastic wackestones characterized by the predominance of pelagic organisms including ammonoids and nautiloids, dacryoconarids, and thin-shelled ostracods. Mesozoic cephalopod wackestones exhibit ammonoids, thinshelled pelagic bivalves and ostracods as well as radiolarians.
Cephalopods
Aptychus limestones: Basinal sequences consisting of fine-grained pelagic limestones with aptychi are common in Late Jurassic and Early Cretaceous sediments of the Alpine-Mediterranean region. Aptychi are present throughout the sequence although not present in all beds. The fossils are concentrated at the base of some beds in striking shell pavements (Zapfe 1963), indicating the existence of bottom currents, or they occur irregularly distributed as debris (Fig. 10.44). The carbonates originated from the deposition of nannoplankton at depths of several hundreds of meters and more (see Fig. 4.4). The depositional setting of the Aptychus
Plate 90 Cephalopods The plate demonstrates the appearance and common preservation of ammonoid shells (–> 1–3, 5), aptychi (–> 4, 6), nautiloid shells (–> 8) and belemnites (–> 7) as seen in thin sections of Paleozoic and Mesozoic limestones. 1 Mesozoic cephalopod limestone. The sample shows various sections through internal molds of ammonoid shells embedded within micrite (black): (1) Median section through the early whorls, (2) transverse section to the plane of the planispirally coiled shell, (3) oblique section, (4) relicts of the shell. Earlier chambers of the shell are filled with calcite cement, later chambers with micritic sediment. Note that the sedimentary infilling of the chambers with pelagic micrite is the same as the enclosing matrix. This indicates infilling of the shells on the substrate surface before burial (Ferdinandez-Lopez and Melendez 1994). The diagnostically important lines of junction of septa and shell walls (suture, S) are still preserved. The originally aragonitic walls and septa have recrystallized to a calcite mosaic. Basinal facies. Late Triassic (Norian, Hallstatt limestone): Northern Alps, Austria. 2 Paleozoic cephalopod limestone. Bedding-parallel surface of a cephalopod coquina consisting of ammonoids (goniatites: Tornoceras, T) and orthocone nautiloids (ON). Note the preservation of the inner prismatic microstructure of the shell (center). The sample characterizes a ‘condensed shell concentration’ defined by high-diversity skeletal elements carrying different taphonomic signatures (pristine shells as well as abraded shells, various stages of shell solution). Open-marine pelagic platform. Late Devonian (Early Famennian): Tafilalt, eastern Anti-Atlas, Morocco. 3 Ammonite preservation. Hardground with molds of ammonoid shells. The sections are transverse to the plane of coiling. Early Jurassic: Murcia region, southern Spain. 4 Aptychi. These Low-Mg calcitic plates are found associated with ammonites and interpreted as parts of the jaws or an opercular structure. Normally the plates occur as pairs and have an ornamented outer surface. Note the cellular-tubular microstructure (white arrow) and the ornamentation of the shell surface (black arrow). Classification of aptychi is based on the kind of ornamentation formed by fine lines, ribs (this sample), tubercles or spines. The matrix of the radiolarian mudstone and the aptychi are selectively silicified (white). Late Jurassic: Monte Generoso, Ticino, Switzerland. 5 Ammonite preservation. The shell has been infilled with dark pelagic micrite (PC) and after breakage and dissolution of the shell, was later embedded in forereef carbonates (FC) with fine-grained bioclastic sediment. The septa were replaced by coarse granular calcite cements followed by the burial filling of the chambers with blocky calcite. O: Ophiurid element. The sample comes from a locality showing the interfingering between basinal sediments and forereef facies. Late Triassic (Norian): Hoher Göll, Northern Alps, Bavaria, Germany. 6 Aptychus. The thin section shows the characteristic ornamentation. White spheres are radiolarians. This microfacies is common in limestones deposited during pre- or post-radiolarite phases of deep-sea sedimentation. Same locality as –> 4. 7 Belemnite guards. Transverse section (TS) and longitudinal section (LS) of the Low-Mg calcite guards (rostrum). In transverse section the rostrum shows a fibrous structure with tiny calcite prisms radiating from an eccentric point. Rings concentric about this point indicate growth stages. Under crossed nicols an extinction cross can be seen resulting from the radial orientation of the individual prisms making up the rostrum. The microstructure of belemnite shells seen in SEM can be used to distinguish belemnite fragments from other mollusk remains (Dauphin 1986). Late Jurassic (Late Oxfordian): High Atlas Mountains, Morocco. 8 Nautiloid shell. Longitudinal section of a straight orthocone shell embedded within a bioclastic wackestone matrix. Note preservation similar to the ammonite shell shown in –> 5. Open-marine deep-water ramp. Early Devonian (Emsian): Erfoud, Morocco. –> 2: Wendt and Belka 1991; 6: Kiessling 1996
Plate 90: Cephalopods
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Aptychus Limestones
Fig. 10.44. Aptychus limestone. The thin section displays numerous fragments of aptychi within a fine-grained matrix. The fragments appear curved due to the slightly domal character of the plates. The exterior ornamentation (white arrows; T: tangential section of ribbed shell surface) is shown by uniformly oriented ribs. Small dark dots in the shells (black arrows) are caused by the porous microstructure. The concentration of the shells in particular areas and their irregular distribution within the sample are characteristics of strong burrowing. Late Jurassic: Karawanken Mountains, Austria. Scale is 1 mm.
limestones corresponds to the standard facies zones 1 and 2 (see Chap.14). ‘Belemnite battlefields’: Accumulations of belemnite rostra (Pl. 90/7) on bedding planes (Fig. 10.45) are important facies criteria indicating current directions and the type of near-bottom water energy in shelf settings (Urlichs 1971; Doyle and Macdonald 1993), reduced and condensed sedimentation and lag deposition. A common feature of belemnite battlefields is the monospecific composition and the predominance of adult individuals. The most likely explanation of the battlefield is post-spawning mortality in shelf areas followed by current transport. Basics: Cephalopods Blendinger, W. (1991): Upper Triassic (Norian) cephalopod limestones of the Hallstatt type, Oman. – Sedimentology, 38, 223-242 Doyle, P., Macdonald,D.I.M. (1997): Belemnite battlefields. – Lethaia, 26, 65-80 Dullo, W.C., Bandel, K. (1988): Diagenesis in molluscan shells: a case study from cephalopods. – In: Wiedmann, J., Kullmann, J. (eds.): Cephalopods. Present and past. – 719-729, Stuttgart (Schweizerbart) Fernandez-Lopez, S., Melendez, G. (1994): Abrasion surfaces
on internal moulds of ammonites as palaeobathymetric indicators. – Palaeogeography, Palaeoclimatology, Palaeoecology, 110, 29-42 Ferretti, A. (1989): Microbiofacies and constituent analysis of Upper Silurian-Lower Devonian limestones from southwestern Sardinia. – Bolletino della Società Geologica Italiana, 28, 87-100 Kaesler, R.L. (ed., 1996): Mollusca 4. Cretaceous Ammonoidea. – Treatise on Invertebrate Paleontology, part L (revised), 362 pp., Lawrence (Geological Society of America) Moore, R.C. (ed., 1957): Mollusca 4 (Ammonoidea). – Treatise on Invertebrate Paleontology, part L, 490 pp., Lawrence (Geological Society of America) Moore, R.C. (ed., 1964): Mollusca 3 (Cephalopoda general features, Endoceratoidea, Actinoceratoidea, Nautiloidea, Bactritoidea). – Treatise on Invertebrate Paleontology, part L, 519 pp., Lawrence (Geological Society of America) Schlager, W. (1974): Preservation of cephalopod skeletons and carbonate dissolution on ancient Tethyan sea floors. – International Association of Sedimentologists, Special Publication, 1, 49-70 Tucker, M. (1974): Sedimentology of Palaeozoic pelagic limestones: the Devonian Griotte (Southern France) and Cephalopodenkalk (Germany). – International Association of Sedimentologists, Special Publication, 1, 71-92 Urlichs, M. (1971): Alter und Genese des Belemnitenschlachtfelds im Toarcian von Franken. – Geologische Blätter für Nordost-Bayern, 21, 65-83 Wendt, J., Aigner, T. (1985): Facies patterns and depositional
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Fig. 10.45. Accumulations of belemnite rostra (predominantly Dactyloteuthis) on a marly-calcareous bedding surface accurately named ‘belemnite battlefield’. The bed overlies the black organic-rich Posidonia shales, interpreted as deposition in dysaerobic-anaerobic basins with fluctuating oxygen conditions at the sediment-water interface. Early Jurassic (boundary Early/Late Toarcian): Mistelgau, southwest of Bayreuth, northern Franconian Alb, Germany. Large specimens up to 10 cm long. Courtesy of C. Schulbert (Erlangen). environments of Palaeozoic cephalopod limestones. – Sedimentary Geology, 44, 263-300 Zapfe, H. (1963): Aptychenlumachellen. – Annalen des Naturhistorischen Museums Wien, 66, 261-266 Further reading: K124
10.2.4.5 Tentaculitids and Other Conical Shells (Cricoconarida; Hyolithida) Tentaculitids are extinct millimeter- to a few centimeter-sized exclusively Paleozoic fossils of uncertain affinities but now considered as a class of the Mollusca. They are characterized by slender conical calcitic shells occurring in rock-building abundance in Ordovician to Late Devonian open-marine and basinal carbonates. Tentaculitids and related fossils are common constituents of Devonian pelagic and hemipelagic limestones. Morphology and classification: The conical shells have a pointed apex, some groups are chambered. Classification is based on the shell sculpture, type and angle of the apex and shell microstructure. The class Cricoconarida is defined by having conical calcareous shells with ringlike exterior sculptures. The exterior of the shell is characterized by periodic concentric thicken-
ings (annular rings) and fine concentric or longitudinal striations. Longitudinal sections show narrow V-shaped figures (apical angle up to 20°); in cross sections circular, elliptical or polyhedral figures can be observed. Excellent thin-section microphotographs can be found in Farsan (1994). Tentaculites may be confused with brachiopod spines (see Pl. 86/6). The class includes two orders - the Tentaculitida and the Dacryconarida. The tentaculitids (Ordovician to Late Devonian II) have concentric ring-like thickenings at the exterior surface but a smooth interior wall surface. The shells (length commonly 15-30 mm) are pointed, the initial part is chambered. The microstructure is finely laminated and can have fine perforations near the apertura. The dacryconarids (Ordovician to Late Devonian II), comprising the nowakiids and styliolinids, are characterized by small shells, apical angles larger than in tentaculitids, a bubble-shaped and unchambered initial shell section, and smooth or longitudinally striated, shell surfaces. This group comprises the styliolinids and nowakiids, the latter characterized by ringed exterior and interior surfaces. The styliolinids have smooth exterior and interior shell surfaces. Styliolina differs from Tentacu-
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Tentaculitids
Fig. 10.46. Pelagic limestones with cricoconarid shells. A: Styliolinid grainstone. Telescoping in some shells is caused by currents (arrows). Late Devonian (Frasnian): Eastern AntiAtlas, Morocco. After Wendt and Belka (1991). B: Tentaculitid wackestone with longitudinal and cross sections of Nowakia characterized by ringed exterior and interior surfaces. Middle Devonian (Eifelian): Prague Basin, Czech Republic. Scale is 1 mm.
lites in possessing a smooth, unchambered, imperforate shell without concentric thickenings. The shell usually terminates in a small bulbous extremity. The microstructure is clear homogeneous calcite. Transverse sections show an extinction cross under crossed nicols. The tentaculitids are valuable index fossils used in regional zonations (particularly Devonian). Ecology and paleoenvironment: The styliolinids and dacryoconarids were pelagic (mainly planktonic) inhabitants of the upper reaches of the oceans contributing to the formation of pelagic carbonates in basins and deep ramp settings (Fig. 10.46). The tentaculitids sensu strictu lived probably as nektobenthos; they are abundant in slightly argillaceous, sometimes nodular packstones and wackestones (e.g. ‘Flaserkalk’ of the German Devonian cephalopod limestone) as well as in shales formed on deep shelves, submarine highs (swells) Plate 91
Devonian Tentaculitids and Styliolinids
Small conical calcitic shells including the tentaculitids, styliolinids and other morphologically similar fossils are common and rock-building constituents of Paleozoic, particularly Silurian and Devonian limestones formed in open-marine environments. The fossils are interpreted as presumably planktonic organisms allowing regional and global biozonations e.g. of the upper Late Devonian. Diagnostic criteria are the shell sculpture (–> 2), type and angle of the apex (–> 3, 5) and shell microstructure (–> 7). 1 Bioclastic wackestone with tentaculitids (T) and trilobites (TR). The sample is from bedded carbonates deposited on a deep platform. Early Devonian (Emsian): Erfoud, Anti-Atlas, Morocco. 2 Longitudinal section (LS) of Nowakia (N) with characteristic ribs on the exterior and interior shell surfaces, and transverse section of Styliolina (S). Note stacked cone-in-cone cross section of Styliolina (center bottom). The micritic matrix is rich in fine-grained bioclastic debris (white spots). Middle Devonian (Eifelian): Prague Basin, Czech Republic. 3 Tentaculitid packstone. Variously sized circular and polyhedral cross sections and V-shaped longitudinal sections. Note ‘cone-in-cone structure’ of some shells (arrows) indicating current activity (Larsson 1979) and the delicate external transversal ornamentation of the shells typical for tentaculitids. SMF 3-TENT. Early Devonian: Frankenwald, Germany. 4 Tentaculites. Note shells stacked one within the other (arrow) and the current-controlled unidirectional orientation of some shells. Same locality as –> 3. 5 Styliolina microcoquina packstone deposited after current transport on a pelagic swell. Note the late diagenetic calcite cement covering the shells (arrows). Styliolinid coquinas occur in pelagic to hemipelagic settings and are known from the Lochkovian to the Late Frasnian. Early Devonian (Emsian): Erfoud, Anti-Atlas, southern Morocco. 6 Pelagic styliolinid wackestone. Styliolinids (arrow) range from the Early Devonian to the Frasnian. The group is characterized by the absence of peripheral ornamentation. Limestones composed entirely of styliolinids are believed to be ancient analogues of modern pteropod-rich carbonates. E: Echinoderm fragment; O: ostracod. White curved shells are trilobite fragments. Middle Devonian (Givetian): Menorca Island, Spain. 7 Preservation of styliolinid shells: The original microstructure of the shell walls is completely replaced by a fibrous calcite displaying distinct curved twin-lamellae (arrows), probably indicating low-grade metamorphism (Tucker and Kendall 1973). This preservation is common in many styliolinid limestones. The fine-grained matrix is microspar. Middle Devonian pebble within Lower Carboniferous flysch deposits: Betic Cordillera, southern Spain.
Plate 91: Tentaculitids and Styliolinids
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540
Hyolithids
Fig. 10.47. Hyolithid shells. A: cross section. B: section along the longest axis. Both photographs show telescoping of various shells explained by wave transport in nearcoast shallow setting. Uppermost Early Cambrian: Northeastern shore of the Dead Sea, Jordan. Scale is 1 mm. After Shinaq and Bandel (1992).
or in basins. Tentaculitids can reflect paleocurrent bottom water patterns. Criteria used are frequency distributions on bedding surfaces, or graded bedding, imbrication and abundance of calcisiltite components (tentaculites) in the lower parts of turbidite beds (Hladil et al. 1996). Very small-sized cones will roll downcurrent, larger cones will point with the apex into the currents. Presupposing a planktonic life habitat, mass accumulations of dacryoconarids (Pl. 91/3) and styliolinids in packstones characterize the Standard Microfacies Type 3 common in open-marine deep shelf and basinal settings. Hyolithids Other conical shells restricted to the Paleozoic and most abundant in the Cambrian are included within the Hyolithida (Fig. 10.47). The group is considered to be a class of the Mollusca or related to annelid worms. The probably aragonitic shell is cone- or pyramidshaped and was closed by an operculum. Sizes range between a few millimeters and about 200 mm. The microstructure corresponds to the crossed-lamellar type. Hyolithids can be abundant in Cambrian shelf carbonates. Basics: Tentaculitids, styliolinids and hyolithids Farsan, N.M. (1994): Tentaculiten: Ontogenese, Systematik, Phylogenese, Biostratonomie und Morphologie. – Abhandlungen der Senckenbergischen naturforschenden Gesellschaft, 547, 1-128 Fisher, D.W. (1962): Small conoidal shells of uncertain affinities - calyptoptomatids. – Treatise on Invertebrate Paleontology, part W, 116-130, Lawrence (Geological Society of America) Hladil, J., Cechan, P., Gabasova, A., Taborsky, Z., Hladikova,
J. (1996): Sedimentology and orientation of tentaculite shells in turbidite lime mudstone to packstone: Lower Devonian, Barrandian, Bohemia. – Journal of Sedimentary Research, 66, 888-899 Larsson, K. (1979): Silurian tentaculitids from Gotland and Scania. – Fossils and Strata, 11, 1-180 Poijeta, J. (1987): Phylum Hyolitha. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 416-444, Palo Alto (Blackwell) Shinaq, R., Bandel, K. (1992): Microfacies of Cambrian limestones in Jordan. – Facies, 27, 263-284 Tucker, M., Kendall, A.C. (1973): The diagenesis and lowgrade metamorphism of Devonian styliolinid-rich pelagic carbonates from West Germany: possible analogues of recent pteropod ooze. – Journal of Sedimentary Petrology, 43, 672-683 Further reading: K124
10.2.4.6 Serpulids and Other Worm Tubes Fossils referred to as worm tubes are known from throughout the Phanerozoic and common in thin sections of marine and marginal-marine limestones. These tubes may be calcareous (Pl. 92/1-8) or agglutinated (Pl. 92/9, 10). Most calcareous tubes are assigned to annelid Polychaeta. Annelids are known since the Ordovician. Morphology: Calcareous worm shells are straight, curved or spirally coiled unpartioned tubes, a few millimeters in maximum diameter and up to several centimeters in maximum length. The tubes consist of Mgcalcite or aragonite or both. The skeletal mineralogy is related to mean water temperature. The tubes can form rock-building gregarious associations (Pl. 92/3, 4) as
Worm Tubes
541
well as serpulid ‘colonies’ which attribute in the construction of ‘reef’ structures. Many serpulids are characterized by a microstructure consisting of two layers which in longitudinal sections show an outer cone-incone pattern and a clear or laminated inner layer (Pl. 92/6, 7). In tangential sections the shell layers appear as concentric rings (Pl. 92/8). Spirorbid calcareous worms are fixed with supporting structures on hard or soft substrates (Pl. 92/5). Paleoenvironmental significance: Serpulids occur in very different environments (including shallow and rocky coasts, sea grass areas, hard and firm sea bottoms, deep-marine habitats). They can live in warm and temperate seas, and are adapted to various salinities. The fossil record comprises scattered bioclasts, rockbuilding accumulations of serpulid tubes or autochthonous reef-like structures (Fig. 10.48). Modern serpulid ‘reefs’ are represented by banks deposited in instable marine habitats (commonly nearcoastal environments; often mixo- or hypersaline) on shallow parts of continental shelves, and by meter-sized framework reefs at intertidal and subtidal depths, down to 30 m depths, mainly in sheltered coastal lagoonal conditions (Ten Hove and Van Den Hurk 1993; Fornos et al. 1997). Buildups dominated by masses of serpulid tubes associated with coralline red algae are known from lagoonal environments of the Bermudas (Pl. 92/ 1), but also occur in cold-water environments of Antarctica. Ancient serpulid reefs are known from the Early Carboniferous of England (Wright and Wright 1981), from the Early Cretaceous of Argentina (Palma 1992) and of central Europe (Berriasian: hypersaline ‘Serpulit’ beds) as well as from Miocene lagoonal or brackishwater environments (Pl. 92/3). Associations of serpulids and microbial carbonates are common. Examples are Late Triassic serpulid/microbial reefs consisting of serpulid aggregates, microbial crusts and abundant synsedimentary carbonate cements (Flügel 2002). In the Late Tertiary, serpulids in association with microbes formed small-scaled reefs in near-coastal environments (Jasionowski and Wysocka 1997). Agglutinated worm tubes are constructed by siliciclastic sand and silt grains, carbonate and organic material (Pl. 92/9, 10). These tubes have been known since the Early Carboniferous, often associated with reef environments. Different names are in use for designating annelid agglutinated mucous sheats tending to early cementation and reworking from the enclosing sediment (Terebella, Thartharella, Prethocoprolithus; see Delvolvé et al. 1994).
Fig. 10.48. Allochthonous and autochthonous serpulid records. A: Scattered and broken serpulid tubes redeposited in quartz sand. Early Cretaceous: Subsurface, Endelhausen 1, southern Bavaria. B: Rock-building concentration of serpulid tubes. Middle Jurassic: Würgau, Northern Franconian Alb, Germany. C: Serpulid ‘colony’ contributing to the formation of centimeter- to meter-sized reefs in near-coast, often brackish settings. Late Tertiary (Miocene): Poland. Scale is 2 mm. Basics: Serpulid worms Brönnimann, P., Zaninetti, L. (1972): On the occurrence of the serpulid Spirorbis Daudin, 1800 (Annelida, Polychaeta, Sedentaria) in thin sections of Triassic rocks of Europe and Iran. – Rivista Italiana di Paleontologia e Stratigrafia, 78, 67-90 Götz, G. (1931): Bau und Biologie fossiler Serpuliden. – Neues Jahrbuch für Mineralogie, Geologie und Paläontologie, B, 66, 385-438
Serpulids
542
Misik, M., Sotak, J., Ziegler, V. (1999): Serpulid worms Mercierella Fauvel, Durandella Dragastan and Carpathiella nov. gen., from the Jurassic, Cretaceous and Paleogene of the Western Carpathians. – Geologica Carpathica, 50, 305-332 Moore, R.C. (1962): Miscellanea (Conodonts, conoidal shells of uncertain affinities, worms, trace fossils, problematica). – Treatise on Invertebrate Paleontology, part W, 259 pp., Lawrence (Geological Society of America) Railsback, L.B. (1993): Original mineralogy of Carboniferous worm tubes: evidence for changing marine chemistry and biomineralization. – Geology, 21, 703-706 Schmidt, W.J. (1951): Die Unterscheidung der Röhren von Scaphopoda, Vermetidae und Serpulidae mittels mikroskopischer Methoden. – Mikroskopie, 6, 373-381 Ten Hove, H.A., Van Den Hurk, A.M. (1993): A review of Recent and fossil serpulid ‘reefs’, actuopalaeontology and the ‘Upper Malm’ serpulid limestones in NW Germany. – Geol. Mijnbouw, 72, 23-72 Weedon, M.J. (1994): Tube microstructure of Recent and Jurassic serpulid polychaetes and the question of Paleozoic ‘spirorbids’. – Acta Palaeont. Polonica, 39, 1-15. Further reading: K125
Plate 92
10.2.4.7 Crustacean Arthropods Arthropods are characterized by segmentation, specialized appendices and a rigid exoskeleton that must be shed or molted if the organisms change size or shape. The skeletons of some arthropod groups also have a good fossil record in carbonate rocks. Three of these groups belong to the superclass Crustacea, a fourth and most important group are the trilobites, regarded as an extinct superclass (see Sect. 10.2.4.8). Crustaceans recorded in limestones include ostracods, cirripedians represented by balanids, and microcoprolites of decapod crabs. Ostracods Ostracods are bivalved mm- to cm-sized crustacean arthropods, known since the Late Cambrian to occur in nearly all types of aquatic environments. They are valuable proxies for paleoenvironmental conditions
Serpulids and Other Worm Tubes
The plate displays calcareous worm tubes (–> 1–8) and worm tubes with walls consisting of agglutinated grains (–> 9, 10). 1
Serpulids and coralline red algae. Modern aragonitic serpulid tubes. Note distinct differences in tube diameter. Arrow points to sessile foraminifera enclosed in crusts made by coralline red algae (black). Reef environment, Bermuda. 2 Temperate setting. Serpulids occurring in modern beachrock, formed in temperate water. Serpulids are common constituents of temperate and cold-water carbonates. Note concentric shell structure (arrows). Black grains are volcaniclastics. Fuerteventura, Canary Island, Spain. 3 Reefs. Serpulids forming small-scale reefs in near-coast settings. Note fibrous cement rim around the annelid tubes. Late Tertiary (Early Sarmatian, Miocene): Prut Valley, Romania. 4 Rock building. Gregarious growth of serpulids in brackish-water environments may result in the formation of serpulid boundstones (serpulite) characterized by monospecific associations (here: Serpula socialis Goldfuss). Late Jurassic: Pinkerdorf, Lower Saxony, Germany. 5 Spirobids. These calcareous worms have distinct supporting structures (SS). The sample comes from a shallow-marine limestone block that was transported into the deeper ocean. During deposition the block was overturned, as indicated by the geopetal infill. Late Jurassic: Northern margin of the Calcareous Alps near Salzburg, Austria. 6 Microstructure: Serpulid shells exhibiting the characteristic two-layered foliated microstructure with the cone-in-cone pattern. Cretaceous: San Pietro, Mussolino, Apennines, Italy. 7 Microstructure. The walls of the serpulid tubes consist of two microstructurally different layers (arrows). Middle Jurassic: Würgau, northern Franconian Alb, Germany. 8 Microstructure. Characteristic cross section of a serpulid shell. Note the concentric, finely laminated structure. Early Cretaceous: Ras al Khaimah, United Arabian Emirates. 9 Agglutinated worm tube. Oblique section. Late Jurassic: Subsurface, Kinsau near Augsburg, southern Germany. 10 Agglutinated worms. Longitudinal (LS) and cross section (CS) of Terebella lapilloides Münster or Prethocoprolithus Elliott. The silt-sized wall material is bound together by microcrystalline carbonate. Note the intensive dark color reflecting the organic content of the ‘wall’. The supporting lining to the burrow is typical of a loose-ground consistency of the sediment. These worms are common in flanking beds of Upper Jurassic sponge reefs and in wackestones of deeper ramp settings. Associations of Terebella and Tubiphytes characterize deeper, low-energy environments and are diagnostic of regionally lowered sedimentation rates (Leinfelder et al. 1993; Schmid 1996; Schlaginweit and Ebli 1999). The burrows are also found as reworked particles in grain- and rudstones (see Pl. 116/1). Late Jurassic: Schammental, Suebian Alb, southern Germany.
Plate 92: Serpulids and Other Worm Tubes
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544
including salinity, oxygenation, substrate and water depths. The group is useful in the biostratigraphic analysis of marine and nonmarine sediments, particularly in freshwater environments where other fossils may be lacking. The calcitic shells may be common in Mesozoic and Cenozoic lacustrine limestones and brackish-water carbonates as well as in marine shallow- and deep-marine limestones all through the Phanerozoic. Morphology: Ostracods are characterized by a bivalved shell (carapace) that is typically ovate or kidney-shaped in outline. The shell is opened by a dorsally located ligament and closed by a set of adductor muscles. The carapace encloses the soft parts and consists of two valves which overlap each other along the ventral margin or at all margins. The valves are commonly equally shaped but of a slightly different size;
Ostracods
they sometimes show a characteristic overlap of one valve by the other. The valves are articulated by teeth and sockets or ridges and grooves along the dorsal hinge line. The carapace is frequently pierced by a series of tiny canals (not obvious in thin sections). Internally many smooth valves show a recurved shelf (duplicature, calcified portion of the inner lamella) along the free margins (best recognized in cross-sectional views). The surface of the valves is smooth (Pl. 93/1) or sculptured (Pl. 93/4) into ridges, ribs, spines or nodes. Fossil ostracods are defined by the shape of the carapace, carapace surface sculpture, type of hinge, placement of pores and muscle insertion scars. Only a few of these criteria can be evaluated from thin sections. Shell structure: The calcified carapace consists of an outer and an inner calcitic lamella. The microstruc-
Plate 93 Crustacean Arthropods Ostracods (–> 1-4) can be recognized in thin sections of limestones by small calcitic shells consisting of more or less equally shaped but slightly differently sized valves (–> 1), sometimes showing finely prismatic microstructures (–> 2) or shell ornamentation (–> 3, 4). Balanids (barnacles) are sessile cirripedians protected by a calcitic multi-plate skeleton consisting of a massive or hollow basal plate (–> 6) and articulated wall plates (–> 6, 7). Balanid shells are often associated with bryozoans and coralline algae (–> 7) as well as benthic foraminifera. Favreinids (–> 5) are fecal pellets of decapod crabs. These microcoprolites can be easily recognized by their sieve-like structure caused by calcite-filled canals that form specific patterns. 1 Ostracod-bivalve wackestone. Ostracods are represented by disarticulated valves (black arrows) and complete tests (white arrows). The percentage of disarticulated carapaces provides a measure of post-mortem modification by taphonomic processes. The bivalve shells (BS) are replaced by coarse calcite. Late Jurassic: Margas de Terenas, Ribadresella, Spain. 2 Ostracod microstructure. The thick shell consists of two valves showing distinct overlap (O). The carapace exhibits a finely prismatic microstructure. Note the contrast between the well-preserved ostracod shell and the strongly recrystallized sedimentary matrix. Silurian: Gotland Island, Sweden. 3 Paleozoic pelagic ostracods. The thin section shows complete and disarticulated valves of highly spinose ostracods belonging to the Entomozoacea (Myodocopida). Entomozoans are common in Devonian basinal sediments and are interpreted as having been planktonic. The sample comes from limestones deposited in an open-marine deep-shelf environment. Late Devonian (Famennian): Menorca Island, Spain. 4 Pelagic entomozoan ostracods. Note the spinose ornamentation (arrows) of the shells. Same locality as –> 3. 5 Favreinids. Fecal pellets of decapod crabs. Poorly sorted grainstone consisting of Favreina (cross sections, CS, and longitudinal sections, LS), tiny peloids and large rounded intraclasts. This microfacies is common in the lagoonal environments of inner ramps affected periodically by storms. Lowermost Cretaceous (Purbeck facies, Berriasian): Subsurface, Kinsau well, Bavaria, Germany. 6 Balanid shell. Cross section of the basal plate. The hollow plate is subdivided by partitions (black arrows). Some wall plates are displayed in the upper part of the cross section. The barnacle and bryozoans (B) are encrusted by coralline red algae (white arrows) forming a macroid. Multitaxon macroids indicate unstable substrates. The sample comes from warm-temperate carbonates (Nebelsick 1989). Late Tertiary (Early Miocene): Eggenburg, Austria. 7 Balanid-bryozoan-coralline algal association. The poorly sorted bioclastic packstone consists of balanids (white arrow pointing to an oblique section showing some wall plates), bryozoans (B) and coralline red algae (black arrow). This association is a common attribute of Cenozoic temperate and cold-water carbonates. The irregularly shaped open vugs (white) are caused by freshwater leaching. Late Tertiary (Early Miocene): Zogelsdorf near Eggenburg, Austria.
Plate 93: Crustacean Arthropods
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546
Balanids
Fig. 10.49. Cross sections of ostracod shells. A: Shell showing the duplicature. The excellent preservation is due to an early diagenetic silicification of the test. Late Carboniferous (Gzhelian): Auernig, Carnic Alps, Austria/Italy. B: Slightly displaced smooth-shelled test exhibiting the finely prismatic microstructure. Early Carboniferous: Poland. C: Sculptured shell of an entomozoan pelagic ostracod. Late Devonian: Menorca Island, Spain.
ture is layered, homogeneous or finely prismatic. The microstructures commonly form only one shell layer, except in some Paleozoic ostracods (e.g. Leperditiidae). Ecology: Most ostracods are found crawling on or burrowing into the sediments at the bottoms of the ocean (Fig. 10.50) or lakes. A few modern ostracods have a pelagic swimming life style.
Distribution: Ostracods appeared in the Cambrian, diversified in the Ordovician and were abundant in the Silurian and Devonian, where they are successfully used for biozonations. Late Devonian and lowermost Carboniferous sediments yield the entomozoan ostracods that can also be recognized in thin sections. New groups also including deep-water ostracods (Fig. 10.50) appeared in the Late Paleozoic, Mesozoic and Cenozoic. Ostracods in thin sections: Ostracods are common in many thin sections due to the repeated shedding that result in accumulations of high numbers of disarticulated valves (Pl. 93/1). Complete shells are small; they rarely exceed a length of 1 mm. Criteria in thin sections include shape, size, shell overlap (Pl. 93/2), exterior ornamentation (Pl. 93/4; Fig. 10.49C) and the prismatic shell microstructure (Pl. 93/2). The duplicature is rarely shown (Fig. 10.49B). In general, these criteria do not allow systematic assignments, except where characteristic shell ornamentation patterns can be recognized.
Fig. 10.50. Deep-marine ostracods associated with benthic foraminifera, mollusk debris and echinoderm fragments exhibiting the characteristic fine net-like structure. The ostracod near the top (infilled with red micrite) displays the duplicature. The large particle near the bottom is a hardground intraclast (see Sect. 4.2.8.1). Early Jurassic: Adnet, south of Salzburg, Austria. Scale is 1 mm.
Balanids Balanids or barnacles (Pl. 93/6, 7) are sessile crustacean arthropods belonging to the Cirripedia. Cirripedians are highly unusual crustaceans that loose many of their typical arthropod features during ontogeny. All adults are sessile. The carapace of the larvae is replaced by a series of calcified plates that form a protective compartment for the body and the appendages. Balanids occur attached to various substrates including rocky coasts, shells, drifting wood or swimming organisms. Attachment is made by a broad basal plate (Pl. 93/6) that is a part of the calcitic test. The shell forms a truncated cone consisting of up to eight calcareous wall plates and other plates that can close the aperture. Modern balanids are abundant in temperate and cold-water environments and occur from intertidal
Favreinids
547
to mid- and high-latitudinal shelf settings. Balanid shells form a large part of ‘bryomol’ and ‘barnamol’ sediments (Sect. 12.2.2.1). Fossil balanids have been known since the Cretaceous and are common constituents of Late Tertiary nontropical carbonates (Kamp et al. 1988; Pl. 106/3).
Favreinids A somewhat curious microfossil characterized by small micritic bodies with numerous spar-filled eyes has been described from Mesozoic shallow-marine limestones as Favreina (Fig. 10.51), which is a microcoprolite. Some thalassinid decapod crabs produce fecal pellets exhibiting characteristic sieve-like patterns of tubules formed by projections from the intestine walls. The shape, arrangement and size of the calcitefilled tubules are used for taxonomic differentiation. Mass accumulations of these fecal peloids, commonly associated with smaller peloids or ooids, contributed significantly to the formation of extended sheats of carbonate sands in the latest Jurassic and earliest Cretaceous restricted lagoons (Pl. 93/5, Pl. 136/ 9). Fossil decapod coprolites assigned to Favreina and similar genera have been reported from Devonian to Tertiary limestones. Most reports deal with Jurassic and Cretaceous favreinids, followed by Triassic findings. Rock-building occurrences are abundant during the Late Jurassic–Early Cretaceous time interval. A bibliography prepared by J. Blau can be found on the Internet. Basics: Crustacean arthropods Bourget, E. (1980): Barnacle shell growth and the relationship to environmental factors. – In: Rhoads, D.C., Lutz, R.A. (eds.): Skeletal growth of aquatic organisms. – Topics in Geobiology, 1, 469-492, New York (Plenum Press) Dall, W., Moriarty, D.J.W. (1983): Functional aspects of nutrition and digestion. – In: Bliss, D.E. (ed.): The biology of Crustacea. Volume 5. Internal anatomy and physiological regulation. – 215-261, New York (Academic Press) Kaesler, R.I. (1987): Superclass Crustacea. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 241-257, Palo Alto (Blackwell) Molinari Paganelli, V., Pichezzi, R., Tilia Zucchari, A. (1980 and 1984): I coprolithi di crostacei. Rassegna bibliografica e annotazioni tassonomiche. Parte I. Genere Favreina. Parte II. Generi Helicerina, Palaxius, Parafavreina e Thoronetia– Bolletino del Servizio geologico d’Italia, 100 (1979), 409-454 and 104 (1983), 309-344 Moore, R.C., Pitrat, C.W. (eds., 1961): Arthropoda 3. Crustacea Ostracoda. – Treatise on Invertebrate Paleontology, part Q, 442 pp., Lawrence (Geological Society of America) Pokorny, V. (1978): Ostracodes. – In: Haq, B.U., Boersma,
Fig. 10.51. Longitudinal (A) and cross section (B) of Favreina, a microcoprolite of decapod crabs. The patterns caused by the arrangement of the calcite-filled canals are used for taxonomic differentiation. Lowermost Cretaceous (Berriasian): Kinsau well, Bavaria, Germany. Scale is 500 µm.
A. (eds.): Introduction to marine micropaleontology. – 109-149, New York (Elsevier) Schweigert, G., Seegis, D., Fels, A., Leinfelder, R. (1997): New internally structured microcoprolites from Germany (Late Triassic/Early Miocene), Southern Spain (Early/ Middle Jurassic) and Portugal (Late Jurassic): Taxonomy, palaeoecology and evolutionary implications. – Paläontologische Zeitschrift, 71, 51-69 Further reading: K127 (ostracods), K139 (favreinids)
10.2.4.8 Trilobites Fragments of trilobites may be common in Early Paleozoic shelf limestones. Trilobites are extinct, exclusively marine arthropods restricted to the Early Cambrian to Late Permian time interval. Most trilobites lived in fairly shallow water and were benthic; a few may have been pelagic, floating in the water column. Benthic forms walked on the sea bottom and fed on detritus. The segmented dorsal exoskeleton is longitudinally differentiated into an anterior head shield (cephalon), the thorax consisting of a few to many segments, and a tail shield (pygidium). Laterally the skeleton is subdivided into a raised axial portion and two lateral portions. The surface of the skeleton is smooth or variously ornamented with tubercles, spines and other elements. Shell structure: The skeleton consists of chitin hardened by impregnation with calcium phosphate and calcite. The microstructure of the finely layered shell is homogeneous prismatic (Pl. 94/ 6). Some trilobite fragments are characterized by fine pores or canals oriented perpendicular to the surface (Pl. 94/1, 3). Trilobite spines may be confused with brachiopod spines, but the latter consist of two calcite layers showing concentric foliated structures in contrast to homogeneous prismatic trilobite spines.
548
Trilobites in thin sections: The most common trilobite fragments seen in thin sections are dissociated thorax fragments (Pl. 94/2, 4). Because the shields and segments are turned under on the ventral side, in thin sections the inflected margins of the segments look like tiny hooks causing the appearance of ‘shepherd’s crooks’ (Pl. 94/1, 3). Basics: Trilobites Kaesler, R.L. (1987): Trilobitomorpha. – In: Boardman, R.S., Cheetham, A.H., Rowell, A.J. (eds.): Fossil invertebrates. – 221-240, Palo Alto (Blackwell) Moore, R.C. (1959): Arthropoda 1. Trilobita. – Treatise on Invertebrate Paleontology. Part I, 560 pp., Lawrence (Geological Society of America) Further reading: K126
10.2.4.9 Echinoderms Echinoderms are marine invertebrates with a multi-plate calcareous internal skeleton embedded in the skin, a marked five-rayed symmetry and a water vascular system through which the water is circulated in the body. With few exceptions, modern echinoderms are benthic bottom dwellers. A few can swim or float. The phylum Echinodermata, known since the Cambrian, embraces many major subgroups, some of which are important rock builders during the Phanerozoic (e.g. crinoids, echinoids). Echinoderm fragments are present in limestones formed in shallow-marine as well as in deep-marine environments. Most Paleozoic echinoderms were sessile.
Trilobites
Microstructure and recognition of echinoderms in thin sections: The echinoderm skeleton is made of plates, spines or sclerites. Each skeleton element consists of a three-dimensional porous network composed of crystallographically uniformly oriented carbonate crystals (stereom: a latticework constructed from lathes of HighMg calcite) and organic tissue (stroma) within the meshes of the network (Fig. 10.52). Skeleton elements exhibit unit extinction between crossed nicols and correspond to a single calcite crystal. The Mg-calcite elements transform to calcite with some changes in the preservation of the original microstructure (Dickson 2001). Echinoderm grains are usually preserved as large single crystals of calcite. The sparry preservation is found irrespective of whether the primary porous network is filled with syntaxial calcite or with pyrite, hematite, glauconite or other minerals. Micritic preservation of echinoderm fragments is related to crystal diminution by diagenetic dissolution (Evamy and Shearman 1965; Neugebauer 1978; Neugebauer and Ruhrmann 1978; Richter 1974). Usually echinoderm skeletons disarticulate and the skeletal plates are scattered in the sediment. Individual plates are commonly millimeter-sized to a few centimeters in maximum dimension. In thin sections, echinoderm fragments are easily recognized by the uniform extinction, the dusty appearance and often gray or yellowish colors in transmitted light, and the calcite latticework. Many echinoderm plates exhibit a finer more dense meshwork towards the plate surfaces, appearing
Plate 94 Trilobites The plate displays trilobite limestones formed in low- and high-energy settings of open-marine platforms. Different skeletal parts occur in texturally different limestones, pointing to transport sorting. Grainstones yield head-shield and thoracal fragments (–> 1, 3). Wackestones commonly contain variously sized thoracal fragments (–> 2). Packstones are formed by accumulations of pygidia (–> 7). 1 Trilobite rudstone. Various trilobite sections: Semicircular section (1) of a trilobite head shield (cephalon). Note welldifferentiated axial region (glabella, G) and long spines (S). Recurved thoracal segments (2). Open-marine platform. Early Devonian: Hamar Laghdad, Erfoud, southern Anti-Atlas, Morocco. 2 Trilobite wackestone. Poorly sorted trilobite debris in micritic matrix. Same locality as –> 1. 3 Microstructure. Characteristic ‘shepherd’s crook’ structure, caused by recurved lateral borders of thoracic segments. The homogeneous prismatic microstructure encloses fine pores (arrows) running perpendicular to the surface. Intergranular spaces between the trilobite fragments are filled with radiaxial calcite cement (RC). Same locality as –> 1. 4 Typical criteria. Recurved trilobite fragment (T) and echinoderm (E) bioclast within a micritic matrix. Same locality as –> 1. 5 Taphonomy. Microbially encrusted trilobite fragment. Late Silurian: Valentin-Törl, central Carnic Alps, Austria. 6 Typical criteria. Trilobite spines within a fine-grained quartz sandstone. Note the fine prismatic microstructure of the thin elements. Middle Cambrian (Paradoxides Sandstone): Öland Island, Sweden. 7 Trilobite packstone. Densely packed and parallel arrangement of trilobite fragments forming a coquina. Most fragments are parts of tail shields of trilobites living near mud mounds. SMF 12-TRILO. Same locality as –> 1.
Plate 94: Trilobites
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Crinoidea
gray in thin sections, and an open meshwork within the plates. Pore-filling cement occurs commonly within the host skeleton. Many echinoderm bioclasts, particularly those in micritic limestones, exhibit syntaxial overgrowth of clear calcite cement, because the echinoderm plates act as nucleation sites for calcite cement (Pl. 34/ 3, 4). The genetic interpretation of this cement is discussed in Sect. 7.4.2.1. Many echinoderm limestones are cemented accumulations of echinoderm plates that have broken along the calcite cleavage planes. Some echinoderm fragments show characteristic shapes in thin sections (Fig. 10.53). Columnal plates of stalked echinoderms (crinoids, cystoids, blastoids) appear round or pentagonal in cross sections (Fig. 10.54). Stem plates are centrally perforated by one or several holes that may have circular, elliptical or pentagonal outlines. Sections parallel to the stem exhibit rectangular figures. Stem plates still preserved in series are subdivided by crenulated joint surfaces. Crinoid arm plates may reveal lunate patterns when cut perpendicular to the arm axis and rectangular patterns or two parallel plates when cut parallel to the arm axis. Pore plates of echinoids cut parallel to the plate surface show characteristic patterns (Durham et al. 1966). Classification: Echinoderms fall into two major groups including forms attached to the seafloor (Pelmatozoa) and crawling and burrowing forms (Eleutherozoa). The Pelmatozoa include the Crinozoa with the Crinoidea (sea lilies) and the exclusively Paleozoic Blastozoa. The Eleutherozoa comprise the Asterozoa with the Asteroidea (starfish) and the Ophiuroidea, the Paleozoic Homalozoa and the Echinozoa with the Echinoidea (sea urchins), Holothuroidea and further extinct groups. Most of these groups appeared in the Ordovician, the crinoids and the Homalozoa already in the Cambrian. Modern echinoderms comprise crinoids, asteroids, ophiuroids and echinoids, holothurians. Identification of dissociated echinoderm fragments: The elements of echinoderm skeletons are commonly disarticulated by taphonomic processes. Echinoderm fragments are frequent constituents of many Phanerozoic carbonates, but generally they are listed without any further assessment except as ‘crinoids’ or ‘echinoids’. Microfacies analyses, however, require more detailed differentiations. This is possible if specific classification systems developed for isolated skeletal elements of crinoids and echinoids are used (Moore and Jeffords 1968; Termier and Termier 1974; see Fig. 10.55, Pl. 135/5-7).
Fig. 10.52. Criteria of echinoderm fragments in thin sections. A: Oblique tangential section of a crinoid stem exhibiting three columnal plates and a side view of the serrate surface between them. Note the parallel cleavage planes caused by recrystallization. Mississippian: Frank’s land slide, Alberta, Canada. B: Side view of two crinoid columnals still exhibiting the original meshwork microstructure. Early Permian: Carnic Alps, Austria. Scale is 1 mm.
Crinoidea Crinoids are passive filter-feeding organisms. The cup-shaped body (theca) with five simple or branched flexible arms is connected to the bottom via a stem located on the underside of the body. The stem is constructed of a series of interlocking disks called columnals each of which is centrally perforated forming a lumen. The lumen houses the nerves involved in the movement of the animal. Stemmed forms frequently have holdfasts composed of small polygonal plates. The calyx carries long mobile arms which support featherlike structures (pinnules) acting as filtration fan. Stalked crinoids were abundant in shelf environments of the Paleozoic and Mesozoic. Crinoids were the most abundant group of echinoderms from the Early Ordovician to the Late Paleozoic. Today the group is common in deep oceans. The common fossil record of crinoid skeletons are dissociated elements, usually of the stems or the arms. Crinoids disintegrate in a mass of arm and cirri fragments (Allison 1990). Highly porous crinoid ossicles are entrained at lower current velocities than
Echinoderms in Thin Sections
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Crinoid limestones Crinoid limestones are commonly formed by mass accumulations of benthic crinoid elements, but can also originate from the deposition of fragments of pelagic free-living crinoids.
hydraulically sorted components of tempestites, in crinoidal mounds (Becker and Dodd 1994) on slopes and in basins. Widely extended ‘regional encrinites’ (Ausich 1997) were deposited below normal wave base but within storm wave base. Crinoid limestones are regarded as a specific standard microfacies type (see Pl. 119/2, Pl. 139/6). Post-Cretaceous crinoid deposits are rare, because of the change in the habitat of stalked crinoids from shallow shelf and upper slope environments to deeper-water basinal environments. Crinoids thrived in carbonate settings to siliciclastic-dominated settings. Limestones rich in crinoids exhibit differences in the abundance and composition of crinoidal elements, indicating various taphonomic processes and differences in the depositional environment: (a) Many primary crinoid packstones consist almost exclusively of isolated, often worn disks and plates. This may indicate in-place degradation or deposition subsequent to transport. These packstones (encrinites) can form thick beds. (b) Crinoid ossicles may also occur within bioclastic wackestones and marly limestones, but to a minor percentage and admixed with other skeletal grains. This may indicate transport prior to deposition. (c) Long stems and even well-preserved crowns occur in marly limestones and marls. This characterizes more or less autochthonous burial in quiet-water settings. Crinoidal limestones (usually wackestones and packstones, rarely grainstones) originated in place by autochthonous or parautochthonous deposition of crinoidal elements or allochthonous subsequent to transport (Cain 1968):
Benthic crinoidal limestones are formed in different marine depositional environments. Mass accumulations of crinoids occur in reef flank beds (Pl. 144/4-6), as
In-place deposition occurred in reef, near-reef (Fig. 10.55; Pl.143/1), and far-reef settings (Ruhrmann 1971a, 1971b), in crinoidal mounds (Becker and Dodd
Fig. 10.53. Schematic drawings of common sections of echinoderm plates. A and B: Cross sections of crinoid columnals differing in the shape of the central lumina. C: Columnal differing from A and B in the pentagonal outline of the stem plate. D: Section of a stem showing the serrate outline of the articulated surfaces. E: Series of stem plates displaying serrated outlines of stem surfaces. F: Same as E but section passing through the lumen. G: Oblique section through a crinoid stem. H: Lunate crinoid arm plate. Sections along a–a’ and b–b’ will create different pictures, I: Echinoid pore plates parallel to the surface of the plates. Modified from Horowitz and Potter (1971).
solid carbonate particles of the same size (Savarese et al. 1997). Stalkless crinoids (feather stars) appeared in the Mesozoic. Today they live in shallow-marine waters, e.g. reefs. Their isolated elements may be conspicuous constituents of Mesozoic open-marine limestones.
Fig. 10.54. Isolated crinoid columnals in thin sections differing in outlines and surface ornamentation. A: Pentagonal stem plate with several ringlike cirri sockets. Early Jurassic (Sinemurian): Klauskogel near Hallstatt, Austria. B: Pentacrinus exhibiting the original meshlike microstructure especially in the center. Early Permian: Southern Anatolia, Turkey. C: Circular stem plate. The original microstructure has largely vanished. Note cleavage planes. The columnal is encrusted by foraminifera. Late Carboniferous: Rosskofel, Carnic Alps, Austria. Scale is 1 mm.
552
Crinoid Limestones
Fig. 10.55. Crinoid limestone. Most authors avoid a more detailed morphologic differentiation of disarticulated elements of benthic crinoids in thin sections. However, the differentiation of dissociated crinoid elements (Moore and Jeffords 1968) in washed samples and thin sections can assist in answering both stratigraphic and sedimentological problems (Stukalina 1988; Ures et al. 1999). The thin section exhibits several morphological features important for the taxonomic differentiation of dissociated parts of crinoid stems. (1) Columnals: Individual components of the crinoid stalk. (2) Pluricolumnals: Two or more columnals attached to each other. (3) Latera: Outer surface of crinoid columnals with variously shaped profiles. (4N) and (4IN): Nodals and internodals: Larger and intercalated smaller columnals of a crinoid stalk. (5) Articula: Smooth or sculptured surface of columnals serving for articulation with contiguous stem element. (6) Sutures of columnal articulation: Externally visible edges of articula. (7) Axial canals: Longitudinal passageway penetrating the columnals, varying in transverse shape and median longitudinal sections; complex axial canals are characterized by successive alternating columnal restrictions and intercolumnal expansion. This sediment was deposited adjacent to biostromal stromatoporoid-coral reefs. Crinoids lived at the flanks of the reefs. The very poor sorting of crinoid fragments associated with well-preserved crinoid columnals indicates burial in quiet-water settings without longer transport. Large angular stromatoporoid fragments (S) were admixed with the echinoderm detritus. Late Silurian (Hoburgen limestone, Hamra/Sundre Beds, Ludlowian): Grötlinbo quarry, Hoburgen, Gotland Island, Sweden.
1994; Pl. 119/2) and also in slope environments (Michalik et al. 1993). Parautochthonous crinoid accumulations were formed in Silurian and Devonian forereef areas. Crinoids were major contributors to the pioneer stage of Paleozoic reefs (see Sect. 16.2.6.2). Another type of autochthonous deposits are regionally extended accumulations of crinoids (‘regional encrinites’; Ausich 1997). Allochthonous deposits formed on subtidal ramps (e.g. in the Early Carboniferous Waulsortian crinoid-
bryozoan facies behind autochthonous accumulations in mud mounds), on platforms, slopes, basins as well as on seamounts. Deposition in subtidal environments is frequently connected with storm-induced transport of crinoid elements contributing to tempestites (Aigner 1984, 1985; Pl. 95/2). Slope and basinal setting of crinoid accumulations may be associated with a downward transport of crinoids from shallow or upper slope environments leading to the formation of crinoidal turbidites (Tucker 1969; Pl. 139/3). An upslope transport of off-reef crinoids owing to the fact that crinoids can
Crinoid Limestones
553
float as long as the skeletal elements are filled with organic tissue, was discussed by Ruhrmann (1971b). Submarine fissures and neptunian dikes formed after drowning of carbonate platforms can act as traps for crinoid accumulations (Pl. 24/4, Pl. 139/6). Deposition on seamounts is believed to have been controlled by submarine currents (Jenkyns 1971). The differentiation of depositional environments of crinoid limestones requires answers to the following questions: • Which elements of the crinoids have been preferentially deposited (stems, columnals, plates, brachials etc.)? • How well preserved are the elements (e.g. totally or partly disarticulated, long crinoid stems or even crinoid crowns)? • Are the crinoid elements broken, worn, abraded or conspicuously altered (Pl. 95/3, 4)? • Do crinoid stems show signs of current orientation? • Are the elements bored or encrusted (Pl. 52/6, 7)? • How abundant are the crinoids with regard to the rock volume? • Are the crinoid elements grain- or mud-supported? • What about the sorting of crinoids and other grains? • Do crinoids contribute to specific depositional fabrics (e.g. gradation, lamination)? • What other skeletal and non-skeletal grains are associated with the crinoids (Pl. 95/1)? Other important questions in the context of the genetic interpretation of crinoid limestones refer to field criteria (geometry, sedimentary structures), lateral and vertical extension of the crinoid limestones, and interbedding with other facies types. Planktonic crinoids Elements of non-stalked planktonic crinoids are important, sometimes rock-forming constituents of pelagic carbonates. These crinoids are known from the Triassic of the Alpine-Mediterranean region (roveocrinoids)
Fig. 10.56. Saccocoma limestone. A: This microfacies is characterized by numerous antler-shaped calcitic structures floating in a fine-grained matrix of pelagic micrite. Late Jurassic (Kimmeridgian): Central Apennines, Italy. B: Cross section of a Saccocoma arm element (brachialia). Late Jurassic: Thurburbo majus, northern Tunisia. Scale is 500 µm.
and occur with Saccocoma Agassiz in great abundance in Tethyan Jurassic open-marine limestones as well as in epicontinental shelf carbonates (Fig. 10.56). Triassic roveocrinoids (Pl. 96/1) occur in wide distribution in the Tethys within a specific time interval comprising the Late Ladinian and the Early Carnian (Kristan-Tollmann 1970). The first record of Saccocoma is from late Middle Jurassic, mass occurrences are abundant within the time interval Kimmeridgian to middle Tithonian. The planktonic life-style of Saccocoma was questioned by Manni et al. (1997) who favored a possible benthic mode of life. Anatomical and taphonomical criteria as well as the distributional pattern argue against this interpretation (Keupp and Matyszkiewicz 1997). Fig. 10.57. Echinoid spines with stereom-filled center. The pattern corresponds to the Cidaris type distinguished by Hesse (1900). A: The reticulate structure of the central zone is surrounded by a zone of densely arranged radial septa. The septa are connected by rings of interseptal bars. B: The general structural pattern resembles that of A but the spine is covered by a distinct peripheral border. A and B: Early Permian: Forni Avoltri, Carnia, Italy. Scale is 500 µm.
554
Echinoids
Fig. 10.58. Holothurian sclerites. A: Thin section exhibiting two different sclerites. Early Cretaceous: Subsurface, Bahama Bank. B: SEM photograph of Theelia florida (Terquem and Bertelin). Early Jurassic: Empede near Hannover, Germany. B: After Fischer et al. (1986). Scale is 100 µm.
Thin sections of Saccocoma elements were described as Eothrix and Lombardia. Verniory (1955) proved that these fossils were sections of facetalia and brachialia of Saccocoma and closely related taxa. Characteristic sections exhibit antler-like shapes (Fig. 10.56; Pl. 96/2, 3). In the Upper Jurassic central European epicontinental platforms, the Saccocoma facies characterizes late transgressive system tract and highstand deposits. Within the Tethyan region, the Saccocoma facies is common in deep-marine basinal sequences. Allochthonous deposition by turbidites on slopes forming Saccocoma packstones/wackestones are known from deep-water settings as well as epicontinental positions (Matyszkiewicz 1996).
Echinoidea Modern echinoids occur in almost every major marine habitat from intertidal zones to depths of more than 5000 meters, and from the poles to the equator. Regularia are epibenthic detritus feeders living especially on hard bottoms. Irregularia occur within or on soft bottoms. Morphology and classification: The skeleton is made of tightly interlocking plates that form a rigid structure (test) whose shape ranges from globular to flattened. Echinoids are traditionally divided into two subgroups. The regular echinoids with almost perfect pentameral symmetry include the Cidaroidea (since Late Paleozoic) and the Echinoidea s.str. (sea urchins;
Plate 95 Echinoderms The plate exhibits Paleozoic limestones rich in crinoid plates (–> 1-4) and remains of cystoidean echinoderms (–> 6) as well as a Triassic limestone with echinoid fragments (–> 5). The figure on this page demonstrates the morphology of a stalked crinoid (after Black 1988). 1 Crinoid fragments. Bioclastic wackestone with echinoderms (E; predominantly crinoids) and shells of pelagic bivalves (‘filaments’, F). The crinoid sections correspond to the sections A, D and H shown in Fig. 10.53. The biofabric is characterized by loose and disperse packing and bimodal sorting; see –> 4. Open-marine deep slope environment. Early Jurassic: Adnet near Salzburg, Austria. 2 Crinoid accumulation. Crinoid stem plates (columnals) within a tempestite. Some crinoids exhibit syntaxial cement rims. Most cementation, however, is related to fractures. SMF 12-CRINO. Early Devonian (Emsian): Erfoud, southern Anti-Atlas, Morocco. 3 Crinoid stem. Longitudinal section through three columnals. Note grayish color and distinct twin lamellae. Early Carboniferous open-marine carbonate ramp (Viséan): Debnik, Cracov Upland, central-south Poland. 4 Crinoid packstone. The columnal (C) is completely recrystallized. Black arrow points to microborings, white arrows to the serrate boundary between stem plates. B: Bryozoan. The biofabric is densely packed and well-sorted. The sample is from the deeper part of a carbonate ramp. Mississippian: Turtle Mountains, Frank Slide, Alberta, Canada. 5 Echinoids. Bored echinoid plates (EP) and echinoid spines (ES) associated with gastropods (G). White grains are quartz. Late Triassic: Gurumugl, Mount Hagen, central Papua New Guinea. 6 Cystoid wackestone. Cystoids were important Early Paleozoic rock builders (Cambrian–Devonian). The thin section shows broken and eroded plates occurring in a strongly recrystallized limestone rich in cystoids. Black arrows point to syntaxial overgrowth rim cement. B: Bryozoan. Note the difference in color between echinoderms (dusty gray) and bryozoans (dark). Late Ordovician (Ashgillian): Central Carnic Alps, Austria.
Plate 95: Echinoderms
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Echinoid Spines
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since Late Triassic). In the irregular echinoids a secondary bilateral symmetry is superimposed upon the primary pentaradiate pattern. This group includes the Cidaroidea (pencil urchins), Clypasteroidea (sand dollars; since the Paleocene), Spatangoidea (heart urchins; since the Cretaceous) and Cassiduolidea (since Jurassic). Echinoid spines in thin sections: Echinoid fragments can be identified using a combination of morphological criteria (Durham et al. 1966; Smith 1984; Nebelsick 1992). Common constituents in thin sections of limestones are echinoid spines. Living echinoids are covered with movable spines, which are anchored in sockets on the top of tubercles in the test and serve for protection, walking and burrowing.
Echinoid spines show a great variety in shape and size. The spines are long and prominent in regular echinoids, but very short and stubby, forming an almost felt-like covering, in irregular echinoids. These spines are hollow and without a cortex. A spine consists of three parts, a long shaft, a short neck and a base. The concave articulation surface that attaches onto the mamelon of the tubercle is the acetabulum which may be either perforate or imperforate. Each spine can be moved individually. The shaft may be smooth, very finely striated or ornamented by ribs, thorns or granules. Sections of echinoid spines exhibit characteristic internal radial patterns that are useful taxonomically. In cross section spines may be hollow (diadematoid type)
Plate 96 Echinoderms The plate displays skeletal elements of non-stalked pelagic crinoids (–> 1-3) and stalked crinoids (–> 11), echinoid spines (–> 6–10, 12, 13), and a holothurian sclerite (–> 4). Taphonomic criteria of echinoderms are shown in –> 14 and 15. The figure on this page shows a recent non-stalked crinoid. After Black (1988). 1–3 Pelagic ‘free-swimming’ crinoids. Skeletal elements of planktonic crinoids are important constituents of Mesozoic open-marine carbonates. Most common elements seen in thin sections are brachial extensions used for swimming. 1 Roveocrinoids. Oblique sections of thin arm plates (brachialia) of Osteocrinus rectus (Frizell and Exline). Late Triassic (Carnian): Feuerkogel near Aussee, Austria. 2 Microfacies with Saccocoma. Accumulation of antler-like brachial plates and other elements of the non-stalked crinoid. Note the fine meshwork structure (white arrow) diagnostic for echinoderms. Black arrows point to zoospores of planktonic algae (Globochaete; see Sect. 10.2.6). This microfacies is widely distributed in Kimmeridgian to Tithonian basinal limestones of the Tethys. Late Jurassic: Margin of the Northern Alps near Salzburg, Austria. 3 Saccocoma. Arm plate. Late Jurassic (Saccocoma limestone, Kimmeridgian): Central Apennines, Italy. 4 Holothurian sclerite. Skeletal elements of holothurians differ from those of other echinoderm groups in that they consist of isolated microscopic single-crystal calcite pieces. Morphology is highly variable and includes anchors, hooks, tables, rods, wheels (this figure) and others. Early Cretaceous: Subsurface, Bahamas. 5–10, 12, 13 Echinoid spines: The tests of echinoids are covered by spines of different size, serving as protection, locomotion, or as means of anchoring. Cross sections exhibit a radial pattern around a hollow or irregularly fabricated center. Differences in radial patterns may be useful taxonomically and for stratigraphic zonation. 5 Echinoid spine. Longitudinal section. The spine projects outward from a tubercle (T). Late Triassic: Kuta, Papua New Guinea. 6–8 Echinoid spines. Cross sections. The three sections correspond to each other in the perforated thin zone surrounding the lumen, and the general pattern of the wall zone. The wall zone consists of wedge-shaped radial elements which widen toward the periphery. The cross sections differ in the shape and the number of the radial septa. The general pattern corresponds to the Diadema type distinguished by Hesse (1900). Cross sections exhibit either low birefringent colors or extinction under crossed nicols. Same locality as –> 5. 9 Echinoid spine. Cross-section and oblique section. Same locality as –> 5. 10 Echinoid spine. Note the syntaxial overgrowth cement. Late Jurassic (Tithonian): Subsurface, Kinsau, Germany. 11 Crinoid arm elements (brachial fragments), characterized by bifurcated lunate patterns (see Fig. 10.53). Early Jurassic: Adnet, Salzburg, Austria. 12 Echinoid spine. Cross section. The spine differs from those shown in –> 6–8 and 13 by the arrangement of the radial septa, distinct ring-like interseptal connections and by the larger size of the spine. Late Permian: Bergama, western Anatolia, Turkey. 13 Echinoid spine. Cross section. Same locality as –> 5. 14 Destruction of echinoderm plates by boring. Arrows point to cross sections of cricoconarid shells (see Pl. 91). Early Devonian (Emsian): Erfoud, Anti-Atlas, Morocco. 15 Echinoderms and mollusks compared. In thin sections, echinoderm (E) elements differ from mollusk (M) shells in color and net-like microstructure. Mollusk shells are commonly replaced by a coarse calcitic mosaic. Same locality as –> 5.
Plate 96: Echinoderms
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or solid (cidaroid type). The hollow center is called the lumen. The solid center filled with stereom is termed the medulla (Fig. 10.57), the surrounding zone is the cortex. A still useful classification of echinoid spines based on the microstructure was established by Hesse (1900) who distinguished six major groups (see Pl. 96/ 6–8, 12, 13 for examples). Distribution: The oldest echinoids appeared in the Late Ordovician but were only a minor element of the benthos throughout the Paleozoic. The roots of modern echinoids can be traced back to the Permian where the first cidaroids appeared and started to diversify in the Triassic. A major diversification of echinoids took place in the Jurassic. The first irregular echinoids specialized in deposit feeding appeared in the Jurassic. Echinoids generally became a major constituent of the shallow-water benthos during the Jurassic and Cretaceous, also in deep-water environments (chalk) in the Late Cretaceous.
Ophiuroidea and Holothuroidea Remains of echinoderm groups other than crinoids and echinoids are extremely rare in thin sections of limestones, although skeletal elements of holothurians and ophiuroids are often abundant in acid residues of carbonates. Cross sections of ophiuroid vertebrae from Mesozoic and Late Tertiary limestones exhibit two coalesced calcite plates, each of which constitutes a single crystal, forming a mirrowed reflection pattern (Glazek and Radwanski 1968). Holothurian sclerites may be represented in thin sections by wheel-like structures (Fig. 10.58; Pl. 96/4) or sieve plates. Basics: Echinoderms Donovan, S.K. (1991): The taphonomy of echinoderms; calcareous multi-element skeletons in the marine environment. – In: Donovan, S.K. (ed.): The process of fossilization. – 241-269, London (Belhaven Press) Glazek, J., Radwanski, A. (1968): Determination of brittle star vertebrae in thin sections. – Bulletin de l’Académie polonaise des Sciences, Série des Sciences géologique et géographique, 16, 91-96 Guensburg, T.E., Sprinkle, J. (1992): Rise of echinoderms in the Paleozoic evolutionary fauna: significance of paleoenvironmental controls. – Geology, 20, 407-410 Hess, H., Ausich, W.E., Brett, C.E., Simms, M.J. (1999): Fossil crinoids. – 275 pp., Cambridge (University Press) Hesse, W. (1900): Die Mikrostruktur der fossilen Echinoideenstacheln und deren systematische Bedeutung. – Neues Jahrbuch für Mineralogie, Geologie und Paläontologie, Abhandlungen, 13, 185-264
Rare Thin-Section Fossils
Keupp, H., Matyszkiewicz, J. (1997): Zur Fazies-Relevanz von Saccocoma-Resten (Schwebcrinoiden) in OberjuraKalken des nördlichen Tethys-Schelfs. – Geologische Blätter für Nordost-Bayern, 47, 53-70 Kiaer, P.M. (1987): Class Echinoidea. – In: Boardman, R.S., Cheetham, A.M., Rowell, A.J. (eds.): Invertebrate fossils. – 596-611, Palo Alto (Blackwell) Kristan-Tollmann, E. (1977): Zur Gattungsunterscheidung und Rekonstruktion der triadischen Schwebcrinoiden. – Paläontologische Zeitschrift, 51, 195-198 Moffat, H.A., Bottjer, D.J. (1999): Echinoid concentration beds: two examples from the stratigraphic spectrum. – Palaeogeography, Palaeoclimatology, Palaeoecology, 149, 329-348 Moore, R.C. (ed., 1966): Echinodermata 2. AsterozoaEchinozoa. – Treatise on Invertebrate Paleontology, part U, 725 pp., Lawrence (Geological Society of America) Moore, R.C. (ed., 1967): Echinodermata 1. General characters Homalozoa-Crinozoa (except Crinoidea). – Treatise on Invertebrate Paleontology, part S, 679 pp., Lawrence (Geological Society of America) Moore, R.C., Jeffords, R.M. (1968): Classification and nomenclature of fossil crinoids based on dissociated parts of their columns. – University of Kansas Paleontological Contributions, 9, 1-86 Moore, R.C., Teichert, C. (eds., 1978): Echinodermata 2. Crinoidea. – Treatise on Invertebrate Paleontology, part T, 411 pp., Lawrence (Geological Society of America) Nebelsick, J. (1992): Echinoid distribution by fragment identification in the Northern Bay of Safaga, Red Sea, Egypt. – Palaios, 7, 316-328 Nebelsick, J. (1995): Uses and limitations of actuopaleontological investigations on echinoids. – Geobios, 18, 329338 Savarese, M., Dodd, J.R., Lane, N.G. (1997): Taphonomic and sedimentologic implications of crinoid skeletal porosity. – Lethaia, 29, 141-156 Smith, A.B. (1990): Biomineralization in echinoderms. – In: Carter, J.G. (ed.): Skeletal biomineralization: patterns, processes and evolutionary trends. Volume I. – 413-443, New York (Van Nostrand) Stukalina, G.A. (1988): Studies in Palaeozoic crinoid-columnals and stems. – Palaeontographica, A, 204, 1-66 Further reading: K128
10.2.5 Rare Thin-Section Fossils In contrast to the preceding chapters, this chapter and plate 97 deal with fossils that may be recognized in thin sections of limestones, but are generally rare. Vertebrate remains seen in thin-sections are predominantly bones, scales and teeth of fish. Bones may have ovate, rectangular or round shapes. Remains of amphibians and reptiles are very rare but may occur in lacustrine carbonates. These phosphatic vertebrate grains display clear, amber or brownish, sometimes black colors in thin sections under ordinary light. Be-
Microproblematica
tween crossed nicols they may exhibit very low birefringence and undulose extinction. Many fish bones show an interior vesicular or open-network structure and a dense more compact outer structure (Pl. 97/2). Fine lamellar microstructures, sometimes with fine tubular pores, are common in bones, teeth and scales (Pl. 97/9). Conodonts (Cambrian to Late Triassic) are small (about 0.2 to about 3 mm in length) tooth-like microfossils consisting of calcium phosphate. They have the shape of serrated bars, ornamented platforms, and blades. They are characteristically brown, gray or black in thin sections (Pl. 97/6-8) and sometimes show a lamellar microstructure. Conodonts are generally rare in thin sections, but may be a minor constituent in limestones occurring in condensation sequences. Graptolites (Cambrian to Early Devonian) are usually found in black shales, but specimens may also be preserved in limestones or cherts (97/5). The skeleton consists of collagen, a fibrous scleroprotein, showing brown or honey colors in transmitted light . Wood can be preserved by infiltration of calcite, silica (Pl. 97/10) or pyrite, leading to preservation of the organic or carbonized material consisting of cells, sometimes with secondarily thickened walls. Wood fragments may be minor constituents of lacustrine and marine near-coastal carbonates.
10.2.6 Microproblematica The last part dealing with the criteria of fossils found in thin sections concerns microfossils that are enigmatic with respect to their systematic position. Microproblematica are micron- and millimeter-sized fossils known from acid residues, washed samples and thin sections, that cannot readily be placed in established phyla or major groups. Some microproblematica, described from thin sections may result from difficulties in recognizing threedimensional geometries from two-dimensional views. But many ‘thin-section microproblematica’ exhibit distinct morphologies, allowing them to be treated taxonomically. Some of these enigmatic fossils have been potentially assigned to the algae, but this assignment is still controversial (Babcock 1986). Some taxa are restricted to specific paleoenvironments and some even represent index fossils (Hoffman and Nitecki 1986).
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10.2.6.1
Inventory of Microproblematica
Box 10.8 gives a list of fossils that have been described only from thin sections and are regarded as microproblematica by many authors. Most Mesozoic taxa are benthic and occur in reef limestones and shallow-marine platform carbonates. Some taxa are planktonic and may be common in pelagic open-marine limestones (e.g. Globochaete; Pl. 99/1). Microproblematica are remarkably abundant in reef communities, predominantly acting as encrusters (see Sect. 9.2.3). Encrusting microproblematica are particularly abundant in reef carbonates, occurring both at the reef surface between and upon reef builders (e.g. Bacinella and Lithocodium; Pl. 99/4, 5) or in small cryptic habitats of reef cavities (e.g. Baccanella and Radiomura; Pl. 99/8, 9). Cavities in modern reefs comprise between 30 and 75 % of the total reef volume. They are inhabited by sessile, boring and encrusting ‘coelobionts’ that contribute greatly to the recycling of nutrients, the production of fine-grained internal sediments and the rapid cementation of reef rocks (Ginsburg 1983; Hutchins 1983; Choi 1984). The composition and abundance of coelobionts are related to water energy and flushing within the reef body. Common coelobionts that are preserved in the fossil record include foraminifera, sponges, bryozoans, mollusks, serpulids, algae, microbes and many microproblematica (Kobluk 1988). The taxonomic composition of microproblematica contributing to coelobiont assemblages has changed over times as demonstrated by Triassic and Jurassic reef communities. Common microproblematica of Middle Triassic and Carnian reefs in the entire Tethyan region are distinctly different from those of Norian and Rhaetian reefs (Flügel 2002). The former are represented by Aeolisaccus, Baccanella, Bacinella, Barbafera, Ladinella, Lamellitubus, Radiomura and Tubiphytes; the latter by Lithocodium, Microtubus, Radiomura and Thaumatoporella. Abundant Late Jurassic encrusting microproblematica are Bacinella, Lithocodium-Troglotella and Tubiphytes (Schmid 1996). The differences become more striking at a species level. Many microproblematica can be assigned to informal groups characterized by the dominating morphology: • Hollow tubes (e.g. Aeolisaccus, Koivaella, Lamellitubus, Macrotubus, Microtubus, Paraeolisaccus) • Isolated or closely packed spheroids consisting of radially arranged calcite crystals or spheroids (e.g. Baccanella, Fusanella, Gemeridella, Globochaete, Muranella, Microcodium, Nuia)
560
• Cell assemblages (e.g. Bacinella, Ladinella, Thaumatoporella) • Differentiated crusts and plates (e.g. Lithocodium, Palaeoaplysina).
10.2.6.2 Discussion of Selected Microproblematica The following alphabetic discussion of the morphology and systematic assignment is restricted to some microproblematica that are of importance for paleoenvironmental and sedimentological interpretations. Archaeolithoporella Endo (Pl. 42/1, Pl. 98/8, Pl. 145/1) Morphology: Millimeter- to centimeter thick encrustations on hard substrates (fossils, clasts, marine carbonate cement) consisting of alternating couplets of dark micritic layers and slightly thicker layers of fibrous or blocky calcite. Micritic layers may contain
Rare Thin-Section Fossils
organic matter. Archaeolithoporella also occurs as nodules (‘tebagite’: Razgallah and Vachard 1991). Distribution: Archaeolithoporella is known from the Sakmarian to the Latest Permian. The genus is widely distributed in the western and southwestern Tethys (Slovenia, Sicily, Greece, Tunisia, Oman), in oceanic microcontinent settings (Japan, China), and in marginal basins of Panthalassa (Texas and New Mexico). The genus is a common constituent of Permian shallow subtidal reefs (e.g. Permian Reef Complex in Texas; Pl. 145/1), often associated with coralline sponges (Fig. 10.59) and with Tubiphytes. Together with syngenetic aragonitic cements, Archaeolithoporella formed waveresistant structures (Fig. 10.59). Archaeolithoporella crusts originated in various settings, including shelf edge, platform margin, forereef, and upper slope. Archaeolithoporella does not seem to be totally dependent on lightened areas, as shown by the occurrence within syngenetic fissures and cryptic cavities. Interpretation: The genus has been interpreted as red algae, stromatolitic algae, remains of microbial mats
Plate 97 Rare Thin-Section Fossils The plate displays fossils which are not common in thin sections of limestones, but, if present, may be important in microfacies analysis. 1
Placoderm bonebed. Placoderms are a group of fresh-water and marine fishes with primitive jaws bearing tooth-like body plates and variously armored heads and bodies. The group appeared in the Late Devonian, flourished in the Devonian and disappeared in the Early Carboniferous. Thin sections of limestones yield bone plates exhibiting characteristic double-layered microstructures. The sample is a calcareous sandstone. Early Devonian (Odenspieler bed): Bergisches Land, Germany. 2 Placoderm microstructure. The placoderm plates are characterized by two layers represented by an inner vesicular (white arrows) and an outer, more compact finely vesicular structure (black arrows). Same locality as –> 1. 3 Fish scales. Cross section of numerous thin and flat scales appearing black in transmitted light. The arrows point to shapes that are characteristic in thin sections. Early Jurassic (Lias epsilon): Hetzles near Erlangen, Germany. 4 Fish tooth and fish bones. The black arrow points to a longitudinal section of a tooth of predatory sharks, the white arrow to oblique sections of fish bones. Shark teeth are characterized by sharply pointed shapes. Late Tertiary (Miocene): Saykamt, Turkey. 5 Graptolites. In contrast to the common deposition of graptolites in euxinic basins, graptolites may also accumulate in shallow-marine pelagic carbonate environments. Silurian (Fluminimaggiore Formation, Wenlock): Southwestern Sardinia, Italy. 6 Conodonts: Oblique section of Polygnathus. The sample comes from a strongly recrystallized limestone consisting predominantly of reef debris. Late Devonian: Rugendorf, Frankenwald, Germany. 7 Conodonts morphology. Free blade (arrow) and platform posterior of Polygnathus. The phosphatic conodonts appear black or brown in transmitted light . The sample is from a pelagic cephalopod limestone. Late Devonian: Elbersreuth, Frankenwald, Germany. 8 Conodont morphology. Oblique section of Polygnathus ex gr. P. glabra. Note carina (ridge with small denticles) and parts of the platform. Late Devonian (Famennian): Drewer, Sauerland, Germany. 9 Fish scale. Note distinctive layering (arrow). Middle Triassic (Muschelkalk, Ladinian): Eschenbach, southwestern Germany. 10 Wood. Length and oblique sections of wood fragments, with a well-defined cellular pattern. Preservation of wood is facilitated by carbonification or silicification (this example). Small grains within the micritic matrix are terrigenous quartz. Near-coast mixed siliciclastic-carbonate environment. Middle Jurassic (Bajocian): Morondava Basin, Madagascar. –> 10: Sample provided by M. Geiger (Bremen)
Plate 97: Rare Thin-Section Fossils
561
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Box 10.8. Generic list of thin-section microproblematica. Many taxa may be cyanobacteria or algae but their systematic position is not yet fully acknowledged. More than 90% of thin-section fossils designated as microproblematica have been described from shallow-marine reef or platform carbonates. Taxa found predominantly or exclusively in open-marine pelagic ‘deep-water’ limestones are marked by an asterisk. Tertiary: Codiomorpha Andrusov and Schalekova 1972, Microcodium Glück 1912, *Pieninia Borza and Misik 1976, Thaumatoporella Raineri 1922 Cretaceous: Aeolisaccus Elliott 1958, Bacinella Radoicic 1959, Carpathoporella Dragastan 1967, Ciprocampylodon Elliott 1963, Cretacicrusta Elliott 1972, *Diademnoides Bonet and Benviniste-Velasquez 1971, *Didemnum Savigny 1816, *Gemeridella Borza and Misik 1975, Getaia Dragastan 1972, *Globochaete Lombard 1945, Koskinobullina Cherchi and Schröder 1979, Lacrymorphus Elliott 1958, Lithocodium Elliott 1956, Microcodium Glück 1912, *Microlaminoides Bonet 1956, Palaeosiphonium Elliott 1985, *Pieninia Borza and Misik 1976, Thaumatoporella Raineri 1922, Tubiphytes Maslov 1956 Jurassic: Aeolisaccus Elliott 1958, Axothrix Nagy 1971, Bacinella Radoicic 1959, *Diademnoides Bonet and Benviniste-Velasquez 1971, *Didemnum Savigny 1816, Enigma Eliasova 1981, Fusanella Eliosova 1981, *Gemeridella Borza and Misik 1975, *Globochaete Lombard 1945, Koskinobullina Cherchi and Schröder 1979, Lithocodium Elliott 1956, Microcodium Glück 1912, Muranella Borza 1975, Palaeosiphonium Elliott 1985, *Palaeothrix Ferasin 1956, Polygonella Elliott 1957, Thaumatoporella Raineri 1922, Tubiphytes Maslov 1956, *Verticillodesmis Dragastan and Borza 1975 Triassic: Actinotubella Senowbari-Daryan 1984, Aeolisaccus Elliott 1958, Baccanella Pantic 1971, Bacinella Radoicic 1959, Barbafera Senowbari-Daryan 1984,*Diademnoides Bonet and Benviniste-Velasquez 1971, *Didemnum Savigny 1816, *Gemeridella Borza and Misik 1975, *Globochaete Lombard 1945, Ladinella Ott 1968, Ladinosphaera Oberhauser 1960, Lamellitubus Ott 1968, Lithocodium Elliott 1956, Macrotubus Fois and Gaetani 1981, Messopotamella Dragastan et al. 1985, Microcodium Glück 1912, Microtubus Flügel 1964, Muranella Borza 1975, Panormidella Senowbari-Daryan 1984, Paraeolisaccus Senowbari-Daryan and Schäfer 1980, Plexoramea Mello 1977, Porferitubus Senowbari-Daryan 1984, Radiomura Senowbari-Daryan and Schäfer 1978, Thaumatoporella Raineri 1922, Tubiphytes Maslov 1956 Permian: Aeolisaccus Elliott 1958. Archaeolithoporella Endo 1959, Koivaella Chuvashov 1974, *Globochaete Lombard 1945, Palaeoaplysina Krotow 1888, Pseudovermiporella Elliott 1958, Tubiphytes Maslov 1956 Carboniferous: Archaeolithoporella Endo 1959, *Globochaete Lombard 1945, Koivaella Chuvashov 1974, Palaeoaplysina Krotow 1888, Tubiphytes Maslov 1956
Microproblematica
and as inorganic structures (Mazzullo and Cys 1977; Wang et al. 1994; Grotzinger and Knoll 1995; Kendall and Iannace 2001). A microbial origin is reliable. Baccanella Pantic 1971 (Pl. 99/8) Morphology: The genus is characterized by small regularly or irregularly developed spheres consisting of radially arranged aggregates of calcite crystals starting from a central point. The spheres are commonly densely packed side by side. The calcite crystal aggregates exhibit shapes resembling leaves of flowers. They have been interpreted as branches that divide into smaller ones near the periphery. The diameter of the individual spheres varies between about 50 and 500 µm. Distribution: The genus is a common constituent of cryptic cavities in the central parts of Middle and Late Triassic reefs. Interpretation: Baccanella has been regarded as algae, bacteria, bacterially induced carbonate precipitate or diagenetic product caused by recrystallization of micritic High-Mg calcite and aragonite on the sea floor (Pratt 1997). The last two explanations are reliable. Bacinella Radoicic 1959 (Pl. 99/4) Morphology: Epibenthic cell assemblages forming crusts or nodules consisting of relatively regular, sometimes parallel irregularly shaped cells. The assemblage is surrounded by an exterior wall. The cells may be circular, rectangular or triangular in cross section. They are filled with sparry calcite. Cell walls are thin. Cell diameter between 100 and 500 µm. Distribution: The genus is known from Middle Triassic to Late Cretaceous deposits. It was originally described from the Cretaceous. Bacinella is a common fossil in Ladinian reef limestones, but occurs also in Late Triassic reef carbonates and is abundant in Jurassic and Early Cretaceous platform carbonates. Bacinella is often intimately associated with Lithocodium. Interpretation: The genus is still enigmatic. It was interpreted as codiacean alga, compared with hydrozoans, and explained as ecotypes of bacterial associations (Neuweiler and Reitner 1992) or as cyanobacteria. Globochaete Lombard 1945 (Pl. 99/1) Morphology: The genus is characterized by spherical and kidney-shaped calcitic bodies with a micritic nucleus, about 30 to 100 µm in size, occurring in clusters, isolated or serially arranged, sometimes attached to thin calcitic plates corresponding to ‘filaments’. An axial cross between crossed nicols indicates a construction of radially arranged calcite fibers. The bodies exhibit a sparry wall, appearing yellowish in transmitted light, and sometimes showing a micritic boundary.
Microproblematica
563
Fig. 10.59. Archaeolithoporella crusts associated with sphinctozoan coralline sponges. The crusts consist of couplets of dark and light irregular and sometimes disturbed layers without distinct microstructures. Note the difference in the preservation of the originally aragonitic sponges and Archaeolithoporella. Both organisms contributed greatly to the formation of extensive Permian reef structures. Middle Permian (Wordian): Oman. Scale is 1 cm.
Globochaete is similar to Gemeridella Borza and Misik, but the latter genus displays a petaloid arrangement of the radial calcitic elements. Distribution: The genus is reported from Silurian to Early Tertiary limestones. Globochaete is abundant in open-marine pelagic carbonates (e.g. in Late Jurassic deep-water carbonates, often associated with Saccocoma), but also occurs in shallow-marine environments (Schäfer and SenowbariDaryan 1980; Bachmann 1987). Interpretation: Originally interpreted as calcified zoospores of chlorophycean alga Globochaete is now regarded as a planktonic single-celled green alga (Skompski 1982). Halysis Hoeg 1932 (Pl. 98/2) Morphology: The genus is characterized by a fanshaped structure forming palisades and chains in cross sections. Distribution: Common in Ordovician boundstones (Munnecke et al. 2001) but also known from Silurian to Middle Devonian reefs and lagoonal limestones. Interpretation: Originally regarded as enigmatic alga, but now assigned to cyanobacteria. Koivaella Chuvashov 1974 (Pl. 98/3) Morphology: The genus is characterized by small, sometimes irregularly shaped tubes with micritic walls and indications of bifurcations.
Distribution: The microproblematicum is abundant in Late Carboniferous and Permian reef and platform carbonates. Interpretation: Commonly interpreted as calcified cyanobacterial tube or attributed to green algae. Lithocodium Elliott 1956 (Pl. 99/5) Morphology: The encrusted partly coiled test exhibits microgranular (calcitic, partly agglutinating) imperforate walls that appear dark in transmitted and brownish in reflected light. The wall has filament-like alveoli (originally probably containing endosymbionts), and is covered by a thin epidermal layer. Chambers show ‘bubble-like’ structures, originally interpreted as algal sporangia but now representing cryptic boring foraminifera. Distribution: Anisian to Late Cretaceous. Often occurring together with Baccanella irregularis. The genus is common in reef limestones. Lithocodium/Bacinella associations contribute to the formation of mud mounds through baffling, binding, frame building and automicrite production (Alsharhan 1987; Neuweiler and Reitner 1992; Schmid 1996). Interpretation: Originally regarded as an encrusting codiacean green alga, the genus is now considered to be a loftusian foraminifera, probably containing photoautotrophic endosymbiotic algae and boring foraminifera (Troglotella incrustans Wernli and Fookes) within the chambers (Schmid and Leinfelder 1995).
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Microproblematica
Microcodium Glück 1912 (Pl. 128/4) Morphology: The name Microcodium refers to structures consisting of aggregates of up to 1 mm long prismatic to pyramidal monocrystalline laths of calcite, with a central cavity in each lath. Depending on the arrangement of the basic units, two types can be distinguished: (a) the units may be radially distributed around a central axis producing a structure resembling an ear of maize, or (b) the units are stacked parallel to each other, along axes that branch in successive planes, thus forming thick crusts. Distribution: First described from Miocene deposits, Microcodium and closely related taxa (Microcodiaceae) have been found in carbonates as old as the Devonian. Many authors consider Microcodium to be restricted to the Cenozoic, associated with lacustrine environments and indicative of pedogenic origin. Interpretation: Older records explained Microcodium as a product of per descensum contamination. However, Paleozoic occurrences demonstrate that microcodiaceans thrived in shallow-marine and supratidal as well as in lacustrine environments (Nassichuk et al. 1986). Microcodium has been assigned to green alga, ‘blue-green alga’ or bacteria (Mamet and Roux 1982).
Plate 98
Some authors have suggested a diagenetic origin. Non-marine occurrences in paleosols and caliche were explained as products of calcite-secreting filamental bacteria living in symbiosis with other organisms (fungi) within the soil (Esteban 1974; Bodergat 1974) or as calcified roots (Klappa 1978; Jaillard et al. 1991). Microtubus Flügel 1964 (Pl. 99/3) Morphology: Small straight or slightly bent hollow tubes with thin walls of microcrystalline calcite. One end of many tubes appears to be enclosed within micritic matrix. The tubes occur isolated or more commonly in aggregates. Tube diameter 0.05 to 0.20 mm, mostly 0.10 mm. Distribution: Abundant in Late Triassic (Norian and Rhaetian) reef limestones, often associated with micritic spongiostromate crusts. Microtubus is known from all parts of the Tethys. It characterizes the central reef facies. Interpretation: The genus was tentatively assigned to serpulid worms but may have cyanobacterial affinity. Nuia Maslov 1956 (Pl. 98/1) Morphology: The genus is characterized by cylindrical bodies consisting of transparent fibrous calcite crystals growing in zones around micritic centers.
Paleozoic Microproblematica
Shown on this plate are common Paleozoic thin-section fossils whose systematic assignment is disputed. 1 Nuia. The arrows point to longitudinal and cross sections of cylindrical bodies characterized by transparent fibrous calcite crystals growing in zones around micritic centers. The fossil is interpreted as a bacterial or cyanobacterial growth form. Nuia is a common constituent of Late Cambrian and Early Ordovician reef mounds. Ordovician: Antarctica. 2 Halysis. The fan-shaped structure exhibiting palisades and chains in cross-sections is now regarded as a cyanophycean taxon. Early Ordovician (Arenigian): Megaconglomerate unit of Cow Head Peninsula, western Newfoundland. 3 Koivaella. Small hollow tubes with well-defined micritic walls are common microfossils in Late Paleozoic reef and platform carbonates. They are interpreted as cyanobacteria or green alga. Early Permian (Sakmarian): Forni Avoltri, Carnia, Italy. 4 Pseudovermiporella sodalica Elliott characterized by attached tubes with perforated walls. The fossil is considered to be a dasyclad alga, a foraminifera or of cyanobacterial origin. Pseudovermiporella is abundant in Late Permian shelf carbonates. Permian: Taurus Range, Turkey. 5 Tubiphytes obscurus Maslov, encrusted by Archaeolithoporella (A). The longitudinal section shows the segmentation of the fossil, the spar-filled cavities (arrows) and the surrounding spider-like network. The current explanation of Late Paleozoic Tubiphytes is a commensalism of different organisms responsible for the internal cavities and tubes and cyanobacteria or other bacteria responsible for the network (see Fig. 10.60). Tubiphytes obscurus is characterized by tubes in each segment. Middle Permian (Murgabian): Sosio, Sicily, Italy. 6 Tubiphytes obscurus Maslov. Cross section. Same locality as –> 5. 7 Palaeoaplysina Krotow, a common Late Carboniferous to Early Permian reef-building organism. The calcareous plates are characterized by an internal differentiation and a canal system. The fossil is regarded as an alga or a sponge. Early Permian (Asselian): Duvansk reef complex, central Urals, Russia. 8 Archaeolithoporella hidensis Endo, an abundant organism in Permian reefs acting as encruster, binder and frame builder. The characteristic criteria are couplets of micritic and microsparitic or sparry irregularly undulated layers. Middle Permian (Wordian): Oman. –> 2: Pohler and James 1989
Plate 98: Paleozoic Microproblematica
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Distribution: Cambrian to Silurian, possibly also Devonian. Nuia is an important rock builder in Late Cambrian and Early Ordovician carbonate mounds. Interpretation: The fossil has been interpreted as codiacean alga, but is now regarded as bacterial or cyanobacterial growth form. Palaeoaplysina Krotow 1888 (Pl. 98/7) Morphology: The genus is characterized by calcareous, originally aragonitic broad tabular plates that appear, wavy in cross sections, with external protuberances and an internal cellular fabric with branched tubes. Length usually 3–10 cm, thickness 0.5–7.0 mm. The plates often occur as parallel layered accumulations. Distribution: Late Carboniferous to Early Permian (Moscovian to Artinskian). Palaeoaplysina built shal-
Microproblematica
low-water mounds in subtropical to temperate paleolatitudes on platforms and in platform-margin areas around the northern margin of Pangea. Some of the reefs are reservoir rocks (e.g. Barents Sea; Stemmerick and Larsen 1993; Ehrenberg et al. 1986). The genus is often associated with phylloid algae. Interpretation: The fossil has been attributed to alga, hydrozoans, and sponges. Most authors prefer an algal interpretation, but a sponge interpretation should be considered as well. Thaumatoporella Raineri 1924 (Pl. 99/2) Morphology: The genus is characterized by small cells arranged in chains forming closed or open irregularly shaped structures, sometimes constructing bridges between sessile organisms, e.g. corals. Some thaumatoporellids exhibit only one chamber, others form sev-
Plate 99 Mesozoic Microproblematica Microproblematica are common in thin sections of Mesozoic limestones and known from shallow-marine and deep-marine carbonates (–> 1, 10). Many taxa are benthic, others planktonic (–> 1, 10). Encrusting microproblematica are particularly abundant in reef carbonates, occurring both, at the reef surface between and upon reefbuilders (–> 4, 5) or in small cryptic habitats of reef cavities (–> 8, 9). The plate shows thin sections of microfossils, whose systematic position is still under discussion, with exception of the calcispheres (–> 10). 1
Globochaete Lombard. These spherical hollow calcitic bodies occur as aggregates or in clusters. They are interpreted as planktonic green algae. The arrows point to a typical serial arrangement. Middle Jurassic (Klauskalk): Grimming, Austria. 2 Thaumatoporella Raineri, characterized by fine cells (black arrow) arranged in chains or enclosing chamber-like structures. White arrows point to globular inner structures. Thaumatoporella is now enclosed within a separate green algal group. Early Jurassic (Pantokrator Limestone): Korfu Island, Greece. 3 Microtubus communis Flügel. These small open tubes commonly occur in aggregates, often embedded within micritic crusts developed on reef builders. Microtubus is a marker fossil indicating the central parts of Norian and Rhaetian reefs. Late Triassic (Oberrhätkalk, Rhaetian): Feichtenstein, Salzburg, Austria. 4 Bacinella. The fossil is characterized by an assemblage of variously formed cells that are sometimes arranged in rows. Bacinella is common in Mesozoic reefs. Late Triassic (Norian): Jabal Kawr, Oman. 5 Lithocodium, encrusting a shell fragment (SF). The peripheral structures indicated by the arrow are considered as ‘pores’ or ‘cortical filaments’ depending on the systematic assignment (foraminifera or udoteacean algae). Lithocodium is now regarded as a foraminiferal consortium. It is often associated with Bacinella (see Pl. 134/2) and occurs in Triassic, Jurassic and Cretaceous reefs. Late Jurassic (Kimmeridgian): Ota, west-central Portugal. 6 Tubiphytes surrounded by a similar but more open network structure called Plexoramea. These fossils are abundant in Middle Triassic and Carnian reef limestones. Middle Triassic (Wettersteinkalk, Ladinian): Blumautal, Germany. 7 Tubiphytes morronensis Crescenti. This species, associated with Terebella (see Pl. 92/10), is abundant in Late Jurassic and Early Cretaceous platform and slope deposits and in limestones with siliceous sponges. The arrow points to a miliolid foraminifera commonly found in the center of the fossil. The Tubiphytes network is interpreted as part of the foraminiferal wall. Late Jurassic (Treuchtlinger Kalk, Kimmeridgian): Southern Franconian Alb, Germany. 8 Baccanella floriformis Pantic. These microfossils (arrow) are common in cryptic cavities of Triassic reefs. They represent most probably bacterially induced carbonate precipitates. Late Triassic (Dachsteinriffkalk, Norian): Gosaukamm, Austria. 9 Radiomura cautica Senowbari-Daryan and Schäfer. The fossil is found commonly attached to frame builders or within micritic cryptic habitats of Triassic and Jurassic reefs. Late Triassic (Oberrhätkalk): Feichtenstein, Salzburg, Austria. 10 Some fossils formerly considered to be microproblematica are no longer problematic with regard to their position. An example are ‘calcispheres’. The thin section shows Stomiosphaera sphaerica (Kaufmann) characterized by spheres, 10100 µm in diameter, with calcitic and organic walls. SEM studies reveal strong morphological differentiation of the wall supporting an interpretation as calcareous dinoflagellate cysts (Sect. 10.2.1.9). Late Cretaceous (Turonian, Seewen limestone): Seewen, Switzerland. –> 8: Gaetani and Gorza 1989
Plate 99: Mesozoic Microproblematica
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Microproblematica: Tubiphytes
Fig. 10.60. Schematic drawing of the internal, commonly spar-filled cavities in Permian Tubiphytes based on thin-section investigations and SEM studies. The network structure is shown in black. Tubiphytes can be differentiated according to the presence (A to C) or absence (D to F) of segmentation and the nature of the axial tube. The latter may correspond to hollow unmineralized cavities restricted to individual segments (A, B) or to multibranched primarily mineralized stems of Rigidocaulis (possibly a calcimicrobe) passing through all segments (C) or the whole non-segmented body (D). In some Tubiphytes the central spar-filled ‘tube’ is a palaeonubeculariid foraminifera (E) or a hollow branched tube with a defined wall, probably identical with Koivaella (F; Pl. 98/3). Type A agrees with the structure of Tubiphytes obscurus Maslov. The Tubiphytes network of the other types differs in the microfabric used for taxonomic discrimination. Not to scale. After SenowbariDaryan and Flügel (1993).
eral chambers and still others are developed as chains without well defined encircled structures (Dragastan et al. 1994). Distribution: The genus is known from Middle Triassic to Late Cretaceous deposits. It is widely distributed in Mesozoic platform carbonates and occurs in lagoonal and reefal settings, frequently with Bacinella. Interpretation: Originally regarded as red alga and later as dasyclad algae or possible sponges, Thaumatoporella is now interpreted as belonging to a separate group of green algae (DeCastro 1990). Tubiphytes Maslov 1960 (Pl. 98/5, 6, Pl. 99/6, 7) Tubiphytes is one of the most common microproblematica in Late Paleozoic and Mesozoic limestones. The name should be replaced by Shamovella for reasons of priority (Riding 1993) but most people still use the familiar old name. Morphology: Oval, lobate, elongated dark encrusting structures exhibiting a fine irregular web-like network surrounding spar-filled cavities (tubes). The fossil appears dark gray in transmitted light and milky white in direct light. The size of individual elements is commonly up to 1.0 mm but may be greater. Taxonomic differentiation is based on central cavities, segmentation and network differentiation (Senowbari-Daryan and Flügel 1993). Vennin et al. (1997) distinguished three morphotypes. Paleozoic and Mesozoic Tubiphytes appear to be different in construction and in the type of the tubes and cavities surrounded by the web-like network. Fig. 10.60 shows the differentiations found in Permian Tubiphytes. In contrast to the Paleozoic Tubiphytes where the spider network surrounding the internal cavities is re-
garded as functionally different from the cavities, the interpretation of Jurassic ‘Tubiphytes’ considers the network as the exterior part of a micropeloidal foraminiferal wall formed by endosymbiotic algae or cyanobacteria (Schmid 1995). Distribution: The Tubiphytes group is a cosmopolitan reef builder ranging from the Carboniferous to the Cretaceous (Fig. 10.61) and abundant in the Permian, Middle Triassic and Late Jurassic.
Fig. 10.61. Distribution and frequency of fossils attributed to Tubiphytes. Frequency is expressed by the relative number of publications dealing with the taxa. After SenowbariDaryan and Flügel (1993).
Microproblematica: Tubiphytes
Permian Tubiphytes are encrusters, binders, bafflers and frame builders in reefs (LeMone 1995; Wahlman 2002; Weidlich 2002) but are also known from bedded shelf carbonates. Tubiphytes occurs in shallow subtidal (and intertidal) settings as well as in submarine fissures indicating relatively large tolerance levels with regard to light, salinity and substrate. Tubiphytes is commonly associated with Archaeolithoporella, sponges and bryozoans. The group flourished in warm-water tropical settings but also lived in deeper shelf and slope environments as well as in cooler settings. Triassic Tubiphytes are major constituents of Middle Triassic reefs but minor elements in Late Triassic reefs. Late Jurassic and Early Cretaceous Tubiphytes occurred in different environments as compared with Paleozoic Tubiphytes, including shallow outer ramp environments where thick specimens lived, and deepermarine and cryptic habitats characterized by thinner specimens (Schmid 1996). Various growth forms reflect different water energies (Kott 1989). Intergrown Tubiphytes morronensis associated with Terebella (see Pl. 92/10) is abundant in Late Jurassic and Early Cretaceous proximal platform talus and slope deposits as well as in siliceous sponge ‘reefs’ formed in deeper ramp settings. Paleowater depths of about 100 to 200 m have been considered for mass occurrences. The species also contributed to the formation of organic frameworks (Pomoni-Papaioannou et al. 1989). Interpretation: Tubiphytes has been attributed to various algal groups, cyanobacteria as well as sponges, hydrozoans and foraminifera (Riding and Guo 1992). The current explanation for Paleozoic and some Triassic Tubiphytes is a symbiotic commensalism of different organisms recorded by the spar-filled ‘cavities’ and tubes and calcified cyanobacteria represented by the surrounding network. Jurassic Tubiphytes are explained as foraminifera. Basics: Microproblematica Babcock, J.A. (1986): The puzzle of alga-like problematica, or rummaging around in the algal wastebasket. – In: Hofman, A., Nitecki, M.H. (eds.): Problematic fossil taxa. – Oxford Monographs on Geology and Geophysics. – 12-26, New York (Oxford University Press) Banner, F.T., Finch, E.M., Simmons, M.D. (1990): On Lithocodium Elliott (calcareous green algae): its paleobiological and stratigraphical significance. – Journal of Micropaleontology, 9, 21-36 Bengtson, S. (1990): Problematic fossil taxa. – In: Briggs, D.E., Crowther, P.R. (eds.): Palaeobiology. A synthesis. – 442-445, Oxford (Blackwell) Cherchi, A., Schroeder, R. (1988): Osservazioni sui microproblematica Paronipora Capeder, Microcodium Glück, Baccanella Pantic et Paleomicrocodium Mamet and Roux. – Bolletino della Società Paleontologica Italiana, 27, 7981
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De Castro, P. (1990): Thaumatoporelle: conoscenze attuale a appoccio all’interpretazione. – Bolletino della Società Paleontologica Italiana, 29, 176-206 Elliott, G.F. (1958): Fossil microproblematica from the Middle East. – Micropaleontology, 4, 419-428 Elliott, G.F. (1963): Problematical microfossils from the Cretaceous and Paleocene of the Middle East. – Palaeontology, 6, 293-200 Flügel, E. (1964): Mikroproblematika aus rhätischen Riffkalken der Nordalpen. – Paläontologische Zeitschrift, 38, 74-87 Flügel, E. (1972): Mikroproblematika in Dünnschliffen von Triaskalken. – Mitteilungen der Gesellschaft der Geologieund Bergbaustudenten in Österreich, 21, 857-988 Glück, H. (1912): Eine neue gesteinsbildende Siphonee (Codiacee) aus dem marinen Tertiär von Süddeutschland. – Mitteilungen der badischen Geologischen Landesanstalt, 4, 3-24 Henssel, K., Schmid, D.U., Leinfelder, R. (2002): Computergestützte 3D-Rekonstruktionen in der Paläontologie an Hand von Serienschnitten. – Mathematische Geologie, 6, 131-142 Hoffman, A., Nitecki, M.H. (eds., 1986): Problematic fossil taxa. – Oxford Monographs on Geology and Geophysics 5, 267 pp., New York (Oxford University Press) Klappa, C.F. (1978): Biolithogenesis of Microcodium: elucidation. – Sedimentology, 25, 489-522 Kobluk, D.R. (1988): Cryptic faunas in reefs: ecologic and geologic importance. – Palaios, 3, 379-390 Munnecke, A. (2001): Halysis Hoeg, 1932 - a problematic Cyanophyceae: new evidence from the Silurian of Gotland (Sweden). – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 2001, 21-42 Nassichuk, W.W., Davies, G.R., Mamet, B.L. (1986): Microcodiaceans in the Viséan Emma Fjord Formation, Devon Island, Arctic Canada. – Current Research, part B, Geological Survey of Canada, Paper 86-1B, 467-470 Perret, M.F., Vachard, D., Aguirre, P., Crasquin-Soleau, S. (1994): Micropaléontologie des calcaires epibathyaux a Globochaete (algue problématique) du Carbonifère des Pyrénées. – Geobios, 27, 659-675 Riding, R., Guo, L. (1992): Affinity of Tubiphytes. – Palaeontology, 35, 37-49 Schmid, D.U.: (1995): ‘Tubiphytes’ morronensis - eine fakultativ inkrustierende Foraminifere mit endosymbiontischen Algen. – Profil, 8, 305-317 Schmid, D.U. (1996): Marine Mikrobolithe und Mikroinkrustierer aus dem Oberjura. – Profil, 9, 101-251 Schmid, D.U., Leinfelder, R. (1996): The Jurassic Lithocodium aggregatum-Troglotella incrustans foraminiferal consortium. – Palaeontology, 39, 21-52 Senowbari-Daryan, B. (1984): Mikroproblematika aus den obertriadischen Riffkalken von Sizilien. – Münstersche Forschungen zur Geologie und Paläontologie, 61, 1-81 Senowbari-Daryan, B., Flügel, E. (1993): Tubiphytes Maslov, an enigmatic fossil: Classification, fossil record and significance through time. Part I: Discussion of Late Palaeozoic material. – In: Barattolo, F., DeCastro, P., Parente, M. (eds.): Studies on fossil benthic algae. – Bolletino della Società Paleontologica Italiana, Special Volume, 1, 353-382 Skompski, P. (1982): The nature and systematic position of the microfossil Globochaete alpina Lombard, 1945. – Acta Geologica Polonica, 32, 47-56 Vennin, E., Vachard, D., Proust, J.-N. (1997): Taphonomie et synécologie du ‘genre’ Tubiphytes dans les biocon-
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structions de Tratau et de Nizhni-Irginsk (Permien inférieur de l’Oural, Russie). – Geobios, 30, 635-649 Further reading: K138 (general), K139 (Paleozoic), K140 (Mesozoic), K141 (Cenozoic)
10.3 Biozonation of Platforms with Thin-Section Fossils The section discusses the potential of microfacies in subdividing platforms sections that are generally poor in guide fossils into biozones. The example used is from the Late Mesozoic, but similar problems exist for Paleozoic, Early Mesozoic and Tertiary platforms. Starting in the 60s of the last century many attempts have been made in subdividing the thick Jurassic and Cretaceous carbonate platform sequences in the circumMediterranean and Alpine regions and in stratigraphically using thin-section fossils, predominantly calcareous algae and benthic foraminifera. Sartoni and Crescenti (1962) studying Mesozoic platforms in the southern Apennines introduced the concept of ‘cenozones’. The authors defined a cenozone as strata containing a characteristic association of microfossils. Later authors combined characteristic association with conspicuous abundance of specific taxa. Doing this, the definition of cenozones and the practical use of these units include both characteristics of assemblage zones as well as abundance zones. Characteristics of assemblage zones and abundance zones are enclosed in the modern definition of biozones which are determined by the stratigraphic and geographic distribution of (1) a single taxon, (2) an diagnostic association, or (3) the relative abundance of taxa (Salvador 1994). In earlier studies, zones have been set up for Jurassic and Early Cretaceous platforms, predominantly based on dasyclad algae and associated benthic foraminifers: Plate 62 and Pl. 69 display dasyclad and foraminiferal taxa used in defining Jurassic biozones. Pl. 63, Pl. 70 and Pl. 71 show taxa characterizing Early Cretaceous zones. The zonations developed in the Apennines were compared with biozones recognized in the Northern Calcareous Alps, the Dinarides, eastern and southern Carpathians, Switzerland, Spain, South France and Greece (see basics and references in K142 and K144). Jaffrezo (1980) tried to correlate the local biozones and suggested a general biozonation of Jurassic and Early Cretaceous platforms in the western Tethys. Pitfalls. The cenozone approach was widely used until about 1990 despite serious pitfalls causing difficulties in correlating and determining the chronostratigraphic age. The age of the biozones and the bio-
Biozonation of Platform Carbonates
stratigraphic ranges of benthic shallow-water organisms must be calibrated with global pelagic biozones (based, e.g., on cephalopods, calpionellids, radiolarians). Isolated skeletal grains, reworked on platforms, transported to basinal environments and deposited within calciturbidites may aid in this calibration. Care must be taken in using microfossils enclosed in reworked platform lithoclasts because this material might represent older, non-synchronous sediment. Using the cenozone concept must consider the following points: • Environmental control. Dasyclad algae and benthic foraminifera are strongly controlled by environmental conditions. Biostratigraphic interpretations based on dasyclads generally suffer from an apparently slow evolution rate, facies-dependence, and poor knowledge of the stratigraphic range of many taxa. • Environmental conditions (e.g. substrate, nutrients, water energy) are different in different parts of carbonate platforms or ramps, and different zonation patterns may be developed in restricted inner platform settings (e.g. shelf lagoons), open-marine parts of platforms and at platform margins (Crescenti 1971; De Castro 1987; Dya 1992). • Mesozoic and Cenozoic dasyclads and also benthic foraminifera exhibit strong provincialism. Biozonations established in one paleogeographical province may not be found again in another province. Future developments. Many current microfacies studies of platform carbonates devoid establishing biozones within a stratigraphic framework and use the range of single taxa in deliminating stratigraphic units. However, biozones based on occurrence, association and abundance of algae and foraminifera provide still an important tool in microfacies analysis of carbonate platforms if the definition of formal biozones considers: • Depositional environment controls. Sections used in studying changes in the composition of algal and foraminiferal assemblages must be differentiated with regard to depositional setting (restricted or open platform, platform margin). • Provincialism of the fossil used in establishing biozones demands rigorous statistical tests indicating which geographical areas can be included within one paleogeographical site and which not. The Jaccard similarity index has been found very useful in comparing Late Jurassic dasyclads from various localities in the Northern Calcareous Alps, and recognizing high or low similarities (Rasser and Fenninger 2002). • Only paleogeographical sites having high similarity indices should be used in correlating biozones within a chronostratigraphic framework. In the case of the Late Jurassic platforms of the Northern Calcareous Alps, the
Biozonation of Platform Carbonates
Carpathians, the Apuseni Mountains in Romania and the Central Apennines form one region in which the platforms appear to exhibit identical or similar biozonation sequences. Basics: Biozonation of platform carbonates based on thinsection fossils Azema, J, Chabrier, G., Fourcade, E., Jaffrezo M. (1977): Nouvelles données micropaléontologiques stratigraphiques et paléogéographiques sur le Portlandien et le Néocomian de Sardaigne. – Revue des Micropaléontologie, 20, 125-139 Arnaud-Vanneau, A. (1980): Micropaléontologie, palécologie et sédimentologie d’un platforme carbonatée de la marge passive de la Théthys: l’Urgonian du Vercors septentrional et de la Chatreuse (Alpes occidentales). – Géologie Alpine, Mémoire, 11, part 1, 267 pp., part 2, 874 pp., part 3, 119 pp. Bodrogi. I., Bona, J., Lobitzer, H. (1994): Vergleichende Untersuchung der Foraminiferen- und Kalkalgen-Assoziationen der Urgon-Entwicklung des Schrattenkalkes in Vorarlberg (Österreich) und der Nagyharssány Kalkstein Formation des Vilány-Gebirges (Ungarn). – Jubiläumsschrift 20 Jahre Geologische Zusammenarbeit ÖsterreichUngarn, Geologische Bundesanstalt Wien, Teil 2, 225-283 De Castro, P. (1987): Le facies di piattaforma carbonatica del Giurassico italiano: diffusione areale e lineamente biostratigrafici. – Bolletino della Società Paleontologica Italiana, 26, 309-325 Dragastan, O. (1975): Upper Jurassic and Lower Cretaceous microfacies from the Bicaz Valley Basin (East Carpathians). – Institut de Géologie et Géophysics Bucuresti, Memoir, 21, 87 pp. Farinacci, A., Radoicic, R. (1964): Correlazione tra serie giurese e cretacee del’Apennino centrale e Dinaridi esterne. - Ricerche Scientifiche, Anno 34 (IIA), 169-300 Garcia-Hernandez, M. (1981): Biozonation du Cretacé inférieur a l’aide des foraminiféres benthiques et algues dasycladacés dand le Prebetique occidental (Cordilleres Betiques, Espagne). – Geobios, 14, 261-267 Jaffrezo, M. (1980): Les formations carbonatées des Corbiéres (France du Dogger a l’Aptien: micropaléontologie stratigraphique, biozonation, palécologie. Extension des resultats a la Meésogé. – Thése Doctorat d’Etat, Université Pierre et Marie Curie, Paris VI, 613 pp. Rasser, M.W., Fenninger, A. (2002): Jurassic/Cretaceous Dasycladalean algal stratigraphy in the Northern Calcareous Alps: a critical review and a paleobiogeographcal approach using similarity indices. – In: Bucur, I.U., Filipescu, S. (eds.): Research advances in calcareous algae and microbial carbonates. - Proceedings of the 4th IFAA Regional Meeting, Cluj-Napoca 2001, 167-190 Salvador, A. (ed., 1994): International stratigraphic guide. – 214 pp., Boulder (Geological Society of America) Sartorio, D., Venturini, S. (1988): Southern Tethys biofacies. – 235 pp., San Donato Milanese (Agip) Sartorio, S., Crescenti, U. (1962): Ricerche biostratigrafiche nel Mesozoico dell’Appennino meridionale. – Giornale di Geologia, 29, 161-304 Schlagintweit, F., Ebli, O. (1999): New results on microfacies, biostratigraphy and sedimentology of the Late JurassicEarly Cretaceous platform carbonates of the Northern
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Calcareous Alps. – Abhandlungen der Geologischen Bundesanstalt Wien, 56, 379-418 Schröder, R., Neumann, M. (1985): Les Grand Foraminiféres du Cretacé moyen de la région méditerranne. – Geobios, Mémoire spécial, 7, 161 pp. Septfontaine, M. (1980): Les foraminifères imperforès des mileux des mileux plate-forme au Mésozoique: détermination pratique, interprétation phylogénetique et utilisation biostratigraphique. – Revue de Micropaléontologie, 23, 169-203 Tisljar, J., Velic, I. (1993): Upper Jurassic (Malm) shallowwater carbonates in western Gorski Kotar area: facies and depositional environments (western Croatia). – Geologia Croatica, 46, 263-279 Further reading: K111, K129, K142, K144
10.4 Where to Look for Thin-Section Pictures of Fossils and Microfacies Types A good picture may say more than many words. Therefore, the survey given in Sect. 10.2 on the various groups of thin-section fossils is supplemented in a more practical way by the following list, which should assist in finding appropriate thin-section photographs. Sect. 10.4 presents a bibliography of books and papers that include good and instructive photographs (at least 10 plates) of skeletal grains and common microfacies types in Phanerozoic carbonates. Not included are specific publications on the major groups of thin-section fossils (foraminifera, calcareous algae). The references are marked by letters that characterize the main topics dealt with: BZ - biozonations based on thin-section fossils, E - environmental conditions, F - facies characteristics and sedimentological criteria, FM - facies models, P - valuable paleontological informations, PD - paleontological determinations that can be used in the generic and/or specific assignment of thin-section fossils. Books covering several time units Bozorgnia, F., Banafti, S. (1964): Microfacies and microorganisms of Paleozoic through Tertiary sediments of some parts of Iran. – 22 pp., 158 Pls., Tehran (Iranian Oil Company). Cambrian, Devonian to Tertiary (Miocene). PD Chiocchini, M., Mancinelli, A. (1977): Microbiostratigrafia del Mesozoico in facies di piattaforma carbonatica dei Monti Aurunci (Lazio Meridionale). – Studi Geol. Camerti, 3, 109-152, 48 Pls., Camerino. Late Triassic to Miocene. BZ, P Cita, M.B. (1965): Jurassic, Cretaceous and Tertiary microfacies from the Southern Alps (Northern Italy). – International Sedimentary Petrographical Series, 8, 99 pp., 117 Pls., Leiden (Brill). Early Jurassic to Late Tertiary (Miocene). PD Christodoulou, G., Tsaila-Monopolis, St. (1975): Eastern Hellenic zone microfacies. – National Institute of Geological
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and Mining Research, 17, 63 pp., 75 Pls., Athens. Early Triassic to Early Tertiary. P Cuvillier, J., Sacal, V. (1951): Corrélations stratigraphiques par microfacies en Aquitaine Occidental. – International Sedimentary Petrographical Series, 2, 16 pp., 90 Pls., Leiden (Brill). Jurassic to Tertiary. P Hagn, H. (1955): Fazies und Mikrofauna der Gesteine der Bayerischen Alpen. – International Sedimentary Petrographical Series, 1, 20 pp., 71 Pls., Leiden (Brill). Triassic to Tertiary. P Hanzawa, S. (1961): Facies and micro-organisms of the Paleozoic, Mesozoic and Cenozoic sediments of Japan and her adjacent islands. – International Sedimentary and Petrographical Series, 5, 117 pp., 148 Pls., Leiden (Brill). Silurian to Cenozoic, PD Lehmann, E.P. et al. (1957): Microfacies of Libya. – 80 pp., 37 Pls., Tripoli (Petroleum Exploration Society of Libya). Cambrian to Miocene. P Mehrnush, M., Partoazar, H. (1977): Selected microfauna of Iran. – Geol. Min. Survey Iran Report, 33, 396 pp., 117 Pls., Tehran. Late Devonian to Miocene. PD Microfacies italiane (1959): Microfacies italiane (dal Carbonifero al Miocene medio). – 35 pp., 145 Pls., Milano (Agip Mineraria). Late Carboniferous to Miocene. P. Misik, M. (1966): Microfacies of the Mesozoic and Tertiary limestones of the West Carpathians. – 269 pp., 191 Pls., Bratislava (Slovenska Akademia Vied). Early Triassic to Miocene. P, F Papp, A., Turnovsky, K. (1970): Anleitung zur biostratigraphischen Auswertung von Gesteinsdünnschliffen (Microfacies Austriaca). – Jahrbuch der Geologischen Bundesanstalt, Sonderband, 16, 50 pp., 88 Pls., Wien. Late Carboniferous to Miocene. P Rey, M., Nouet, G. (1958): Microfacies de la région prérifaine et de la moyenne Moulouya (Maroc Septentrional). – International Sedimentary Petrographical Series, 3, 41 pp., 97 Pls., Leiden (Brill). Early Jurassic to Miocene. P Sampo, M. (1969): Microfacies and microfossils of the Zagros area, southwestern Iran (from Permian to Miocene). – International Sedimentary Petrography Series, 12, 102 pp., 105 Pls., Leiden (Brill). Early Permian to Miocene. PD Sartoni, S., Crescenti, U. (1962): Richerche biostratigrafiche nel Mesozoico del Apennino meridionale. – Giornale di Geologia, 29, 161-304, 52 Pls., Bologna. Early Jurassic to Paleocene. BZ, PD Sartorio, D., Venturini, S. (1988): Southern Tethys biofacies. – 235 pp., numerous microphotographs, Milano (Agip Mineraria). Cambrian to Tertiary (Pliocene), predominantly Mesozoic and Tertiary. PD, E Tsaila-Monopolis, St. (1977): Micropaleontological and stratigraphical study of the Tripolitza (Gavrovo) zone in the Peloponnesus. – Inst. Geol. Mining Res. Geol. Geophys. Res., 20, 106 pp., 56 Pls., Athens. Late Triassic to Early Tertiary. P Paleozoic Carozzi, A.V., Textoris, D.A. (1967): Paleozoic carbonate microfacies of the Eastern Stable Interior (U.S.A.). – International Sedimentary Petrographical Series, 11, 41 pp., 100 Pls., Leiden (Brill). Ordovician to Late Carboniferous. F Glintzboeckel, C., Rabate, J. (1964): Microfaune et microfacies du Permo-Carbonifère de sud Tunisien. – International Sedimentary Petrographical Series, 7, 45 pp., 108 Pls.,
Pictures of Thin-Section Fossils
Leiden (Brill). Middle Carboniferous to Late Permian. P Kreutzer, L.H. (1992): Photoatlas zu den variszischen Karbonat-Gesteinen der Karnischen Alpen (Österreich/Italien). – Abhandlungen Geologische Bundesanstalt Wien, 47, 1-129, 46 Pls. Ordovician to Carboniferous. E, F, P Mamet, B.L. (1970): Carbonate microfacies of the Windsor Group (Carboniferous), Nova Scotia and New Brunswick. – Geological Survey of Canada Bulletin, 70-21, 121 pp., 19 Pls., Ottawa. – Early Carboniferous. P Mamet, B. (1975): An atlas of microfacies in Carboniferous carbonates of the Canadian Cordillera. – Bulletin of the Canadian Geological Survey, 255, 1-131, 95 Pls., Ottawa. Carboniferous. E, P Sacal, V., Cuvillier, J. (1964): Microfaciès du Paléozoique Saharien. – Compagnie Francaise du Petrol, Notes et memoirs, 6, 1-30, 100 Pls., Paris. Cambrian to Carboniferous. P Tellez-Giron, C. (1970): Microfacies y microfosiles paleozoicos del area ciudad Victoria. – Instituto Mexicano Petroleo, Serie Monografia, 1, 127 pp., 57 Pls., Mexico. Silurian, Carboniferous and Permian. F Vachard, D. (1980): Tethys and Gondwana au Paléozoique supérieur les données afghanes. Biostratigraphie, micropaléontologie, paléogéographie. – Documents et Travaux IGAL, 2, 463 pp., 35 Pls., Paris. Late Carboniferous and Permian. E, F, P Zamarreno, I. (1972): Las lithofacies carbonatadas del Cambrico de la zona cantabrica (NW Espana) y su distribucion paleogeografia. – Faculdad de Ciencias, Universidad Oviedo, 102 pp., 17 Pls., Oviedo. Cambrian. F Mesozoic: Triassic Bissell, H.J. (1970): Petrology and petrography of Lower Triassic marine carbonates of Southern Nevada (U.S.A.). – International Sedimentary Petrographical Series, 14, 27 pp., 82 Pls., Leiden (Brill). Early Triassic. F Gazdzicki, A. (1974): Rhaetian microfacies, stratigraphy and facial development in the Tatra Mts. – Acta Geologica Polonica, 24, 17-96, Pl. 29-52, Warszawa. Late Triassic. P Pantic-Prodanovic, S. (1975): Les microfaciès triassiques des Dinarides. Le Montenegro, la Bosnie occidentale et Herzegovine et Serbie occidentale. – Societé des Sciences et des Arts Montenegro, Classe des Sciences Naturelles, Monograph, 4, 257 pp., 100 Pls., Titograd.Triassic. PD Sudar, M. (1986): Triassic microfossils and biostratigraphy of the Inner Dinarides between Gucevo and Ljubisna Mts., Yugoslavia. – Annales Géologiques de la Péninsula Balkanique, 50, 151-394, 30 Pls., Beograd. Triassic. E, PD Mesozoic: Triassic and Jurassic Fabricius, F.H. (1966): Beckensedimentation und Riffentwicklung an der Wende Trias/Jura in den Bayerisch-tiroler Kalkalpen. – International Sedimentary Petrographical Series, 9, 143 pp., 27 Pls., Leiden (Brill). Late Triassic and Early Jurassic. E, F Perconig, E. (1968): Microfacies of the Triassic and Jurassic sediments of Spain. – International Sedimentary Petrographical Series, 10, 63 pp., 123 Pls., Leiden (Brill). Triassic and Jurassic. F, PD Mesozoic: Jurassic Carozzi, A.V., Bourroullec, J., Deloffre, R., Rumeau, J.L. (1972): Microfaciès du jurassique d’Aquitaine. – Bulletin Centre de Recherche Pau. SNPA, Volume Spécial, 1,
Pictures of Thin-Section Fossils
594 pp., 200 Pls., Pau Early to Late Jurassic. E, F, FM Derin, B., Reiss, Z. (1966): Jurassic microfacies of Israel. – Israel Institute of Petroleum, Special Publication, 43 pp., 320 microphotographs, Jerusalem. Early to Late Jurassic. P Farinacci, A. (1959): Le microbiofacies giurassiche dei Monti Martani (Umbria). - Pubblicazione dell’ Istituto Geologico Università Roma, 8, 1-38, 17 Pls., Roma. Jurassic. BZ, Mesozoic: Jurassic and Cretaceous Barattolo, F., Pugliese, A. (1987): Il Mesozoico dell’Isola di Capri. – Quaderni dell’Academia Pontiana, 8, 1-172, 66 Pls., Napoli. Early Jurassic to Late Cretaceous. F, P Borza, K. (1969). Die Mikrofazies und Mikrofossilien des Oberjura and der Unterkreide der Klippenzone der Westkarpaten. – 301 pp., 88 Pls., Bratislava (Slowakische Akademie der Wissenschaften). Middle Jurassic to Albian. P. Bucur, I.I. (1997): Formatiunile mesozoice din zona ResitaMoldova-Noua. – 214 pp., 32 Pls., Cluj (University Press). Carras, N. (1995): La piattaforma carbonatica del Parnasso durante di Giurassico superiore-Cretacico inferiore (stratigrafia ed evoluzione paleogeografica). – 232 pp., 69 Pls., Athens. Middle Jurassic to Aptian. BZ, F Dragastan, O. (1975): Upper Jurassic and Lower Cretaceous microfacies from the Bicaz Valley Basin (East Carpathian). Institut de Géologie et de Géophysique, Mémoires, 21, 187, 95 Pls., Bucurest. Late Jurassic to Aptian. BZ, P Farinacci, A., Radoicic, R. (1964): Correlazione tra serie giurese e cretacee dell’Àppennino centrale e delle Dinaridi esterne. – Ricerche Scientifiche Anno 34 (II A), 7, 269300, 15 Pls., Roma. Early Jurassic to Maastrichtian. BZ Radoicic, R. (1966): Microfaciès du Jurassique des Dinarides externes de la Yougoslavie. – Geologija, 9, 377 pp., 165 Pls., Ljubljana Ramalhao. M.M. (1971): Contribution à l’étude micropaléontologique et stratigraphique du Jurassique supérieur et du Crétacée inferieur des environs de Lisbonne (Portugal). – Servicos Geologicos Portugal, Memoir, 19, 1-212, 39 Pls., Lisboa. Late Jurassic to Valanginian. F, P Velic, I. (1977): Jurassic and Cretaceous assemblages zones in Mt. Velika Kapela, Central Croatia. – Acta Geologica, 9, 15-37, 32 Pls., Zagreb. Mesozoic: Cretaceous Arnaud-Vanneau, A. (1980): Micropaléontologie, paléoécologie et sédimentologie d’une plate-forme carbonatée de la marge passive de la Téthys. L’Urgonien du Vercors septentrional et de la Chartreuse (Alpes occidentales). – Géologie Alpine, Mémoire, 11, 1: 267 pp., 2: 874 pp., 3: 119 pp., 115 Pls., Grenoble. Early Cretaceous. E, PD Berthou, P.Y. (1973): Le Cénomianian de l’Estrémadura portugaise. – Memoir Servizio Geologico Portugal, 23, 169 pp., 67 Pls., Lisboa. Late Cretaceous. P Berthou, P.Y., Bengtson, P. (1988): Stratigraphic correlation by microfacies of the Cenomanian-Coniacian of the Sergipe Basin, Brazil. – Fossil and Strata, 21, 88 pp., 48 Pls., Oslo. Late Cretaceous. BZ, F Bilotte, M. (1984): Le Crétacée supérieur des plate-formes est-Pyrénéennes. Atlas. – Strata, 1, 45 Pls., Toulouse. Late Cretaceous. P Bilotte, M. (1985): Le Crétacée supérieur des plate-formes est-Pyrénéennes. Atlas. – Strata, 1, 1-438, Toulouse. Late Cretaceous. PD Cestari, R., Sartorio, D. (1995): Rudists and facies of the Peri-
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adriatic domain. – 207 Pls., many microphotographs, Milano (Agip). Early and Late Cretaceous. E, F, PD Ford, N., Houbolt, J.J. (1963): The microfacies of the Cretaceous of Western Venezuela. – International Sedimentary Petrographical Series, 6, 52 pp., 55 microphotographs, Leiden (Brill). Early and Late Cretaceous. FM Gusic, I., Jelaska, V. (1990): Upper Cretaceous stratigraphy of the island of Brac with the geodynamic evolution of the Adriatic plarform. – Djela Jugoslavenske Akademije Znanosti i Umetnosti, 69, 160 pp., 20 Pls., Zagreb. Late Cretaceous. F, FM Hamaoui, M., Saint-Marc, P. (1970): Microfaunes et microfaciès du Cénomanien du Proche Orient. – Bulletin Centre des Rechereche Pau-SNPA, 4, 357-352, 21 Pls., Pau. Cenomanian. E, P Höfling, R. (1985): Faziesverteilung und Fossilvergesellschaftungen im karbonatischen Flachwasser-Milieu der alpinen Oberkreide (Gosau-Formation). – Münchner Geowissenschaftliche Abhandlungen, A, 3, 1-241, 10 Pls., München. Late Cretaceous. E, F, P Luperto Sinni, E. (1976): Microfossili senoniani delle Mnurge. – Rivista Italiana della Paleontologia e Stratigrafia, 82, 293-416, 59 Pls., Milano. Late Cretaceous. P Mesozoic, Cenozoic: Cretaceous and Early Tertiary Bignot, G. (1972): Recherches stratigraphiques sur les calcaires du Crétacé supérieur et de l’Éocene d’Istrie et des régions voisines. Essai de révision du Liburnien. – Thèse Travaux Laboratoire de Micropaléontologie Université Paris VI, 2, 353 pp., 50 Pls., Paris. Late Cretaceous and Eocene. PD Radoicic, R. (1960): Microfaciès du Crétacé et du Paléogene des Dinarides externes de la Yugoslavie. – Inst. Rech. R.P. Crna Gora, Geol. Ser., 4, 172 pp., 67 Pls., Titograd. Early Cretaceous to Eocene. P Samuel, O., Borza, K., Köhler, E. (1972): Microfauna and lithostratigraphy of the Paleogene and adjacent Cretaceous of the Middle Vah Valley (West Carpathians). – 246 pp., 180 Pls., Bratislava (Geologicky ustav Dionyza Stura). Late Cretaceous and Early Tertiary. F, PD Cenozoic Bonnefous, J., Bismuth, H. (1982): Les faciès carbonatées de plate-forme de l’Éocene moyen et supérieur dans l’ offshore tunisien nord-oriental et en Mer Pélagienne: complications paléogéographiques et analyse micropaléontologique. – Bulletin de la Centre des Recherches Pau SNPA, 6, 337-403, 16 Pls., Pau. Early Tertiary (Eocene). FM, P Carozzi, A.V., Reyers, M.V., Ocampo, V.P. (1976): Microfacies and microfossils of the Miocene reef carbonates of the Philippines. – Philippine Oil Development Company, Special Volume, 1, 79 pp., 20 microphotographs, Manila. Late Tertiary. E, FM, P Herbig, H.-G.(1991): Das Paläogen am Südrand des zentralen Hohen Atlas und im Mittleren Atlas Marokkos. Stratigraphie, Fazies, Paläogeographie und Paläotektonik. – Berliner geowissenschaftliche Abhandlungen, A, 135, 285 pp., 40 Pls.. Paleogene. F, FM, PD Massieux, M. (1973): Micropaléontologie stratigraphique de l’Éocene des Corbières septentrionales (Aude). – Cahiers de Paléontologie, Centre national de recherches scientifique, 139 pp., 29 Pls., Paris. Eocene. P Further reading: K144
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Fig. 10.62. Exercise: Which fossils are shown in this figure? For answers see Appendix.
Exercise: Fossils
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MICROFACIES INTERPRETATION The second part of this book explains the characteristics of microfacies types (Chap. 11), discusses diagnostic features of paleoenvironmental conditions (Chap. 12) and the importance of integrated facies analysis (Chap. 13), examines the potential of generalized facies models and standard microfacies (Chap. 14), and illustrates the significance of microfacies in the recognition of depositional settings (Chap. 15) and specific sedimentation patterns (Chap. 16).
11 Summarizing Microfacies Criteria: Microfacies Types
Microfacies studies aim for the recognition of overall patterns that reflect the history of carbonate rocks, by means of a thorough examination of their sedimentological and paleontological characteristics. The previous chapters have shown which thin-section criteria are used in describing the microfacies of limestones. The evaluation of microfacies in the context of facies interpretation requires a synopsis of microfacies data observed in various samples into a microfacies type (MFT). Microfacies types and facies associations are fundamental to the development of models for carbonate sedimentation.
11.1 MFT Concepts Microfacies, microfacies types and standard microfacies types The microfacies recorded by a thin section is not necessarily identical to a microfacies type. The approach to microfacies is largely descriptive, and microfacies types should be defined by those microfacies criteria whose existence and abundance are determined by specific environmental factors and are linked to specific depositional settings. This requires careful selection of the textural and compositional criteria used. Basic prerequisites for defining MFT are a clearcut discrimination of grain categories discussed in Sect. 4.2, an understanding of limestone classifications based on textural criteria (Sect. 8.3), the recognition of depo-
sitional fabrics (Chap. 5) and the ability to attribute thinsection fossils to major systematic groups and taxonomic units. Identical biota and comparable depositional environments give rise tof time-specific microfacies types, some of which have been generalized, for example the facies types developed for Devonian shelf and reef carbonates (Hladil 1986; Machel and Hunter 1994). Similar or identical criteria in individual microfacies types from limestones not restricted to specific time intervals are used to define standard microfacies types (SMF types), which describe major depositional and biological controls and suggest major depositional settings (Wilson 1975). The SMF concept and its application in reconstructing depositional systems are discussed in Sect. 14.3. Microfacies types, microfacies variability and facies hierarchy Texture, as well as qualitative and quantitative microfacies characteristics, may vary considerably within a limestone bed, both laterally and vertically. Variability depends partly on sample size and sampling procedures (see Sect. 3.1.2.1), but may be real and reflect small-scaled depositional events or rapidly fluctuating environmental conditions. Several MFT may occur within a single thin section (Morrow and Webster 1991; see Pl. 127) or in samples taken within lateral distances of only a few tens of centimeters (Egenhoff et al. 1999). The evaluation of microfacies variability requires pre-
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_11, © Springer-Verlag Berlin Heidelberg 2010
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cise and reproducible definitions of MFT, but the question then arises of which rank to ascribe to MFTs within a hierarchically structured facies concept (Aigner 1984; Hüssner 1985)? Microfacies types for allochthonous carbonates, defined by texture and composition and based on smallscale analyses (e.g. bed-by-bed studies), define the fundamental depositional unit and reveal the depositional dynamics that were operating during transport and deposition (e.g. storms). Facies analyses based on depositional lithofacies (e.g. wackestone, grainstone) and the comparison of MFTs from sedimentary sequences record both vertical and lateral changes. Vertical changes may indicate shallowing-up or deepening-up and reveal facies dynamics (e.g. regressive/transgressive events). Lateral changes (e.g. following the dip of a ramp) may indicate differences in water depths and hydrodynamics. Associations of MFTs occurring within the same lithofacies and deposited in the same general environment (e.g. MFTs of grainstones) describe local sedimentary subenvironments (e.g. soft and firm substrates) or local processes.
11.2 How to Differentiate Meaningful Microfacies Types Microfacies types for autochthonous limestones are defined by different criteria from those for other limestones Autochthonous limestones (Sect. 8.2) include both reef limestones and microbial carbonates. Reef limestones: MFTs are differentiated according to the type of reef-building fossils. Generally, the most common reef-building taxa in combination with the dominant texture (framestone, bafflestone, bindstone or boundstone) and matrix type are used in defining a MFT. The example shown in Fig. 10.36 (subsequently called sample A) is a sponge boundstone, or more precisely a coralline-sponge boundstone. Because the latter designation does not reflect the diversity of the coralline sponges contributing to the boundstone fabric, the MFT would be better characterized as ‘mediumdiverse coralline boundstone’. The sample depicted on Pl. 81/1 (sample B) yields the same coralline sponge groups and corresponds to sample A in an inhomogeneous micrite matrix. The two samples represent the same microfacies type even though sample B exhibits spongiomorphids not shown in A. This additional constituent is of minor importance for the definition of the
Differentiation of Microfacies Types
MFT because its consideration does not change the environmental interpretation of the two samples. Both samples reflect the existence of an organic reef framework formed by coralline sponges within low-energy environments. However, the depositional setting of each of these environments is different. Both samples are Late Triassic in age. Sample A comes from small patch reefs in the marly-calcareous Zlambach beds of the Northern Calcareous Alps which were formed in front of the large platform margin reefs represented by sample B which comes from the Dachstein limestones. The similarity in biotic composition and texture does not express congruent depositional sites, but does express similar nutrient controls in habitats that were comparable with regard to the ecological constraints of nutrient supply, available substrate and reduced water energy. Another sponge limestone is shown in Fig. 10.34. This Late Jurassic sample differs distinctly from the sponge limestones discussed above in the sponge group (siliceous hexactinellid sponges), the association of sponges with microbial crusts at the top of the sponges, and the spotty matrix. The sample represents a specific MFT characterized essentially by biotic composition. The spotty and peloidal texture of the matrix indicates microbial controls. Other Late Jurassic sponge limestones may differ from the sample shown in growth forms of the sponges, the absence of microbial crusts, or matrix texture. These differences define specific MFTs that reflect differences in ecological and depositional constraints. Some criteria are of minor or no relevance for the characterization of the MFT. The existence of brachiopods (figure center) should be taken into consideration if the group is a conspicuous constituent of the limestone. Fracturing and differential compaction are not diagnostic criteria for definition of the MFT. Plate 88 exhibits Cretaceous limestones with rudists. The term rudist limestone would be too general for establishing a defined MFT because different rudist groups prefer different habitats and environments and occur in different depositional settings. The definition of a MFT must take into consideration the major taxonomic groups (which can be recognized in thin section by specific microstructural patterns), their diversity (e.g. mono- or multispecific) and their association with other biota. Evaluation of diversity may be difficult or misleading on microfacies scale for this reef limestone, as well as for other reef carbonates, without knowledge of outcrop and field data. Many reef limestones require a two-fold typification as demonstrated by the Figs. 11.1 (sample A) and 11.2 (sample B) which show Late Triassic coral lime-
Microfacies Types of Reef Limestones
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Fig. 11.1. Coral limestone with Retiophyllia clathrata (Emmrich). Coral thicket formed on the upper part of a carbonate ramp. Late Triassic (Rhaetian): Plattenkogel, Steinplatte near Waidring, Tyrol. Scale is 1 cm.
stones from the Austrian Calcareous Alps (Stanton and Flügel 1989). Both figures show cross sections of colonial scleractinian corals exhibiting the same fasciculate growth form. The samples differ in taxonomy and in the composition of the fine-grained matrix between the coral stems. The corals in sample A belong to Retiophyllia clathrata (Emmrich), the corals in sample B to Stylophyllum polycanthum (Reuss). The first species is one of the most abundant Norian and Rhaetian corals, forming coral thickets and contributing to the formation of patch reefs in high- and low-energy settings. The latter coral is relatively rare and occurs in small isolated patches in low-energy environments. Retiophyllia clathrata is a high-growing form, Stylophyllum polycanthum a low-growing form. Systematic assignment and different growth forms attribute sample A and sample B to separate biofacies types. Their separation into different MFTs is indicated by their distinct differences in matrix. The sediment in sample A consists of microbial micrite, partly forming irregular crusts on and Fig. 11.2. Coral limestone with Stylophyllum polycanthum (Reuss). Coral patch formed on the deeper and distal part of a carbonate slope. Late Triassic (Rhaetian): Cliff section, Steinplatte near Waidring, Tyrol. Scale is 1 cm.
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between the coral stems, and fine-grained bioclasts and peloids. Most bioclasts are of coral debris. The sediment in sample B consists of homogeneous micrite with echinoderm and shell debris. These differences are of greater significance for the definition of MFTs than are the taxonomic differences because they reflect different environmental and depositional conditions. The corals of sample A grew on the inner part of a carbonate ramp. The sediment between the coral stems was formed by physical and biological erosion of corals and by the microbial precipitation of fine-grained carbonate. Baffling of sediment took place in protected interspaces within the coral colony. The MFT should be called ‘coral (Retiophyllia) bafflestone with finegrained coral skeletal debris/microbial micrite matrix’. The corals of sample B developed on the lower and deeper part of a slope between a platform and a distally steepened ramp. Carbonate production took place by accumulation of mollusk shells, echinoderms and finegrained carbonate on the slope. The MFT is a ‘coral (Stylophyllum) bafflestone with fine-grained echinoderm-micrite matrix’. The message of this example is that less conspicuous differences (such as matrix types) may be more relevant in generating MFTs than more conspicuous features (such as coral types). Microbial carbonates: The MFTs of carbonate textures believed to be due to microbially induced processes (see Sect. 9.1) should be differentiated using the criteria summarized in Box 9.1 and following the classification of microbialites and stromatolites presented in Box 9.2. Microbial MFTs exhibit both layered and non-layered fabrics (see Pl. 50). Most layered fabrics correspond to bindstones. The names of MFTs should include both textural and compositional criteria. The latter can be indicated by the dominant constituents, e.g. thrombolitic bindstone (Pl. 8/6), or peloidal laminated bindstone (Pl. 50/2) corresponding to a ‘laminated fine-grained agglutinated stromatolite’. A particular microbial MFT characterized by an intimate association of abundant carbonate cement with microbiota is depicted in Fig. 8.3. Which criteria are suitable for establishing the microfacies types of allochthonous limestones? Both simple and more sophisticated methods can be used to distinguish between separate MFTs. Simple definitions are based on a combination of Dunham’s classification and the relative abundance of different grain types. Most commonly used criteria are depositional texture and fabric, as well as qualitative and quantitative compositional data. Box 11.1 summarizes these criteria.
Diagnostic Microfacies Type Criteria
Box 11.1. Thin-section criteria used in establishing microfacies types of allochthonous carbonates. Depositional texture of the rock • Mudstone, wackestone (floatstone), packstone, grainstone (rudstone); Sect. 8.3.2 • Grain size, sorting and packing; Sect. 6.1.2, e.g. Pl. 93 Qualitative composition • Fine-grained matrix, grain types, carbonate cement (sparite); Sect. 4.1 • Type of the fine-grained matrix (micrite, microspar, calcilitite); Sect. 4.1.1, 4.1.3, 4.1.4 • Basic micrite types (auto- and allomicrite, seafloor and internal micrite); Sect. 4.1.2 • Micrite texture (homogeneous, inhomogeneous, peloidal, fine bioclastic, laminated, silty); Sect. 4.1.5 • Grain categories (skeletal grains, peloids, cortoids, oncoids, ooids and pisoids, aggregate grains, resediments); Sect. 4.2 Quantitative composition • Relative proportions of fine-grained matrix, grains and interparticle carbonate cement (modal analysis); Sect. 6.2.1 • Abundance and percentage of major grain categories; Sect. 6.2.1 • Quantitatively dominant grain categories • Ranking of grain categories according to frequency; Sect. 6.2.1.4 Depositional fabrics • Biofabrics; Sect. 5.1.2 • Bedding and lamination; Sect. 5.1.3 • Burrowing and bioturbation; Sect. 5.1.4 • Fenestral fabrics; Sect. 5.1.5
The significance of depositional texture Depositional texture types, particularly those based on the original or expanded Dunham classification (see Sect. 8.3.2) permit a preliminary subdivision of limestone samples. However, the name determined by the application of the limestone classification rarely corresponds to a microfacies type. Consider the examples illustrated by Pl. 100/1, 2 and Pl. 101. All samples are grainstones following the Dunham classification, and all samples are skeletal grainstones. Using the Folk classification, the thin sections would be designated as biosparites, but taking into account compositional criteria, the three samples must be placed into two separate microfacies types. Grain size and sorting can be helpful in differentiating textural groups. Examples are shown in Pl. 15. The depicted grainstones correspond in composition, but size and sorting of the aggregate grains varies indicating slight hydrodynamic differences in the depositional environment.
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Have a look at specific criteria of matrix and grains! Matrix specifics: Differences in matrix texture reflect different controls on the formation of fine-grained matrix. The distinction between seafloor micrite and internal micrite is particularly important for packstones, densely packed wackestones and floatstones as well as for reef limestones. Consider Pl. 44/3: most of the fine-grained sediment in the interstices between the larger grains of this coralalgal packstone is an internal micrite that was infilled into a grain-supported fabric. This infilling indicates short-term events that transported fine-grained sediment into depositional areas characterized by fluctuating hydrodynamic conditions. The recognition of sedimentary infilling is crucial in the definition of MFT based on baffling processes. An example can be found in Pl. 41/1.
Limestones composed of lithoclasts, intraclasts and extraclasts (Sect. 4.2.8) often yield a wide spectrum of genetically indicative MFTs. Again, a texturally oriented classification would lump together genetically different carbonates. Pl. 16/1, Pl. 16/3 and Pl. 16/7 are intraclast wackestones, but the first two samples represent a MFT defined by intraclasts corresponding to black pebbles indicative of near-coast depositional sites, whereas the intraclasts of the sample shown in Pl. 16/7 are indicative of a tidally influenced depositional site.
Grain specifics: A purely descriptive approach to some of the grain categories discussed in Sect. 4.2 results in a loss of information. In contrast with peloids and cortoids, which commonly exhibit only few morphological criteria that would allow further differentiation, the morphological features of other grain types offer possibilities for the discrimination of genetically defined subtypes. Oncoids (Sect. 4.2.4.1) formed in different environments differ in shape and growth forms, internal growth patterns, and the biotic composition of coatings. The latter feature is a diagnostic criteria for MFT differentiation. Consider Pl. 12, which exhibits limestones with oncoids formed by cyanobacteria (cyanoids) and by red algae (rhodoids). Clearly limestones with cyanoids and rhodoid limestones represent different MFTs. A comparison of Pl. 12/3, Pl. 12/4 and Pl. 12/7 shows marked compositional differences between cyanoids, which attribute the samples to three different MFTs. The differentiation of these three MFTs does not result from a preference for splitting but is meaningful in the context of microfacies interpretation: the cyanoids of Pl. 12/3 characterize a reef environment, those of Pl. 12/4 a platform environment and the cyanoids depicted in Pl. 12/7 are typical for brackish-water settings. Ooids (Sect. 4.2.5) and pisoids (Sect. 4.2.6) originate in different environments and are subject to different controls. The descriptive criteria summarized in Box 4.15 offer the possibility of defining MFTs for oolitic limestones that reflect particular paleoenvironmental conditions (water energy and transport processes, salinity, water depth) and depositional settings (Box 4.17). The samples in Pl. 13/4, 5, 6 represent three well-defined microfacies types although all samples would be rather simply classified as oolitic grainstones.
Skeletal grains are of prime importance in defining microfacies types Limestones are predominantly biogenic sediments (Sect. 2.1.2) formed by biologically controlled processes (Chap. 9). Fossils are significant proxies for paleoenvironmental conditions: skeletal grains are highly sensitive to processes characterizing specific depositional environments. The composition of bioclastic carbonates depends on many factors including skeletal mineralogy and the path of taphonomic and diagenetic processes examined in Sect. 4.2.1 and referred to in the context of the discussion of the various fossil groups (Chap. 10). The taphonomic filter as well as the diagenetic loss of paleontological data must be thoroughly considered in the evaluation of MFTs based predominantly on the occurrence and frequency of skeletal grains (see Pl. 9/7). Nevertheless the type, association patterns, and systematic classification of fossils, as well as their abundance, are the basic criteria for the discrimination of significant microfacies types. A differentiation of MFTs based on skeletal grains should include answers to the questions listed in Box 11.2. Combining the criteria summarized in Box 11.2 results in a wide range of MFTs and allows for environmental and depositional interpretations on different levels. Consider Fig. 11.3 and try to answer the questions in Box 11.2: The thin section exhibits a poorly sorted skeletal packstone. Fossil groups represented by skeletal grains are echinoderms, brachiopods and mollusk shells (low thin-section diversity; question 1). Echinoderms and shell fragments are common, brachiopods are rare (question 2). Echinoderms are better preserved
Microfacies types of Proterozoic limestones lacking skeletal grains are defined by microbialite types, by differences in intraclast categories and ooid constructional types, by compositional patterns of wackestones and packstones consisting of extraclasts, and by depositional fabrics (Sami and James 1993).
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Box 11.2. Skeletal grain criteria relevant in defining microfacies types of limestones. 1
Which fossil groups are represented by the skeletal grains (describes the ‘thin-section diversity’)?
2
Which fossil groups are abundant, common and rare?
3
If calcareous algae or foraminifera occur, which major systematic units (e.g. dasyclad green algae or corallinacean red algae; orbitolinid or nummulitinid foraminifera) are present?
4
Are there differences in the preservation (e.g. fragmentation, solution, recrystallization) of the skeletal grains?
5
Is there an apparent loss or predominance of grains of a specific skeletal mineralogy (e.g. no aragonitic grains)?
6
Are the sizes of the skeletal grains within comparable ranges or are some grains conspicuously larger or smaller than the average grains?
7
Do the skeletal grains represent benthic (mobile, sessile) or planktonic organisms?
8
Do the grains exhibit biogenic encrustations?
9
Which fossil groups appear commonly associated (e.g. radiolarians and sponge spicules, or echinoderms and bivalves)? Are there striking grain associations (that might point to specific paleoclimatic conditions)?
10 Show the fossil concentrations in your sample (e.g. shell coquinas, Sect. 3.1.1.3; accumulations of sponge spicules, Pl. 78.2, or of crinoids, Pl. 95.4)! Are these concentrations allochthonous or autochthonous (Fig. 10.48)?
than shells which are strongly fragmented; some echinoderms exhibit borings (question 4). Predominantly High-Mg calcite echinoderms and calcitic brachiopod shells occur but are differently preserved; no indication of aragonitic grains (question 5). Grain size varies within a wide range; it reflects growth size for echinoderm elements but the shell size is a result of mechanical destruction (question 6). The skeletal grains indicate the existence of sessile benthos (crinoids, brachiopods) and mobile benthos (echinoids, recorded by rare spines), question 7. Biogenic encrustations are absent (question 8). The sample represents an example of a fossil concentration (question 10). No answer is possible for question 9. Which of the criteria will define a specific microfacies reflecting environmental and depositional conditions? The biofabric characterized by a concentration of echinoderms and shells and the admixture of variously fragmented bioclasts indicate a deposition of skeletons of organisms from a single environment, their
destruction and their redeposition. Redeposition must be one of the criteria used in the definition of the MFT. The qualitative biotic composition together with the predominance of echinoderms and mollusk shells is another criterion. Further criteria valuable in comparing the sample with texturally and compositionally similar samples are the presence or absence of brachiopod shells, and the presence not only of crinoids but also of echinoids. Borings in the echinoderm fragments may have taken place prior to or subsequent to deposition of the sediment, and are therefore less useful in defining the MFT. The differences in preservation of echinoderm, mollusk and brachiopod grains tell a lot about burial diagenesis but do not assist in MFT discrimination. In contrast, the infilling of interstices between grains by inhomogeneous micritic matrix is an important MFT criterion. The matrix type together with the packstone texture, reflects the deposition of allochthonous bioclastic sediment by a mass flow event. This explanation is in agreement with field data: the sample comes from carbonate beds dipping with a paleoslope that separates a non-rimmed platform from a distally steepened carbonate ramp (Stanton and Flügel 1989). Quantitative composition and multivariate discrimination of microfacies types Many authors differentiate MFT according to the frequency of compositional constituents using point counter data or abundance estimations (see Chap. 6). The abundance of micrite, sparite and different grain categories determined by modal analysis characterizes lithofacies types that are commonly expressed by limestone classification groups. Establishing MFT requires a different approach focused on the abundance of grains and the grain types. Abundance can be expressed by the frequency (or relative frequency) of grain categories, by considering the most abundant grain category, or by ranking grain categories according to their frequency in the sample. Abundance and frequency form basic data for multivariate discriminations of MFT and facies types (Sect. 6.3.2). Hierarchical methods (resulting in a dendrogram) as well non-hierarchical numerical classifications (producing a number of unaffiliated clusters) are used to discriminate between samples and establish facies groups. Hierarchical methods are either agglomerative (forming clusters by combining the most similar individuals first, and continuing to combine similar clusters into partitions of progressively higher rank until all individuals are in a single set) or divisive (subdividing an initial set into classes of progressively lower hierarchical rank). The use of hierarchical methods for the discrimination of ancient carbonate rocks can be
Microfacies Types
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Fig. 11.3. Bioclastic echinoderm-shell packstone used as an example for discriminating microfacies types. B - brachiopod shells, C - crinoid fragments, E - echinoid spine, M - mollusk shells. Late Triassic (Rhaetian): Steinplatte near Waidring, Tyrol, Austria. Scale is 1 mm.
tested on data sets collected from modern carbonates. Compositional and distributional data for shallow-marine carbonates of the Great Bahama Bank (Purdy 1963a, 1963b) have been used several times to test the value of multivariate discrimination methods (Gill 1993). Frequency data for twelve sedimentary and biogenic constituents per sample, derived from pointcounting of thin sections of plastic-impregnated sediment samples, were grouped using agglomerative cluster analysis to obtain a grouping of constituents and factor analysis to distinguish the main facies groups. The statistical division into facies groups (corresponding roughly to microfacies types) showed good agreement with the sediment grouping based on geological field data, except for those carbonates characterized by high amounts of carbonate mud. Presence or absence data, multistate nominal data, and binary data obtained by dichotomizing each vari-
able about its mean value (Shi 1993) are often regarded as sufficient to present meaningful classifications of multivariate facies data (Schott 1983, 1984; Caracuel et al. 1998). Microfacies types and sedimentary fabrics and structures MFTs based on textural and compositional criteria should be tested with respect to variations caused by differences in depositional fabrics. Specific biofabric features (Sect. 5.1.2) including grain orientation (Pl. 18/1, 5) and small-scale cross-bedding (Pl. 18/3) may be useful MFT criteria. Graded bedding layers may exhibit different MFTs, particularly if textural criteria are considered as being of higher importance than grain composition (see Pl. 17/1). Biolamination (Pl. 18/4) and lamination fabrics caused by bottom currents (Pl. 18/ 7) are other MFT characteristics. Wackestone MFTs can
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be differentiated by the intensity and distribution of burrowing and bioturbation (compare Pl. 19/7 with Pl. 19/ 6). Fenestral fabrics are important environment-sensitive MFT indicators. Have a look at Pl. 20/1, Pl. 20/2 and Pl. 20/3! These photographs illustrate the differences with respect to the geometry and frequency of open-space structures. Sedimentation on carbonate ramps is strongly controlled by lateral variations in hydrodynamic conditions resulting in a wide range of sedimentary structures (e.g. bedding types, channels, erosion marks) and lithofacies textures (Dunham types). A specific characteristic is the close occurrence of areas with in situ sedimentation and areas dominated by the deposition of resediments and transported grains. These differences in depositional patterns and sedimentary structures assist greatly in defining environment-sensitive MFTs for carbonate ramps (Bachmann 1994; Rüffer 1995). Diagenesis and microfacies types Diagenesis may or may not change textural and compositional criteria used for the definition of MFTs. Compare the samples depicted in Pl. 28/1 and Pl. 28/2! The original structures of the skeletal grains of the reef limestone (Pl. 28/1) are strongly altered by recrystallization. The depositional microfacies is, however, still recognizable. The constituents of the lagoonal limestone in Pl. 28/2 are, in contrast, considerably obliterated by burial cementation and dolomitization, making MFT definition difficult (another example is shown in Fig. 7.21). Solution enlargement of interparticle pores (Pl. 29/1) and selective dissolution of grains (Pl. 29/8) can
How Many Microfacies Types?
greatly alter depositional microfacies and make MFT establishment more difficult. On the other hand, the combination of compositional and textural microfacies criteria with diagenetic criteria can assist in defining MFTs. Early diagenetic marine and meteoric carbonate cements indicate specific conditions within different parts of a reef complex characterized by different microfacies types (Städter and Koch 1987). Diagenetic microfacies is commonly defined using observations from petrographic (thin section, SEM, CL) and isotopic analyses, as well as from outcrop and core data, focusing on carbonate cements, porosity types and diagenetic textures (Chap. 7). Combination of these criteria results in the definition of ‘diagenetic microfacies types’ that are critical to the development of reservoir models in carbonate sediments (see Chap. 17). How many microfacies types are reliable? A survey of about hundred papers shows that carbonates of different depositional environments are described by different numbers of microfacies types. Peritidal carbonates have been characterized by only a few MFTs (up to 5); lagoonal carbonates by up to 10 MFTs. Reef complexes including back-reef, central reef and forereef were described by 10 to 15 MFTs. Carbonate ramps and carbonate platforms have been differentiated into up to 22 MFTs. Platform margin carbonates (excluding reef carbonates) display 5 or less MFTs. Similar low numbers are reported from carbonates formed on open-marine slopes, but the number of MFTs is highly variable because of differences in the fre-
Plate 100 Defining Microfacies Types The interpretation of depositional and environmental constraints from microfacies data requires a synopsis of those elements that are characteristic of the sediment. A generalization of microfacies types leads to categorization of common microfacies data into ‘standard microfacies types’ (Sect. 14.3). The plate displays two samples of a skeletal grainstone from a carbonate platform which are used to demonstrate the methodology used to define microfacies types (MFTs). 1 This sample A yields calcareous algae (udoteacean green algae: Neoanchicodium - N; dasyclad green algae - D), foraminifera (fusulinid larger foraminifera - F, palaeotextulariid smaller foraminifera - P), echinoderms (crinoids - C, echinoids E), and gastropods (G). Composite grains are represented by microbial and encrusted aggregate grains (AG). Sorting is moderate. Interparticle pores are occluded by carbonate cements. The MFT is typified not by the depositional texture (moderately sorted skeletal grainstone) but by the dominant skeletal grain category (Neoanchicodium) and the association of the alga Neoanchicodium, microbial and encrusted aggregate grains, echinoderms and fusulinids. Early Permian (Asselian) platform carbonates: Zweikofel, Carnic Alps, Austria. 2 Sample B contains calcareous algae (Neoanchicodium - N, dasyclads - D), foraminifera (fusulinids - F), and echinoderms (crinoids - C). Large irregularly shaped micritic grains are aggregate grains (AG). Sorting is moderate. Interparticle pores are filled with thin rims of marine cements (arrow) and gray crystal silt (CS). The microfacies criteria agree with those of sample A. A striking difference exists with regard to the type of interparticle cement The coincidence of compositional criteria attribute the sample to the same MFT as sample A. The different cement types indicate a different diagenetic history for limestone beds represented by the samples A and B, but not a different depositional history. Same locality as –> 1.
Plate 100: Defining Microfacies Types
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quency of redeposited reef- or platform-derived material. Basinal limestones are commonly described using only a few microfacies types (maximum 8 MFTs). The difference in the number of MFTs for shallow-marine and deep-marine limestones reflects the greater number of microfacies studies dealing with shelf and reef carbonates and also the greater diversity of shallowmarine benthic biota used in defining microfacies types.
11.3 Making Microfacies Types Plate 100/1 (subsequently called sample A), Pl. 100/ 2 (sample B) and Pl. 101 (sample C) show three thin sections of Early Permian shelf carbonates from the Southern Alps. The samples A and B are from the same section, sample C is from a different but coeval locality. The three samples are similar in depositional texture (grainstone), sorting and gross composition (more skeletal grains than composite grains). They are also similar in the occurrence of calcareous algae, fusulinids, smaller foraminifera and microbial and encrusted aggregate grains, as well as in the generic attribution of the fusulinids. These coincidences indicate deposition in a shallow-marine, high-energy environment. The shared criteria would unite the three samples in a single facies type defined by grain composition and depositional texture. A closer look, however, reveals differences: several biota occur only in some samples - gastropods only in sample A, echinoderms only in samples A and B, and thin shells, corals and tetrataxiid foraminifera only in sample C. These differences might appear negligible but there are also additional differences concerning the quantitative ranking of the constituents and the systematic attribution of algae. Ranking in samples A and B is similar with regard to Neoanchicodium and aggregate grains, both of which are the most abundant grains. In sample C, however, the first two positions are occupied by dasyclad algae and aggregate grains, followed by fusulinids.
Examples for Microfacies Types
These differences have to be taken into consideration in the definition of MFT. Samples A and B comprise one microfacies type (MFT 1), whereas sample C represents another microfacies type (MFT 2). The microfacies types are discriminated by environmentally sensitive criteria which allow the recognition of specific habitats. MFT 1 is dominated by Neoanchicodium grains, MFT 2 is dominated by dasyclad grains. The dasyclads comprise at least three taxa. Echinoderm grains occur in MFT 1 but not in MFT 2. MFT 2 differs from MFT 1 by a greater frequency of fusulinids, although this feature may not be relevant for the discrimination of the microfacies types, but may simply indicate very localized concentrations of tests by sporadic currents. Does the detailed differentiation of MFTs make sense? If you are satisfied with the statement that the samples indicate shallow-marine deposition in a highenergy environment, further studies will not be warranted. If, however, you would like to interpret this shallow-marine high-energy depositional environment in terms of subenvironments and habitats, you will need MFTs whose diagnostic criteria reflect specific settings and environmental conditions. Such additonal criteria can be derived from taxa and major groups of Early Permian calcareous green algae which exhibit specific distribution patterns. Facies analyses of several Early Permian platforms have shown that udoteacean algae (Neoanchicodium) dominate in the distal parts of openmarine platforms, whereas the Gyroporella-MizziaConnexia dasyclad association of sample C is a common feature of Permian proximal platform areas. Another example of making microfacies types is shown in Fig. 11.4 which shows three thin sections of Mesozoic sediments with serpulids. Using the occurrence of serpulids only, the three samples would be attributed to a single MFT. This, however, leads to a loss in information. The scarcity of serpulid tubes in sample A separates this sample from samples B and C which would be placed in an other MFT. The attribution of sample A to a distinct MFT is also supported by the
Plate 101 Defining Microfacies Types This plate displays an example of how microfacies types can be defined. The sample yields some microfacies criteria in common with samples A and B on Pl. 100, but differs in environmentally-sensitive characteristics. Skeletal grainstone. Sample C displays calcareous algae (dasyclads - D), foraminifera (fusulinids - F, palaeotextulariid - P) and tetrataxiid -T, and smaller foraminifera), and thin shells (S). Composite grains are represented by microbial and encrusted aggregate grains (AG; see Pl. 15/4, 5). Interparticle pores are filled by carbonate cements. The diagnostic microfacies criteria that attribute the sample to a different MFT different from that of samples A and B (on Pl. 100) are the abundance of specific dasyclads comprising an assemblage of Gyroporella (GY), Mizzia (MI) and Connexia (CO) that typify proximal platform habitats. Early Permian (Asselian) platform carbonates: Seikofel, Sexten Dolomites, Italy.
Plate 101: Defining Microfacies Types
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Examples for Microfacies Types
Fig. 11.4. These three samples of serpulid limestones represent three different microfacies types defined by frequency, allochthonous occurrence (A, B), and autochthonous growth (C). This is a reduced picture of Fig. 10.48. Scale is 2mm.
association of broken serpulid tubes with quartz sand, indicating a quite different near-coastal depositional environment compared with that of samples B and C. The abundance of serpulid tubes in sample B is caused by hydrodynamic concentration, that in sample C by serpulids growing in colonial aggregates that can form reefs. These differences in the origin of abundance place the samples into two separate microfacies types. Basics: Microfacies types Aigner, T. (1984): Dynamic stratigraphy of epicontinental carbonates, Upper Muschelkalk (M. Triassic), South-German Basin. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 169, 127-159 Bachmann, M. (1994): Die Karbonatrampe von Organya im oberen Oberapt und unteren Unteralb (NE-Spanien, Prov. Lerida): Fazies, Zyklo- und Sequenzstratigraphie. – Berichte aus dem Fachbereich Geowissenschaften der Universität Bremen, 54, 202 pp. Flügel, E. (1972): Mikrofazielle Untersuchungen in der Trias: Methoden und Probleme. – Mitteilungen Gesellschaft der
Geologie- und Bergbaustudenten in Österreich, 21, 6-64 Gill, D. (1993): Discrimination of sedimentary facies by association analysis. – Mathematical Geology, 25, 471-482 Hladil, J. (1986): Trends in the development and cyclic patterns of Middle and Upper Devonian buildups. – Facies, 15, 1-34 Machel, H.G., Hunter, I.G. (1994): Facies models for Middle to Late Devonian shallow-marine carbonates, with comparisons to modern reefs: a guide for facies analysis. – Facies, 30, 155-176 Rüffer, T. (1995): Entwicklung einer Karbonat-Plattform: Fazies, Kontrollfaktoren und Sequenzstratigraphie in der Mitteltrias der westlichen Nördlichen Kalkalpen (Tirol, Bayern). – Gaea heidelbergensis 1, 1-282 Schott, M. (1984): Mikrofaziell-multivariate Analyse einer rhäto-liassischen Karbonatplattform in den Nördlichen Kalkalpen. – Facies, 11, 229-280 Städter, T., Koch, R. (1987): Mikrofazielle und diagenetische Entwicklung einer devonischen Karbonatfolge (Givet) am SW-Rand des Briloner Sattels. – Facies 17, 215-230 Wilson, J.L. (1975): Carbonate facies in geologic history. – 471 pp., Berlin (Springer) Further reading: K146, K205 - K208 Fig. 11.5. Exercise: Look for bottom/top structures - improve orientation! Which sort of clasts? Which fossils can be recognized? What do they indicate? How are they preserved? Some of them are indicative for the age of the sample! Which depositional environment? What about pore space and cement? Which differences exist between here and Pl. 100 and Pl. 101? Which criteria are effective in defining an environmentallysensitive MFT? For answers see Appendix!
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12 Recognizing Paleoenvironmental Conditions
This chapter provides an overview of microfacies data that assists in recognizing environmental parameters that control carbonate deposition and the distribution of organisms. Section 12.1 summarizes paleoenvironmental proxies. Specific attention is given to paleoclimate-indicative criteria (Sect. 12.2) and to paleo-depth criteria (Sect. 12.3). The last section of this chapter deals with microfacies data that allow seismic events to be recognized. Prime marine environmental parameters are nutrient supply, salinity, water energy, oxygen, turbidity, substrate, rate of sedimentation, water depths, light and temperature. Modern world oceans are characterized by three all-important ecological gradients: (a) depth gradient; light intensity, temperature and oxygenation decreasing with increasing water depth, (b) shore to open ocean gradients, e.g. nutrient levels decreasing from inland waters to the open ocean, and (c) latitudinal gradients, e.g. higher productivity at low latitudes. These gradients are commonly modified by local and regional controls as well as changes during time. The reconstruction of carbonate paleoenvironments and the conditions that controlled deposition and lithification need a critical and combined assessment of paleontological, sedimentological and geochemical data.
12.1 Reconstructing Environmental Constraints Methods for paleoenvironmental reconstructions are: • An actualistic approach, which compares fossil organisms with Recent relatives. • The geochemical approach that uses information locked in the isotopic or trace element record of biogenic hardparts (Sect. 13.2). • The functional approach, which interprets the morphology of biogenic hardparts according to the real or inferred function of the skeleton. • An ecological principle approach, which applies basic ecological concepts (e.g. the relationship between species diversity and the degree of environmental stress; or the relationship between trophic group
composition and energy level) to fossil assemblages. • An integrated approach, which combines different methods and uses information from the sedimentary record including microfacies. Most microfacies criteria aimed at a discrimination of ancient environments are qualitative. They refer to the (1) modes of composition, degradation and preservation of carbonate grains (e.g. growth continuities and discontinuities, coating, worning, breaking, rounding), (2) textural differences (grain size, packing, sorting, maturity level) as well as (3) homogenization, bioturbation and bio-retexturing of the sediment. Limitations to paleoenvironmental reconstructions concern the interplay of environmental parameters and environmental condensation. The latter term describes the projection of two or more environments onto a single stratigraphic level during times of reduced or missing sedimentation. The time involved ranges from several months and years to hundreds or thousands of years. Replacement of one environment by another will cause autochthonous mixing of different communities, specifically in shallow shelf environments. Environmental condensation may be recognized by sedimentological and taphonomic criteria as well as the principle of ecological incompatibility (Fürsich and Aberhan 1990; Kidwell 1991). The paleoenvironment is the sum of non-biogenic parameters that provide the framework of where ancient organisms lived. Because strong interactions exist between organisms and their environment, fossils are the most important indicators of ancient environments. The interpretation of paleoenvironments rests strongly on palecologic analyses. Several texts, listed below, deal with the palecologic approach and also provide information on relevant proxies which will be briefly referred to in the following sections. Basics: Paleoecology and environmental proxies Bosence, D.W.J., Allison, P.A. (eds., 1995): Marine paleoenvironmental analysis from fossils. – Geological Society of London, Special Publication, 83, 272 pp. Boucot, A. (1981): Principles of benthic marine paleoecology. – 463 pp., London (Academic Press)
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_12, © Springer-Verlag Berlin Heidelberg 2010
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Brett, C.E., Baird, G.C. (eds., 1997): Paleontological events: stratigraphical, ecological, and evolutionary implications. – 604 pp., New York (Columbia University Press) Dodd, J.R., Stanton, R.J. (1990): Paleoecology. Concepts and applications. Second edition. – 502 pp., New York (Wiley) Etter, W. (1994): Palökologie. Eine methodische Einführung. – 294 pp., Basel (Birkhäuser) Fischer, G., Wefer, G. (1999): Use of proxies in paleoceanography. – 735 pp., Berlin (Springer) Fürsich, F.T. (1995): Approaches to paleoenvironmental reconstructions. – Geobios, Mem. Spec., 18, 183-195 Martin, R.E. (1999): Taphonomy. A process approach. – 524 pp., Cambridge (Cambridge University Press) Nybakken, J.W. (1993): Marine biology. An ecological approach. Third edition. – 462 pp., New York (Harper Collins College Publisher) Further reading: K109
12.1.1 Hydrodynamic Controls Paleoenvironmental interpretations of carbonate deposits rely strongly on estimations of hydrodynamic conditions and processes, whose definitions are listed in Box 12.1. The effects of severe storms are discussed in Sect. 12.1.2. The kinetic energy of motion existing in the water at the depositional interface or a few tens of centimeters below it is caused by waves or current action. Mechanisms that drive the flow of water across the sediment-water interface and within the pore spaces include wave-induced advection, density driven convection, groundwater seepage and active pumping by benthic organisms (Libelo et al. 1994). Movement of water across the sediment-water interface and within near surface sediments control the cycling of nitrogen, phosphorus, carbon, sulfur and trace metals between sediment and the water column. Low-energy and high-energy levels: Depending on various oceanographic factors (e.g. water depth, wave amplitude) the energy level may be fairly constant or may vary strongly over time (e.g. in tidal zones or in areas affected by storms). The energy of a water mass capable of moving sedimentary particles varies between low (low-energy sediments) and high (high-energy deposition). Generally high-energy environments include the intertidal zone, shallow-marine open platforms, inner ramps and the leeward parts of framework reefs. Lowenergy environments occur in protected areas, such as enclosed bays, estuaries, lagoons, inner platform areas, outer ramps and in deeper-water settings. Benthic organisms exhibit specific adaptions to high or low-energy conditions (Riedl 1969; see Nybakken 1993 for a review). High-energy conditions provide a constant
Hydrodynamic Controls
Box 12.1. Terms used for discussing depositional water energy. • Depositional interface: Interface between the water and the bottom where sediments are deposited in relation to the energy level at the interface. • Energy index: Inferred degree of water agitation in the sedimentary environment of deposition. • Energy level: The kinetic energy (due to wave or current action) that existed or exists in the water of a sedimentary environment either at the interface of deposition or above it. Depending on the water depth, wave amplitude and storm activity, the energy level at the depositional interface may be fairly constant, or may range within wide limits as a function of time. • High-energy environment: An aqueous sedimentary environment characterized by a high energy-level and by turbulent action (such as that created by waves, currents or surf) that prevents the settling and accumulation of fine-grained sediment. • Low-energy environment: An aqueous sedimentary environment characterized by a low-energy level and by standing water or a general lack of wave and current action, thereby permitting fine-grained sediment to settle and accumulate. • Quiet-water sedimentation: Deposition within a near zero energy level at the depositional interface. Quietwater does not necessarily imply the complete absence of water movement or the existence of stagnant conditions. • Storm wave base: The water depth down to which storm-generated waves and resulting bottom currents rework bottom sediments (generally lime mudstones and wackestones) and produce specific texture types (tempestites, see Sect. 12.1.2.1). The depth of the storm wave base varies strongly (about 50 m down to 250 m in shelf seas, and about 20 to 30 m in epicontinental seas). • Surf: (a) The wave activity in the area between the shoreline and the outermost limit of breakers. (b) A collective term for breakers. • Wave base: The water depth below which surface wave action no longer stirs and moves the sediment. Also called fair-weather wave base. The wave base ranges widely, depending on wave amplitude and fetch, bottom topography and activity of storms. • Tsunami: A gravitational sea wave produced by largescale, short-duration disturbance of the ocean floor, principally by earthquakes. Characterized by great speed of propagation, long wave lengths and long periods. May pile up waves of 30 m height and more, causing much damage in shallow-marine water, along coasts, and on land. Ancient tsunami deposits may be recognized by extended sheets of marine sand overlying terrestrial or lacustrine sequences (Minoura et al. 1994; Shiki et al. 2000).
source of nutrients and oxygenated seawater and reduce the settling rate of fine-grained sediment. The reconstruction of water energy is based on sedimentological and paleontological criteria. Paleontological criteria include specific morphologies of skeletal
Energy Levels
elements, life habits, the trophic composition of paleocommunities and trace fossil assemblages. Microfacies data play an essential role in evaluating the results of these reconstructions.
12.1.1.1 Hydrodynamic Energy Levels Energy levels are reflected by microfacies criteria summarized in the following text. The recognition of energy levels based on depositional textures is discussed in Sect. 12.1.1.2. Fig. 12.1 gives an idea of the ideal relations between the winnowing of fine-grained particles, sorting and rounding of grains and the intensity of hydrodynamic processes. Microfacies are proxies for hydrodynamic depositional conditions of carbonate sediments. Matrix Micrite: The occurrence and amount of seafloor micrite (Box 4.2) as well as the micrite : grain ratio are generally regarded as an indication of low-energy depositional environments. Limitations to this concept are discussed in Sect. 4.1.6. Allomicrites and automicrites
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must be evaluated separately. Automicrites (e.g. microbialites) can originate in low-energy as well as in highenergy environments. Grains Oncoids originate in low-energy and high-energy environments (Box 4.10). Low-energy environments are characterized by single or multiple micritic oncoids exhibiting discontinuous growth layers, often forming floatstones and wackestones. High-energy oncoids are small, exhibit rather regular laminations, and commonly form packstones and grainstones (rudstones). Rhodoids formed under low-energy and high-energy conditions differ significantly in shape, external and internal growth, as well as the type of ecological successions (Sect. 4.2.4.2). Ooids are formed in high-energy and low-energy environments (Sect. 4.2.5; Fig. 4.20). High-energy environments are indicated by concentric (tangential) or micritic ooids, and broken or distorted ooids (Fig. 4.23). Low-energy environments are characterized by radial (radial-fibrous) ooids or asymmetrical (eccentric ooids; Fig. 4.23). The maximum and average ooid size and the thickness of ooid coatings can be used as an approximate measure of water energy (Fig. 4.25).
Fig. 12.1. Winnowing, sorting and rounding are controlled by the intensity of hydrodynamic processes. Winnowing of lime mud taking place at low energy levels may be absent (A) in low-energy environments, incomplete (B) in environments with fluctuating energy conditions, or complete (C and D) in high-energy environments. An increase in the intensity of wave action and currents leads to changes from unsorted (C) to sorted (D) sediments with layers of different grain size. An increase in rounding (E) may reflect an increase in water energy, e.g. in the surf zone. Sorting and rounding of carbonate grains will decrease (F and G) when extreme water energy (e.g. severe storms) affects the sediment. The example shown refers to shelly limestones. Black areas between skeletal grains represent fine-grained matrix (e.g. micrite), white areas between skeletal grains intergranular pores filled with calcite cement. After Folk (1962).
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Water Energy Criteria
Pisoids formed under low-energy conditions exhibit irregular shapes and discontinuous lamination. Highenergy pisoids are characterized by regular spherical shapes and concentric laminations (Sect. 4.2.6). Aggregate grains are usually formed in low-energy to moderate-energy environments with fluctuating water circulation (Sect. 4.2.7). Intraclasts originate both in low-energy and highenergy environments (Sect. 4.2.8.1). Imbrication patterns point to controls by weak currents (Sect. 5.1.2). Flat-pebble conglomerate patterns indicate medium to high-energy currents. Morphometric criteria of grains Changes in grain roundness and sphericity (Sect. 4.3.2) in association with changes in grain size and sorting may be used in recognizing different hydrodynamic levels (Fig. 4.32, Fig. 8.6). Of particular interest are rounded fragments of fossils that were not originally round (Fig. 4.31) and angular or rounded resediments (intraclasts and extraclasts). Note that the roundness of skeletal grains depends on many factors, including the architecture and microstructure of the shells, grain size, water energy and the length of transport. Grain size Covariation of grain size and sorting (Sect. 6.1.1.2), grain size parameter diagrams and the maximum grain diameter (Sect. 6.1.2.2) give some clues to depositional energy levels if the limitations discussed in Sect. 12.1.1.2 are considered. Fabrics Biofabrics and grain orientation are valuable indicators of water energy levels (Sect. 5.1.2) and bottom currents. High-energy conditions are reflected by imbrication (Pl. 18/1)), cross-bedding (Pl. 18/3) and stacked shells. Low-energy and moderate energy conditions are indicated by parallel oriented thin platy skeletal grains (Pl. 18/5). Orientation patterns of intraclasts and flat-pebble conglomerates can be used in recognizing flow-current directions. Some depositional breccia are indicators of highenergy conditions, e.g. peritidal breccia and forereef breccia (Sect. 5.3.3.3). Water Energy Indices Textural criteria (size, sorting and roundness of grains, matrix types) and compositional data (fossils, carbonate and non-carbonate admixtures) as well as sedimentary structures allow inferred water energy to be recognized by five energy indices that describe carbonates deposited under quiet, intermittently agitated,
Fig. 12.2. Controls on growth forms and growth patterns of colonial reef builders by water energy and sedimentation rates. The relationship between the external shape and the external growth banding geometry of the colonies can be used to infer water roughness and the relative rates of sedimentation. Enveloping and non-enveloping growth forms result from the position of the living surface during growth. These positions are expressed in corals by successive growth bands, in stromatoporoids by latilaminae (Sect. 10.2.3.1). Enveloping forms indicate increasing stability to water roughness by a decreasing height to width ratio. Progressive burial in sediment increases stability related to ragged margins. The latter is a measure of the capacity of the organism to keep pace with sedimentation. Ragged margins point to high-energy conditions, smooth margins to low-energy conditions. After James and Bourque (1992).
slightly agitated, moderately agitated and strongly agitated conditions (Sect. 12.1.1.2). Growth forms of sessile organisms The external geometry and the growth patterns of reef-building organisms have received considerable attention as proxies for prevailing hydrodynamic conditions. Growth forms of corals and stromatoporoids
Low and High-Energy Environments
are used in assessing paleo-waterdepth and differentiating high-energy and low-energy environments of Paleozoic and Mesozoic reefs, assuming that growth form types are predominantly related to adaption to the physical environment. Major limitations to this concept were discussed by Stearn (1982), who stresses that variations in shape are the result of interactions between various environmental factors and genetically controlled growth patterns of the organisms. Nevertheless, some useful generalizations concerning the relationships between organism growth forms and hydrodynamic controls are possible (Fig. 12.2). Relations between water energy and nutrient supply The distribution of sponges and suspension-feeding bivalves is limited by turbidity, those of crinoids and blastoids by the energy conditions of the environment. These limitations cause different nutrient supplies and potentially different distribution patterns. See Ausich (1983) for an example.
12.1.1.2 Classifying Low-Energy and HighEnergy Environments All limestone classifications based on texture and focusing on the relative abundance of matrix and grains are at least in part genetic (see Sect. 8.3). The basic subdivisions are oriented to interpretative, energy levels‘ distinguishing between sediments deposited in quiet waters (‘low-energy’ sediments) and turbulent waters (‘high-energy’ sediments). The basic philosophy behind these classifications is that differences in water energy should be reflected in sediment texture and biota. This rather simple dual differentiation does not consider the broad variations in wave- and current-controlled sedimentation patterns and resulting depositional textures. Environmental and depositional interpretations based solely on Dunham’s or Folk’s rock types may, therefore, be misleading. The Energy Index Classification discussed below offers a better possibility for recognizing depositional conditions controlled by low or high water energy conditions. Energy Index Classification Differences in water energy are reflected by a limestone classification proposed by Plumley et al. (1962). It is based on the inferred degree of water agitation in the depositional environment. Textural and compositional data are used in grouping limestones into genetic categories. Textural criteria include size, sorting and roundness of grains, and the nature of the matrix. Compositional data refer to mineralogy (carbonate and non-
591
carbonate admixtures) and fossils (types, abundance and associations). In addition, sedimentary structures are important for augmenting the interpretation of textural data. The term ‘clastic particle’ is used for (carbonate) grains that have been mechanically transported by wave action. A crucial point of this concept is the distinction between sedimentary particles that have been transported and those that have not. The authors offer a list of criteria that may infer the existence of agitated water environments or at least some transportation: • Angular or rounded fragments of partially endured sediment or of preexisting rock (intraclasts and extraclasts). • Distinctly rounded fragments of fossils. • Poorly sorted matrix, e.g. sand grains within a finegrained matrix (can also be caused by bioturbation). • Carbonate grains mixed with terrigenous clastics (e.g. detrital quartz) of the same size. • Mixtures of ecologically incompatible biota (e.g. dasycladacean algae in crinoidal limestones). • Ooids (excluding ‘quiet-water ooids’!). • Wave-resistant colonial organisms in place (not always decisive, because of the adaption of growth forms to dominating water energy; see Fig. 12.2). • Sedimentary structures (e.g. small scale cross-bedding, imbrication). Basic classification: Plumley et al. (1962) distinguished five major limestone categories I to V, representing a grading spectrum from quiet water to strongly agitated water deposits. The arbitrary boundaries between these types are determined by semiquantitative and qualitative textural criteria. Each major limestone type is subdivided into three subtypes (I-1, I-2, I-3). These subtypes indicate genetic similarities among limestones with different textural criteria, and genetic differences among limestones with similar textural criteria. Type I represents the end of the spectrum with minimum water agitation. The principal feature of these quiet-water limestones is the lack of recognizable transported particles. Subtype I-1 (Pl. 44/1) is characterized by conspicuous argillaceous admixtures. Subtypes I-2 (Pl. 45/2) and I-3 (Pl. 119/1) are relatively pure carbonates (< 15% clay). Subtype I-2 is essentially nonfossiliferous; subtype I-3 corresponds to a fossiliferous limestone. Fossil assemblages are simple and consist of many individuals but few species. Fossils are not transported and not rounded; many fossils are unbroken.
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Type II, representing alternate deposition in agitated and quiet water, is characterized by a mixture of quietand agitated water criteria. Microcrystalline calcite matrix comprises more than 50% of the rock. The admixture of > 50% transported grains is considered as evidence for intermittent water energy. Fossil assemblages are more highly complex than those of Type I. Subtype II-1 (Pl. 44/2) and II-2 (Pl. 44/3) differ in the size of transported grains (fine- to medium-grained grains and coarse- to very coarse transported grains, respectively). Grains are subangular to rounded. Subtype II-3 (Pl. 18/2) is characterized by alternating laminae composed of lime mud and coarser clastic particles. Clay content is below 15%. Type III describes deposition in slightly agitated water. The size of the transported grains varies from silt through fine-grained sand. Subtype III-1 (Pl. 118/ 1) corresponds to a calcisiltite, subtype III-2 to a very fine-grained calcarenite and subtype III-3 (Pl. 121/1) to a fine-grained calcarenite. Matrix (homogeneous microcrystalline in III-1 and III-2; inhomogeneous and varying in size) may constitute < 50 percent of the rock bulk. Type III includes nonfossiliferous, sparsely or moderately fossiliferous limestones (III-1 and III-2) as well as limestones with abundant and diverse fossils (III-3). Grains are subrounded to rounded. No conspicuous clay content. Type IV indicates deposition in moderately agitated water. The limestones are medium-grained to very coarse-grained calcarenites. The matrix is poorly sorted. Fossils are more common than in Type III. Subtypes (IV-1: Pl. 122/1; IV-2: Pl. 118/3; Fig. 4.32; IV-3: Pl. 45/3) differ predominantly in the grain size of transported particles which are generally broken and abraded fossils. Grains are subrounded to well rounded. No conspicuous clay content. Type V characterizes deposition and growth in strongly agitated water. Fossil fragments and rock debris are coarse-grained and lie within the range of calcirudites. Subtype V-1 (Pl. 115/2) corresponds to a limestone consisting of subrounded to rounded gravelsized fossils and rock fragments. Subtype V-2 (Pl. 9/7) is described as being ‘composed predominantly of conglomeratic and brecciated rock fragments’. Grains are angular to well rounded. Some clay may be present. Subtype V-3 corresponds to a ’reef’ limestone with in situ frame-building organisms’. Modified energy classification: While studying Triassic limestones of Bulgaria, Catalov (1972) differentiated between the energy index log of the energy minima and the energy index log of the energy maxima. These logs are obtained by connecting the respective
Energy Index Classification
points of EI logmax and EI logmin of a geological section. The field between these points is called the differential energy index field (DEI field). It describes the textural variation of carbonate rocks. A wide DEI field indicates large variations, a narrow field smaller variations. In contrast to Plumley et al. (1962), Catalov distinguished only four limestone types: Type I (deposited in quiet water), type II (deposited in slightly agitated water), type III (deposited in moderately agitated water), and type IV (deposited in strongly agitated water). The differences are indicated by different degrees of winnowing as well as frequency, size, sorting and roundness of grains; type and frequency of terrigenous admixtures, fossil associations and substrate (muddy, sandy). The main difference between the classification of Plumley et al. and Catalov is that the frequency and size of grains as well as fossils are much more pronounced in Catalov’s classification. Discussion: The original and the modified Energy Classification have proved successful in describing lithologic variations in time and space, changes in energy levels during time as well as in correlating outcrops and boreholes (e.g. Bolliger and Burri 1970; Skupin 1973; Rasser 1994). Assessment of subgroups may sometimes be difficult, if information on non-carbonate contents of the limestones is missing. Pitfalls in the use of the Energy Index can arise because • the size of grains depends on water energy, but also on the original size ranges (e.g. of fragmented green algae), the growth rates of organisms (e.g. in oncoid formation) or the cementation stages of grains, • the abundance of detrital quartz and clay is not necessarily dependent on water energy. Quiet-water carbonates may exhibit much higher quartz contents than suggested by Plumley et al., due to eolian input, • the occurrence of microcrystalline calcite matrix does not prove a priori quiet-water conditions, because mud can be fixed by organisms also in high-energy settings as well, • the use of ‘characteristic fossils’ in determining energy index types is dangerous as long as only groups or higher taxonomic categories are considered. Energy levels and resulting limestone types are not necessarily related to water depth: Quiet-water Type I limestones may occur in either deep (e.g. basinal) or shallow (e.g. lagoonal) water. Intermittently agitated carbonates of Type II should be more common in the transition zone between deep water (depositional interface below wave base) and very shallow water, where
Paleocurrents
the depositional interface is constantly above wave base. Agitated-water sediments corresponding to Types III, IV and V are more typical of the shallow water of shelf areas, reef and bank settings as well as coastal regions. The differences between the more generalized Plumley et al. classification and the more detailed Energy Index Classification developed by Catalov (1972) point to the necessity of revising the energy index concept by considering those microfacies criteria that indicate the degree of energy levels at the depositional interface with reasonable safety. An attempt in this direction was made by Bissell and Chilingar (1967), who combined the Energy Index Classification with the grain/micrite ratio, proposed by Leighton and Pendexter (1962) as well as with grain percentages. The grain-micrite ratio is the sum of the percentages of all grains divided by the percentage of micrite. Facies number: Based on the system proposed by Plumley et al. (1962), Grötsch (1993) introduced ‘facies numbers’ characterizing the degree of water energy and subdividing the mudstone to grainstone spectrum. Facies numbers are successfully used together with compositional data in statistical analysis, e.g. of cyclic carbonates (Grötsch et al. 1994).
12.1.1.3 Paleocurrent Data Currents are major factors influencing the distribution of sediment and of organisms. Global currents control the distribution of larvae and of nutrients and water temperature (scale: up to >1000 km). Currents control the geometry of reefs (e.g. spur and groove structures), influence the growth forms of reef builders (scale: about 10 to 1000 m, lateral; up to 100 m, vertical) and assist in the transport of sand-grade sediment. Small-scale currents are responsible for the transport of nutrients and waste, and control the settlement of larvae (scale: 0.1–1.0 cm). Smallest scale currents are produced by filtering organisms (scale: 0.1–10 cm). Both small-scale and smallest-scale currents may also be common in lowenergy ‘quiet-water’ environments. Paleocurrent indicators used in facies studies of carbonate rocks are essentially the same as those applied to siliciclastic rocks (Potter and Pettijohn 1977; Miall 1990). Criteria allowing the existence and directions of paleocurrents to be interpreted may indicate only the sense of current flows (groove marks, channel structures, parting lineation) or both the sense and direction of currents (cross-bedding, cross-lamination and ripple marks, flute marks, slump folds; orientation of fossils
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on bedding planes). Paleocurrent analysis is primarily an outcrop study, but some textural criteria can also be studied in thin sections of oriented rock samples. Paleocurrent patterns of ancient shoreline grainstones based on cross-bedding data were described by Knewtson and Hubert (1969), and Stricklin and Smith (1973). The orientation of clasts (Pl. 16/4, 7; Pl. 18/2; Pl. 139/8) and fossils are useful criteria (Sect. 5.1.2). Look for microphotographs showing the orientation patterns of calcareous algae (Pl. 57/7, Pl. 58/1), foraminifera (Pl. 74/5, 8; Pl. 113/1), sponge spicules (Pl. 78/2; Pl. 112/1), mollusk shells (Pl. 18/1; Pl. 102/1, 6; Pl. 113/2), and tentaculitids (Pl. 91/3, 4). Basics: Hydrodynamic controls Bollinger, W., Burri, P. (1970): Sedimentologie von SchelfCarbonaten und Beckenablagerungen im Oxfordian des zentralen Schweizer Jura. Mit Beiträgen zur Stratigraphie und Ökologie. – Beiträge zur Geologischen Karte der Schweiz, Neue Folge, 140, 1-96 Catalov, G.A. (1972): An attempt at energy index (EI) analysis of the Upper Anisian, Ladinian and Carnian carbonate rocks in the Teteven Anticlinorium (Bulgaria). – Sedimentary Geology, 8, 159-175 Grötsch, J., Koch, R., Buser, S. (1994): Fazies, Gildenstruktur und Diagenese des nördlichen Randes der Dinarischen Karbonatplattform (Barreme-Apt, W-Slowenien). – Abhandlungen der Geologischen Bundesanstalt Wien, 50, 125-153 Hamblin, W.K. (1969): Marine paleocurrent directions in limestones of the Kansas City Group (Upper Pennsylvanian) in Eastern Kansas. – Geological Survey of Kansas Bulletin, 194, 3-25 Skupin, K. (1973): Stratigraphy and microfacies in the crinoidal limestones (Trochiten limestone, Triassic) of SW Germany. – Sedimentary Geology, 9, 1-19 Plumley, W.J., Risley, G.A., Graves, R.W., Kaley, M.E. (1962): Energy index for limestone interpretation and classification. – In: Ham, W.E. (ed.): Classification of carbonate rocks. – Anerican Association of Petroleum Geologists, Memoir, 1, 85-107 Potter, P.E., Pettijohn, F.J. (1977): Paleocurrents and basin analysis. 2nd edition. – 425 pp., Berlin (Springer) Riedl, R. (1969): Marinbiologische Aspekte der Wasserbewegung. – Marine Biology, 4, 62-78 Stearn, C.W. (1982): The shapes of Paleozoic and modern reef builders: a critical review. – Paleobiology, 8, 228-241 Further reading: K164, K165
12.1.2 Storms Storms leave distinct facies criteria in marine coldwater carbonates (examples in James and Clarke 1997) and warm-water carbonates as well as in coastal lake sediments (e.g. Liu and Fearn 1993). The distribution and taphonomic characters of fossils may be indicative of storm events, because of a specific response of animals and plants to hurricanes (Walker et al. 1991).
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High-energy storms generate waves and currents that affect sea bottoms, stir up sediment and create characteristic sedimentary structures (e.g. hummocky crossstratification) and textures (e.g. tempestites). Severe storms erode coasts, destroy tropical islands and reefs, take part in onshore and offshore sediment transport and may contribute to extensive supratidal sedimentation. Shallow-marine waters respond to wind energy by • coastal swell (storm floods) resulting from upland winds that raise the sea level, • surface waves producing breakers at the coast, and • ground waves that become transformed into onshore and offshore bottom currents. Severe storms are represented by cyclones and hurricanes, occurring in tropical latitudes, and by intense winter storms (blizzards) in middle and high latitudes forming at fronts between cold and warm air masses. The development of strong storms is controlled by changes in atmospheric pressure and wind intensity. The importance of storm deposits in the geological record has attracted the attention of many workers. The papers listed in Box 12.2 describe the effects of modern storms known as cyclones, hurricanes and typhoons. A cyclone is an intense cell of low atmospheric pressure with a central eye surrounded by a circular wind system that rotates clockwise in the Southern Hemisphere and anti-clockwise in the Northern Hemisphere. The system follows tracks that generally curve away from the equator. Tropical cyclones form over the tropical oceans and range from 100 to 1000 km in diameter. Wind velocity may increase to up 230 mph and more. Tropical cyclones transport fine sediments and nutrients from shore to reef zones. The term typhoon designates a tropical cyclone occurring in the Pacific region. A hurricane is a tropical cyclone in which the wind speed equals or exceeds 118 km/h and may reach more than 300 km/h as in the example of Hurricane Gilbert described below. Locations crossed by the central eye experience winds from different directions that generate storm surges causing catastrophic damages within a zone of approximately 30 km on either side of the eye. The varying intensity, size and strength as well as the forward speed and duration of strong storms are critical constraints on the storm’s capacity to modify shallow-marine environments (Boss and Neumann 1993). Impacts from hurricanes are proportional to the area and the period over which hurricane conditions persist. Slower-moving hurricanes will cause more extensive and geologically relevant surface damage and
Storm Deposits
Box 12.2. Selected references discussing the impacts of hurricanes on organisms and sediments in modern shallow-marine carbonate environments. Frequent hurricanes occur today in belts 7° to 25 °N and S of the equator. Major reef areas suffering regular hurricanes occur in the northwestern and southwestern Pacific, the northern and southern Indian Ocean, and the northwestern Atlantic Oceran including the Caribbean Sea. Reefs: Many papers compare pre- and post-hurricane situations of specific reefs. A recent overview was given by Scoffin (1993). The list here includes only papers not cited by Scoffin: Blanchon and Jones 1997; Braithwaite et al. 2000; Davis and Hughes 1983; Gagan et al. 1987, 1988, 1990; Guozhong 1998; Heap et al. 1999; Hubbard 1992; Kaufman 1983; Knowlton et al. 1988; Kobluk and Lysenko 1987; Lugo-Fernandez et al. 1994; Moran and Reaka-Kudla 1991; Rogers et al. 1982; Russ and McCook 1992; Schönberger 1989; Woodley 1981; Yamano et al. 2001; Young et al. 1999. Carbonate platforms and ramps: Aigner 1985; Boss and Neumann 1993; Bourrouilh-Le Jan 1979, 1980, 1998; Davaud and Strasser 1984; Galli 1989; Hine 1982; Seguret et al. 2001; Vermeer 1963; Wanless et al. 1988. Beaches: Ball et al. 1967; Coch 1994; Collins et al. 1999; Davis et al. 1982; Knowles and Davies 1991; Lee et al. 1994; Siringan and Anderson 1994.
sediment redistribution than the quick passage of a weakened hurricane. The major outcomes of severe storms are (a) transport and deposition of storm sediments on shelves and export to slope and basinal environments, and (b) the destruction of reef structures. These processes are discussed in the following two sections.
12.1.2.1 Storm Deposits (Tempestites) on Shelves, Ramps and Platforms Storm beds are most common on shelves and ramps in windward settings, because of the generation of sediment-laden storm surges that transport sediment from onshore to deeper offshore outer ramp settings via unidirectional return flows. After the storm is over, the sediment is redistributed across the ramp and finally deposited in the form of sheets of limited extension in low-energy environments below storm-wave base (Fig. 12.3). These storm deposits exhibit distinct proximalitydistality trends as shown, for example, in the Arabian Gulf. Storm deposits of attached platforms (e.g. Florida) display strong variations in internal composition. The effects of storms on isolated platforms (e.g. Bahamas) are concentrated at platform margins (Hine et al. 1981). Sediment transport and redeposition produce millimeter- to meter-thick storm layers, contribute to the
Tempestites
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Fig. 12.3. Storm processes and storm beds. Major storms will interact with the seafloor down to the storm wave base occurring in depths of several tens of meters. Storms directed onshore lead to a buildup of water masses and, when the storm abates, to the generation of offshore-directed return flows causing bottom currents that transport sediment down below the storm wave base into low-energy depositional environments forming tempestite beds. Shoreface sediment put into suspension can be transported offshore by density currents down to beyond storm wave base. The figure displays storm processes affecting ramps and platforms. Reefs are different. After Tucker and Wright (1990).
formation of lag deposits, and change the distribution patterns of benthic organisms. The effect of storms on sedimentation is controlled by water depths and the locations of the fair-weather wave base and the stormweather wave base. Storms set up waves and currents below the fair-weather wave base, cause reworking of sediment and develop characteristic bed forms. Onshore-directed storms generate currents leading to the buildup of water in nearshore regions or in lagoons. During the decline of storm intensity, offshore-directed bottom currents transport shallow shoreface sediment to deeper offshore settings (Fig. 12.3).
Tempestites Limestone and sandstone beds deposited by storms are called tempestites (Kelling in Ager 1974). The base of these storm beds is sharp and erosional, and exhibits a variety of sole marks. Common internal structures include accumulations of shells, graded bedding, flat bedding, parallel and cross lamination. Ripple bedding and burrowing may be common at the top of the layer. Proximal and distal tempestites: Storm beds exhibit much variation in thickness, grain size and internal structures, depending on the proximity and the intensity of the storm (Aigner 1985; Myrow and Southerd 1996). Similar to turbidites, tempestite bed thickness varies proportionally with the maximum grain size of the bed. Proximal storm beds are thick-bedded, bioclast-dominated and coarse grained, and include many amalgamated and composite beds containing ecologically and stratigraphically mixed biota. The base of the beds is an erosion surface. Distal tempestite beds tend to become thinner, finer, rarer, better preserved and more biogenic in composition toward deeper depositional environments. They are
commonly mud-dominated. The base of the distal tempestite beds is plane. Proximal tempestites often contain many shell layers, distal tempestites only few. Criteria of ancient tempestites: The following text summarizes field and laboratory criteria denoting the storm-related deposition of carbonate sediments. The facies analysis of ancient storm deposits requires the integration of lithofacies, biofacies and microfacies. Diagnostic features of carbonate tempestites can be derived from observations of modern storm sediments. Some of the criteria listed in the following text are often difficult to recognize in ancient calcareous storm deposits, e.g. erosive soles and rippled tops of the tempestite beds, because of bioturbation and diagenetic overprint. Sedimentary structures and sequences of tempestites: • Tempestite beds are intercalated in marly, shaly or carbonate beds. They are common in shallowing-up sequences. Storm beds range in thickness between a few millimeters and several tens of centimeter. Thickness often differs markedly from that of underlying and overlying beds. • Storm waves and currents deposit layers of packstones and grainstones often exhibiting characteristic undulated bedding. This hummocky cross stratification (HCS; Harms et al. 1975; Swift et al. 1983; Duke 1987) shows gently curved, low-angle cross lamination and laminae with convex-upward (hummock) and concaveupward (swale) developed curvatures. Hummocky cross-stratified pack- and grainstones occur as 0.1 to about 2 m thick units interbedded with lime mudstones and wackestones or shales. Many storm sediments form amalgamated beds. Hummocky limestone beds exhibiting packstone and grainstone textures are common in
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outer ramp settings (Markello and Read 1981; Kreisa 1981; Wu 1982; Aigner 1982; Wright 1986). • Base of storm beds: The base is erosional and sharp. Erosive contacts of tempestites composed of packstones with the underlying bed indicate erosion of the sea bottom (commonly mudstone and wackestone; see Pl. 102/ 1, 3). Sharp and plane boundaries between graded packstone of tempestites and the underbeds point to reworking by storms followed by rapid transport of reworked material into muddy areas (Pl. 102/2). • Sole marks: Channels and scours cutting into finegrained limestones and filled with skeletal packstones and grainstones are often associated with hummocky beds. Infilling material is heterogeneous and unsorted. Gutter casts (U-shaped depressions) indicate paleocurrent directions. • Burrowing is common at the top of storm beds. Strong bioturbation indicates breaks in storm activity. • Internal structures: Storm beds may exhibit characteristic vertical sequences (Dott and Burgeois 1982; Duke 1985; Aigner 1985; Monaco 1992). A complete storm wave sequence consists of a basal lag unit characterized by an erosional base and coarse-grained resediments and skeletal grains, overlain by parallellaminated, hummocky, flat-laminated, cross-laminated and mudstone units (Pl. 102/2). This sequence reflects deposition during waning flows. Note that this sequence is an ideal one that is by no means developed in all tempestites and differs for proximal and distal parts. Proximal storm layers comprise beds consisting from bottom to top of an erosional base, parallel to low-angle lamination with minor internal discordances, wave ripples, and a mud layer. A common variation is an erosive basal coarse-grained lag layer, a central well-sorted and matrix-poor calcisiltite or fine-grained calcarenite layer, and an upper fine-grained rippled calcilutite layer. Distal storm layers exhibit erosive or non-erosive bases, internal flat layers, rare grading, and parautochthonous accumulations of shells. Microfacies criteria • Distinct differences in grain sizes of tempestites and of under- and overlying beds (grain- and packstones intercalated within mudstones or wackestones). The primary grain-size spectra may be obscured due to bioturbation. • Graded bedding is often common. The graded parts of the tempestite beds are commonly succeeded by finely laminated bioclastic packstones and wackestones (Pl. 102/3) which in turn are overlain by dark lime mudstones or marls. • Parallel lamination associated with ripple laminated intervals in beds with hummocky cross stratification indicate a high current regime during the deposition of
Tempestites: Microfacies
the tempestite. Parallel lamination may indicate unidirectional flows (Pl. 102/3). • Millimeter- to decimeter-sized micrite clasts and intraclasts (Pl. 6/1) in wackestones, packstones and grainstones may indicate redeposition of storm-derived material. • Lithoclasts comprising distinctly different microfacies, e.g. fenestral bindstones and coral boundstones, pointing to high-energy events and erosion in peritidal and reefal environments. • Reworked black pebbles (Pl. 16/1) can indicate storm-induced coastal erosion. • Non-carbonate extraclasts may represent storm-derived material (Pl. 16/4). • Flat-pebble conglomerate fabrics. Clast-supported calcirudites composed of coarse flat-shaped intraclasts typically occur in subtidal environments together with storm deposits, but also originate in storm-affected tidal environments (Pl. 125/3). Many examples have been described from Cambrian and Ordovician storm deposits. The preservation of flat-pebble structures is facilitated by early lithification enhanced by microbial activity (Ito and Matsumoto 2001). • Edgewise conglomerate fabrics (Pl. 58/1) may indicate storm-influenced sedimentation. Fossils in tempestites • Distinct quantitative differences exist in the composition of tempestite beds and under- and overlying beds (Stel 1975). • Skeletal concentrations are common features of many modern and ancient carbonate and siliciclastic shelves. Diagnostic criteria of storm-controlled bivalve shell layers are mixtures of epifaunal and endofaunal elements, bimodal sorting of complete shells and comminuted shell debris (Pl. 102/5), random orientation, packstone and grainstone matrix between the shells and an erosive base of the millimeter- to decimeter-thick shell layers (Pl. 102/1, 3). The complete or disarticulated shells are well preserved (Pl. 17/4). Abrasion, bioerosion and encrustations are absent. Distal tempestites often exhibit telescoped mollusk valves (Pl. 102/5). Common fossil groups contributing to the formation of storm-related skeletal concentrations are bivalves, brachiopods and crinoids. • Trace fossils and different types of bioturbation in storm-induced beds provide clues to the estimation of the depositional water depths of tempestites (Wu 1982; Aigner 1985). Bioturbation is commonly strong in the upper part of storm beds. The microfacies of tempestites is characterized by packstone textures, densely packed grainstones and rudstones. Abundant grains are reworked bioclasts and
Ancient Storm Deposits
lithoclasts. Shell beds are common. Frequent microfacies types are • Non-graded, poorly sorted skeletal packstones and densely packed grainstones and rudstones, consisting of coarse skeletal debris, e.g. shell accumulations (bivalves, brachiopods; Pl. 102/1) or accumulations of crinoids (Pl. 95/2). The limestone is often dark-colored. • Graded skeletal packstones, grainstones and rudstones consisting of skeletal debris including mollusks, brachiopods and echinoderms. • Biolithoclastic graded, densely packed floatstones and rudstones with coarse fragments of mollusks and echinoderms, together with centimeter-sized intraclasts. • Lithoclastic packstones, rudstones and densely packed floatstones. A common microfacies of proximal tempestites is characterized by angular micritic clasts resulting from reworking of beds underlying the storm bed. • Thinly-laminated bioclastic mudstones and wackestones, sometimes with pelsparite laminae and graded fine shell debris (Pl. 102/2). Some of the characteristic microfacies criteria of tempestites are shown in Plate 102. Phanerozoic storm patterns: Indications of stormrelated sedimentation are abundant in Phanerozoic limestones formed on ramps and in mixed carbonate-siliciclastic environments (Box 12.3). Fluctuations in the number and intensity of storms during the Phanerozoic have been suggested, based on global paleogeographical distributional patterns of storm deposits (Marsaglia and Klein 1983), paleoclimatic models (Parrish and Curtis 1982; Barron 1989) and variations in the wavelength of hummocky cross stratification (Ito 2001). Surprisingly, the overlap between the spatial distribution of ancient reefs with large amounts of fore-reef detritus and inferred paleowind directions is rare for most of the Phanerozoic except for the Late Triassic (Kiessling 2002). Event Deposits and Eventstones Tempestites are a class of event deposits. Event deposits include tempestites, turbidites, mass flow deposits, inundites (flood deposits) as well as seismites described in Sect. 12.4 (see Seilacher 1991 for a discussion of concepts). Event beds are the result of shortterm sedimentation processes that produce characteristic sedimentary signatures. These signatures may bear considerable similarities. Differentiation of tempestites and turbidites: Storm deposits are physically similar to deposits from other
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Box 12.3. Selected case studies of storm deposits in ancient carbonate environments. Most papers deal with tempestites in ramp settings. Tertiary: Martin et al. 1993 Cretaceous: Banerjee 1981; Chel and Leckie 1992; Garcia-Ramos et al. 1989; Hunter and Clifton 1982; James 1979, 1980; Valecka 1984 Jurassic: Ager 1974; Badenas and Aurell 2001; Fürsich and Oschmann 1986, 1993; Galli 1990; Ghosh et al. 1986; Hamblin and Walker 1979; Hüssner 1985; Molina et al. 1997; Monaco 1992; Moore et al. 1992a, 1992b; Specht and Brenner 1979 Triassic: Aigner 1979, 1982, 1985; Aigner et al. 1978; Aljinovic 1995; Blendinger and Blendinger 1989; Calvet and Tucker 1988; Duringer 1984; Galli 1989; Hagdorn 1982; Hips 1998; Masetti et al. 1988; Michalik 1997; Michalik et al. 1992; Perez-Lopez 2001; Röhl 1986; Török 1993 Permian: McKie 1994 Carboniferous: Ball 1971; Buatois and Mangano 1994; Engelbrecht 2000; Graham 1982; Handford 1986; Kelling and Mullin 1975; Reynolds 1992; Simpson 1987; Stevens 1971; Wright 1986; Wu 1982 Devonian: Dreesen 1982; Gischler 1995; Hubmann 1999; Kazmierczak and Goldring 1978; Lavoie 1992; Miller et al. 1988; Rehfeld 1986; Stel 1975 Silurian: Brunton and Copper 1994; Cotter 1990; Johnson 1989; Sami and Descrochers 1992 Ordovician: Bakush and Carozzi 1986; Brenchley and Newall 1982; Brookfield and Brett 1988; Drummond and Sheets 2001; Jennette and Pryor 1993; Kreisa 1981; Lee 1988; Lee and Kim 1992; Lehman and Pope 1989; Paik and Lee 1989 Cambrian: Liang and Friedman 1993; Markello and Read 1981; Mount 1982; Mount and Kidder 1993; Sepkoski 1962
genetically unrelated processes, such as turbidity or slump events. Turbidite sedimentation may be storminduced (Wu 1982). A table summarizing the distinguishing criteria of tempestites and turbidites (Einsele and Seilacher 1991) as well as a discussion of diagnostic thin-section criteria can be found in Sect. 15.7.2.5. Erosional sole marks, gradation, lamination and ripple bedding occur in both types of event beds, but differ in their frequency. The main discriminating criteria of proximal tempestites and proximal turbidites are biofacies, fossils and stratigraphic context. Eventstones: The term designates carbonate sediments whose deposition took place in rapid short-term events as compared to the long-term average background sedimentation (Hüssner 1985). Eventstones include tempestites and turbidites and imply the existence of a relief and gravitative sediment transport (turbidites) or of episodically high water energy (e.g. storms; tempestites).The important criteria of event-
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Tempestites
Plate 102 Storm Deposits (Tempestites) Limestone and sandstone beds deposited by storms are called tempestites. These storm deposits form beds and lenses intercalated in marls and shales or in limestones of different texture and microfacies. Tempestites are common in subtidal ramp sediments. The plate displays fabrics of tempestites formed under high-energy (proximal) and low-energy (distal) conditions. Proximal tempestites (–> 1, 3, 4) are often bioclast-dominated and coarse-grained, and composed of many amalgamated and composite beds. Distal storm beds (–> 2, 5, 6) are thin-bedded and mud-dominated. The pictures 1–3 and 5 on this plate show thin sections of a case study. The Late Triassic limestones formed in the intraplatform setting of the Kössen Basin (Northern Calcareous Alps) contain many limestone beds interpreted as tempestites (Kuss 1983). The samples are from a sequence consisting of alternating marls and limestones. Sedimentation took place in subtidal low-energy environments of a protected platform influenced by storm events. Some of the limestone beds exhibit erosional sedimentary structures (gutter casts and pot casts), cross bedding, lamination and ripple marks as well as distinct accumulations of shells. The latter occur together with thin, fine-grained peloidal grainstone layers. Field criteria and microfacies suggest high- and low-energy storm sedimentation. High-energy deposition is indicated by the deposition of complete and disarticulated shells (–> 4). Low-energy deposition by the lamination textures of fine-grained peloidal grainstones (–> 4, 6). The shell accumulations (lumachelles, coquinas) are caused by storm-induced bottom flows. 1 Tempestite. Brachiopod coquina formed by storms. An interpretation as proximal tempestite is supported by the preservation of complete shells, the current orientation of shells as well as a sharp lower erosive boundary and an irregular upper boundary. Brachiopod shells are geopetally infilled with micrite. The lower part of the figure shows a mudstone matrix with parallel laminae composed of very small peloids and thin compacted shells. The joint occurrence of these low-energy laminae and the high-energy fabric of the coquina layer points to rapidly changing depositional conditions. Late Triassic (Kössen Formation, Rhaetian): Loferbachgraben, Salzburg, Austria. 2 Distal tempestite bed intercalated within calcimudstone (CM). Both the lower and the upper surface of the tempestite are flat. Slight compositional differences occur from bottom to top: a - densely-packed wackestone with abundant recrystallized involutinid foraminifera (white) and small peloids; b - wackestone with bivalve shells, involutinid foraminifera and peloids; c - wackestone with sparse bivalve shells and peloids. The involutinid foraminifera indicate sedimentation as platform-derived sediment. The low-angle cross-bedding boundary between a and b (arrows) suggests the effect of strong currents in the upper flow regime. Same locality as –> 1. 3 Proximal tempestite. Storm layer intercalated in lime mudstone. A cross-bedded coquina (CBC) at the base is overlain by thin parallel laminae (GL) grading up into mudstone (M). The discontinuous lens-like geometry of the laminae suggests deposition from unidirectional flows. The arrows point to the distinct erosional basal boundary. Note the chaotic texture of the upper coquina consisting of disarticulated bivalves and gastropods (G). Broken shells point to deposition of material that has been fragmented prior to the storm event. Late Triassic (Kössen Formation, Rhaetian): Gaissau near Hallein, Austria. 4 Storm layer intercalated within bioclastic shelf carbonates adjacent to organic mud mounds (see Pl. 143). The tempestite layer differs from the background sediment by the composition and grain size. From bottom to top: a - bioclastic wackestone (below and above the tempestite layer) with fine-grained skeletal debris consisting of ostracods, styliolinids, trilobites and echinoderms, b - coarse-grained echinoderm packstone representing the lower part of the tempestite layer. Arrows point to syntaxial rim cement around echinoderm fragments. c - bioclastic packstone with echinoderms and tabulate corals (TC). Note the stromatactis-like fabric (S) at the upper boundary of the storm layer. Compare Pl. 94/7 and Pl. 143. Early Devonian (Emsian): Erfoud, southern Morocco. 5 Concentration of thin-shelled bivalves and a few gastropods within the distal tempestite bed. Densely-packed, wellsorted, bioclast-supported fabric. The carbonate mud is only preserved as post-storm infillings between the bioclasts. Compare the very similar fabric formed by trilobite shells shown in Pl. 94/7. Late Triassic (Kössen Formation, Rhaetian): Steinplatte, Tyrol, Austria. 6 Uppermost part of a thin distal tempestite layer. The parallel orientation of the shells indicates control by weak bottom currents. Note similarities with –> 2. Middle Triassic (Hauptmuschelkalk): Garnberg, Künzelsau, Germany. –> 1–3: Kuss 1983
Plate 102: Storm Deposits (Tempestites)
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stones are (a) mixing of grains derived from source areas that may be different with regard to facies and age, and (b) vertical sequence of sedimentary structures and textural types indicating changes in energy levels during deposition.
12.1.2.2 Impact of Tropical Storms on Reefs Geologists like to relate the origin of allochthonous reef material, e.g. fore-reef breccias (Pl. 27/2) or isolated reef boulders found in slope position to catastrophic storms devastating the reefs. This is a pleasant interpretation, but how do severe storms actually affect reefs and how can we trace their activity in the geological record? Modern tropical storms and hurricanes cause significant perturbations that affect reef communities and destroy reef architecture (Hughes 1993; Scoffin 1993). Damage by hurricanes is common in shallow reefs. The effect of hurricanes and cyclones may be significant over tens to hundreds of square kilometers, but reef destruction is often patchy. Hurricane waves and storm surges erode reef crest, corals, reef rocks and sediments down to more than 20 m depths and produce mud- to boulder-sized fragmented reef material. This material accumulates • as talus on and at the foot of the forereef slope, in grooves on the reef front and on submarine terraces, • as storm ridges consisting of shingle, boulders or single blocks deposited on intertidal reef flats, • as beach ridges formed by leeward migrating shingle, • as scattered debris in back-reef lagoons, and • as drapes of carbonate sand and mud in forereef and in lagoonal areas. Common and in the geological record recognizable results of strong storms are the disintegration of reef framework, coral breakage, dislodgment of heads of massive corals, transport of variously-sized reef blocks and of carbonate sand, and scarring on standing corals caused by waterborne debris. The extent of damage that will in turn influence the chance of recognizing storm effects in ancient reefs depends on many factors including • the morphology and size of the reefs, • the violence of strong storms relative to fair-weather effects and the state of tides, • the angle under which the cyclone affects the reef slope, • the shelter afforded by differences in water depths, topography or by the structure of the reefs,
Impact of Tropical Storms
• the growth form, tendency of attachment, strength and community composition of reef-building organisms, • the length of time after a hurricane before a subsequent storm disturbs reef recovery. Recurring storms may selectively destroy fragile corals. This can result in gaps in otherwise distinct ecological zonations (Geister 1992). The recovery period of modern reefs (the time taken for the reef to return to pre-hurricane conditions) lasts between a few years and some tens of years. Scoffin (1993) reports numbers between 5 and 40 years. Resettlement is selective and depends on the available substrates. On soft muddy substrates resettlement often starts with sponges and soft corals, on hard substrates with cyanobacteria and sponges. Changes in diversity over time influenced by severe storms have been explained in the context of the ‘intermediate disturbance hypothesis’ (Connell 1978), suggesting that the highest number of corals species will be reached at intermediate levels of the frequency and size of natural disturbances. The effects of hurricanes on reef diversity vary with depths, reef zones, the time since the last disturbance of the ecosystem, and the morphology of dominant species (Rogers 1993). Application of the intermediate disturbance hypothesis to modern ancient reefs struck by severe storms requires data on pre- and post-hurricane species number and cover.
The Hurricane Gilbert - Cozumel Case Study In September 1988, Hurricane Gilbert, on its devastating path through the western Caribbean, passed directly across the Mexican Island of Cozumel, off the eastern shore of Yucatan Peninsula. The eye of the hurricane was situated directly over the northernmost tip of the island. The storm reached an extraordinary strength (about 320 km/h) and a minimum surface pressure (885 hPa), damaged coral reefs on the shelf and at the shelf margin, and also destroyed the tropical forest on the island (Sanchez et al. 1999). The shelf of Cozumel Island had been intensively studied only a few months before with regard to platform and reef morphology and distribution patterns of benthic organisms. Priority was given to the carbonate-producing fauna and flora, and the quantitative assessment of the shallow-water biotopes on the insular shelf (Muckelbauer 1990). With these detailed observations of the pre-hurricane situation at hand, a new survey was started (Muckelbauer 1991, unpublished research report, Institute of Paleontology Erlangen) in
Hurricane Gilbert
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Fig. 12.4. Impact of Hurricane Gilbert on a reef on the western shelf off Cozumel Island. A: Pre-hurricane situation. The reef slope is densely covered by scleractinian corals (Porites, Montastraea, Agaricia), octocorals, and some yellow sponges (in the foreground). The photograph was taken at a depth of about 15 m. B: Post-hurricane situation after the reef destruction by Gilbert. The photograph, taken on September, 18, 1989, almost exactly one year after the hurricane, shows the same location as Fig. A. Almost all former reef builders have been removed or fragmented. The number of hard bottom areas has dramatically increased. A part of the former reef area is now covered by coral debris (mainly Porites), gravel-sized rubble and meter-sized coral blocks (in the background). Large areas of the slope are coated by sand-sized sediment reworked from inter-reef terraces. This sand acted similar to sandblasting equipment, destroying the living reefs and scratching out newly formed free hard bottom areas. Resettlement is weak except for thin cyanobacterial mats growing on a few coral blocks. Photographs: Courtesy of G. Muckelbauer (Erlangen).
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Pre- and Post-Hurricane Comparison
Fig. 12.5. Cozumel Island east of Yucatan Peninsula, Mexico. A: Map showing the narrow shelf region. Most coral reefs are situated on the western shelf and at the shelf margin, only a few on the near-coast eastern shelf. B: Schematic profile of the western shelf of Cozumel Island. The reefs are located on the morphological steps (after Muckelbauer 1990).
A
B
order to compare the post- and pre-hurricane situations one year after Gilbert. Location: Cozumel is located approximately 18 km offshore the northeastern coast of Yucatan Peninsula. The narrow shelf (width 200 to 1500 m; Fig. 12.5A) is organized into steps and terraces (Fig. 12.5B). The insular slope drops down at very steep angles to depths of 400 m and more than 1000 m. The steps are situated at depths of about 4 m, 10 m, 20 m, and in places 30 m and 50 m. The terraces between the steps vary in extension from a few meters to some hundreds of meters. Most coral reefs occur on hard bottoms along the steps and at the shelf edge on the western shelf. The terraces are covered with calcareous sands and sea grass. Reef distribution is controlled by the topography of the shelf, which originated in the Pleistocene and was changed during the Holocene, by the distribution of hardground and softground substrates, and by strong ocean currents (2 m/sec) from southern directions representing an offshoot of the Caribbean Current. Pre- and post-hurricane comparisons: Comparisons of damages caused by hurricanes with the pre-hurricane situation have been made in the Caribbean and other regions (e.g. Endean 1976; Stoddard 1985; Mergner 1985; Mah and Stearn 1986; Hubbard 1992; Boss and Neuman 1993), but most reports are focused only on certain organism groups, or use different quantifying methods that can significantly affect the outcome of pre- and post-hurricane comparisons.
These drawbacks are avoided in the Cozumel study, which is based on identical methods both in the field and for the statistical evaluation of quantitative data. The studied locations were grouped to represent morphological sites of reef growth in different water depths (reefs I, II and III) at the steps between terraces 1, 2, and 3. Patch reefs and the upper insular slope were also studied. Using the photographic transect method (Dodge et al. 1982), numerous photographs taken along transects running parallel to the reef structures were evaluated by point-counter analyses. Corals were identified down to species, forms and even morphotypes. Algae could be classified to the generic or species level. Altogether a total of 95 taxonomic categories was distinguished, and the percentage of coverage was calculated for each one. The resulting data were summarized for the whole transect, thus obtaining the mean coverage as well as the standard deviation for different reefs. Coverage, diversity and species richness were computed by statistical tests. Post-hurricane changes: Hurricane Gilbert devastated the shelf-edge reefs and many smaller reefs located on the shallow shelf. The reef structures were destroyed and eroded. Large areas were abraded by finegrained calcareous sand flushing over the newly created bare hard bottoms. Broken corals and up to metersized reef blocks were transported onshore as well as offshore.
Hurricane Gilbert
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Fig. 12.6. Changes in the substrate, coverage and composition of shelf-edge and shelf reefs off Cozumel caused by Hurricane Gilbert. A: Overall changes. B: Changes in the coverage of calcareous and non-calcareous organisms.
The impact of the hurricane on substrate conditions, coverage and biotic composition are summarized in Figs. 12.6 and 12.7. The histograms give an idea of the overall changes in the reefs studied (Fig. 12.6) and of the variations in the changes in reefs located in different water depths (Fig. 12.7). Major changes with implications for the interpretation of ancient reef destructions relate to the frequency of hard and soft substrates, the distribution of sediments, and the abundance of carbonate-producing and other organisms.
Statistically significant results of the comparison of the pre-hurricane and the post-hurricane situation are: • The size and frequency of bare hard bottoms (reef rocks), not yet resettled by organisms have dramatically increased in shallow subtidal settings (Fig. 12.6A). These hardground areas were almost doubled in some reefs as a result of the hurricane (Fig. 12.7A). • The frequency of sand-sized loose sediment (soft bottoms) has increased in shelf-edge reefs (Fig. 12.7A)
Fig. 12.7. Changes caused by Hurricane Gilbert in shelf-edge reefs (III), shelf reefs (I and II, isolated patch reefs) and reefs on the upper slope (15-30 m). A: Coverage changes. B: Changes in the frequency and coverage of coral morphotypes.
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due to the current-transport of sediment from the shelf to the shelf margin during and after the hurricane. Shelf reefs show no notable changes in the frequency of soft sediments. • Carbonate-producing organisms are distinctly reduced in frequency (Fig. 12.6B). • Corals and hydrozoans are reduced to almost one half their pre-hurricane level (Fig. 12.6A). • Growth forms of corals: Branched corals exhibit a decisive reduction in all reef sites (Fig. 12.7B). The coverage of massive and platy corals has not been reduced except platy corals in shelf-edge reefs. • Considering all corals in regard to the relative coverage of all corals dendroid corals have lost out to massive corals. • Total coral diversity (species richness) has been maintained, but the ranking of species according to their frequency in the reefs has changed considerably. • Differences in coral diversity between reefs located in different water depths have become more pronounced due to great reductions in species dominating in prehurricane time. • Reefs in deeper water are more greatly destroyed than those in shallow water, due to the higher vulnerability of dendroid corals dominating in deeper water. • Different algal groups reacted differently: The coverage of brown algae has increased significantly, that of green algae decreased. Red algae show no significant change. • Calcareous algae have not declined quite as drastically as corals but still significantly. • The frequency of calcareous green algae has been variously reduced. Halimeda is reduced by one half. Other calcareous green algae show a slight increase. • Surprisingly, the coverage of calcareous red algae has not changed. • The coverage of cyanobacteria growing on transported coral fragments and dead coral heads has increased significantly. • The destruction of the non-calcareous sessile reef organisms (mainly sponges, octocorals, hydroids, soft algae) has been hit by the hurricane similar to that of the scleractinian coral fauna. • High-growing finger-, bowl-, tube- and vase-shaped sponges have clearly lost ground. Low-growing encrusting and boring sponges have gained in distribution both in near-coast reefs and in shelf-edge reefs, because many of these sponges live in protected cryptic habitats. The Cozumel study describes the situation just one year after the hurricane. Most of the changes observed reflect the biotic loss and shifts in the dominant taxa
Storm Deposits: Significance
rather than the reestablishment of the reef community. The shifts as well as the creation of large no longer organically covered areas show that the impact of hurricanes on reefs is not solely destructive, but that the ecosystem also benefits from severe storms. The initial recovery stage is indicated by the strong increase in the frequency and distribution of cyanobacteria. How to recognize storm effects in ancient reefs ? Box 12.4 reviews criteria that indicate disruption of reef growth by severe storms. Most criteria must be evaluated in the outcrop, but some should also be evident on a microfacies scale. 12.1.2.3 Significance of Carbonate Storm Deposits Severe storms acting in shallow-marine environments change sedimentation patterns, influence the distribution of organisms, and modify the course of diagenesis. Ancient storm sediments provide information on the following topics: • Depositional processes: Tempestites and reefs offer the possibility of understanding the physical controls on debris production. • Stratigraphy: The recognition of the imprint of hurricanes and other strong storms on the sediment record provides a useful stratigraphic marker for the study of ancient reef systems. • Paleoclimate: Because the frequency and the predominant tracks of intense hurricanes are expected to have varied in time as a function of global climatic changes (Emanuel 1987; Barron, 1989; Gray 1990), hurricane-related tempestites are used as paleoclimatic proxies. • Paleogeography: Storm-dominated, windward platform margins are characterized by sand bodies trending parallel to the platform margin. Windward-leeward orientations of ancient reef complexes can be deduced from distributional patterns of fore-reef sediments. These criteria as well as modeling of windward-leeward effects based on paleowind directions allow largescale and medium-scale paleogeographic maps to be reconstructed. • Diagenesis: Hurricanes are frequently accompanied by heavy falls of rain. Strong rain may trigger vadose dissolution and karstification processes of carbonate platforms (Bourrouilh-Le Jan 1998) and influence dolomitization in tidal flats (Bourrouilh-Le Jan 1979). • Reservoir rocks: Onshore directed storms can accumulate shoals and sheets of porous carbonate sands of reservoir quality along coastlines. Draping of storm-
Basics: Storms
Box 12.4. Diagnostic features of storm effects on ancient reefs. Criteria that have also become evident in the Cozumel study are marked by an asterisk. Reef detritus • Fore-reef detritus: Extended slopes and aprons consisting predominantly of reef-derived material are valuable indicators of storm effects on reefs growing in shallow waters. The limestones are breccias, rudstones or floatstones. Angular and rounded clasts are intermixed with isolated fossils. Marked differences in the microfacies of the clasts are due to different source areas of the eroded material. • Reef breccias, rudstones and floatstones deposited in near-coast position (*). • Large, variously sized coral blocks or reef blocks occurring within fine-grained lagoonal sediments (*). • Accumulation of fragments of corals or other reefbuilders acting as a base for encrusting organisms (*). Compositional biotic changes • Striking vertical and lateral differences of paleontological criteria on a decimeter or meter scale. • Conspicuous reduction of branching growth forms in favor of massive growth forms (*). • High-growing reefbuilders followed by a level of lowgrowing reefbuilders along a conspicuous boundary. • Abrupt differences in the taxonomic composition and frequency of reefbuilder associations (*). Discontinuities and textural changes • Discontinuity planes and fissures disrupting the reef structure. • Beds with abundant, poorly sorted coral debris overlain by microbial carbonates (*). • Conspicuous changes in the composition and grain size of the sediment deposited between and within reefbuilders (*). • Fine-grained fossil-free layers that may correspond to carbonate mud settled down after the abatement of storms. • Layers or lenses with high organic content can be caused by accumulation of destroyed plants.
induced porous skeletal sand waves over dense mound facies can form hydrocarbon traps, as shown by hydrocarbon reservoirs in capping beds of Carboniferous Waulsortian mounds. Basics: Storms Aigner, T. (1985): Storm depositional systems. – Lecture Notes in Earth-Science, 3, 174 pp., Stuttgart (Schweizerbart) Barron, E.C. (1989): Severe storms during earth history. – Geological Society of America, Bulletin, 101, 601-612 Bourrouilh-Le Jan, F.G. (1998): The role of high-energy events (hurricanes and/or tsunamis) in the sedimentation, diagenesis and karst initiation of tropical shallow water carbonate platforms and atolls. – Sedimentary Geology, 118, 3-36 Dott, R.H., Bourgeois, J. (1982): Hummocky stratification: significance of variable bedding sequences. – Geological Society of America, Bulletin, 93, 663-680
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Duke, W.L. (1985): Hummocky cross stratification: tropical hurricanes, and intense winter storms. – Sedimentology, 32, 167-194 Gischler, E. (1995): Current and wind induced facies patterns in a Devonian atoll: Iberg reef, Harz Mts., Germany. – Palaios, 10, 180-189 Harms, J.C., Southard, J.B., Spearing, D.R., Walker, R.G. (1975): Depositional environments as interpreted from primary sedimentary structures and stratification sequences. – Society of Economic Paleontologists and Mineralogists, Short Course, 2, 161 pp. Hubbard, D.K. (1992): Hurricane-induced sediment transport in open-shelf tropical systems – an example from St.Croix, U.S. Virgin Islands. – J. Sed. Petrol., 62, 946-960 Hughes, T. (ed., 1993): Special issue: Disturbance: effects on coral reef dynamics. – Coral Reefs, 12, 117-233 Ito, M., Ishigaki, A., Nihikawa, T., Saito, T.: (2001): Temporal variations in the wave lengths of hummocky crossstratification. Implications for storm intensity through Mesozoic and Cenozoic. – Geology, 29, 87-89 Myrow, P.M., Southard, J.B. (1996): Tempestite deposition. – Journal of Sedimentary Research, A, 66, 875-887 Nummedal, D. (1991): Shallow marine storm sedimentation - the oceanographic perspective. – In: Einsele, G., Riecken, W., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 227-248, Berlin (Springer) Pielke, R.A. (1990): The hurricane. – 228 pp., London (Routledge) Rogers, C.S. (1993): Hurricanes and coral reefs: the intermediate disturbance hypothesis revisited. – Coral Reefs, 12, 127-137 Scoffin, T.P. (1993): The geological effects of hurricanes on coral reefs and the interpretation of storm deposits. – Coral Reefs, 12, 203-221 Seilacher, A., Aigner, T. (1991): Storm deposits at the bed, facies, and basin scale: the geologic perspective. – In: Einsele, G., Riecken, W., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 227-248, Berlin (Springer) Shinn, E.A., Steinen, R.P., Dill, R.F., Major, R. (1993): Limemud layers in high-energy tidal channels: a record of hurricane deposition. – Geology, 21, 603-605 Further reading: K032, K165
12.1.3 Marine Carbonate Substrates The composition and type of carbonate substrates depend on depositional, biological and diagenetic factors. The substrate type affects feeding modes as well as the penetration and attachment of benthic organisms. Limestones represent former soft substrates and former hard substrates and various modifications of these two basic categories. Soft bottoms consist of loose grains exhibiting different consistencies depending on the grain size and water content. Hard bottoms are immobile substrates that are firm or hard enough to take encrusting and boring organisms (Sect. 5.2.4.1 and Sect. 9.2). They represent a non-depositional surface and are resistant to submarine erosion. Differences in the states of con-
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sistency and stability of soft and hard substrates have been designated by the terms soupground, softground, looseground, firmground, hardground, rockground (rocky sea bottom), and shellground (Ekdale 1985; Goldring 1995). Softground refers to a muddy sediment, looseground to a silt or sand grade sediment, the first type with indefinite burrows with irregular outlines, the second type with well-defined burrows. The last three names describe hard bottoms characterized by encrusting and boring biota, and frequently a mineralized crust. Firmgrounds are stiff carbonate substrates in which the sediment has begun to consolidate through initial compaction, physical cohesiveness of fine particles, biological agglutinants (e.g. biofilms, organic mucus), or from early cementation. Firmgrounds may be initial stages of hardgrounds; they require some effort to become eroded, but can be ripped up by storms resulting in the formation of lithoclastic limestones. Hardgrounds are formed at times of extremely slow or intermittent sedimentation by lithification of the nondepositional surface prior to the deposition of the overlying bed. The hard surface requires substantial mechanical effort to break down and can be destroyed only by boring or by chemical processes (Sect. 9.3; Pl. 52).
12.1.3.1 Carbonate Substrate Types and Organism-Sediment Interactions Copper (1993) differentiated seven types of common soft and hard bottoms. • Muddy soft bottoms: Formed by sedimentation of microorganisms or breakdown of larger skeletons. Bottom dwellers are predominantly epibenthic deposit feeders because of high mud fractions associated with organic matter. Species diversity is often low. Fine lamination may be common. Muddy soft bottoms often occur at the base of sedimentary cycles, in protected quiet-water environments, in lagoons, and in deep-water environments. Known since the Cambrian. • Mangrove-sea grass muddy bottoms: (a) Mangroves: Fine-grained muds, formed in low-energy intertidal-supratidal tropical or temperate coastal environments by baffling of sediment, and deposition of epibionts and organic matter (Plaziat 1995). The mangrove ecosystem is also called ‘mangal’. Known since the Late Cretaceous. (b) Sea grass: Fine-grained muds, formed by baffling, and deposition of epibionts. Carbonate grains are stabilized by sea grasses. The sea grass ecosystem dates back to the Late Eocene.
Carbonate Substrate Types
• Shell bed bottoms: Autochthonous (biological) or allochthonous (physical) concentrations of the skeletons of shelly organisms. Hydraulic accumulations may be enhanced by storm events, tidal currents and turbidity currents. Common in inter-supratidal and shallow subtidal environments. Known since the Cambrian. • Carbonate sand to gravel bottoms: Mobile and unstable substrates, formed by bioclastic and non-bioclastic carbonate grains. Redistribution of the grains by currents and waves and winnowing of the fines are common. Burrowing deposit feeders as well as vagile epibionts may be abundant. Oolitic bottoms are known since the Precambrian, bioclastic sand bottoms since the Cambrian. • Coral-sponge-bryozoan-algae-dominated substrates: Muddy to sandy carbonate substrates in quiet waters, covered by basal skeletons of ‘coralline’ organisms, most of which are clones of the first individual resulting from initial larval settlement. Platy, hemispherical and cylindrical growth forms act in protecting and baffling sediment. The skeletons form hard substrates for epibionts and cryptic organisms. Known since the Ordovician. • Hard bottoms are substrates that are firm enough to become encrusted and bored and that are resistant to normal erosion (include firm grounds, hardgrounds, and conglomeratic to skeletal and rocky grounds). Sedimentation rates are low to very low. Known since the Cambrian. • Rocky bottoms: Hard substrates formed by rocks exposed at the sea floor, large boulders and bedrock surfaces; predominantly in shoreline environments. The specialized biota is characterized by abundant borers and encrusters. Known since the Cambrian but more records known since the Late Mesozoic. Interactions between organisms and carbonate substrate Benthic organisms require the organic matter within substrates as food resources and the sediment as larval settlement sites, shelter and for anchoring and fixation. Suspension-feeders (e.g. many bivalves, sponges, corals, bryozoans, brachiopods, crinoids) favor firm and hard or coarse-grained substrates whereas deposit feeders (e.g. worms, echinids, some bivalves, many gastropods) prefer silty and muddy bottoms, probably because of the higher organic content of muds. Organisms influence substrates through bioturbation and by providing skeletons, soft organic tissues and fecal pellets. Organisms may have a strong impact on (1) grain size and grain textures (by digestive activity and burrowing), (2) grain composition (by providing different types of bioclasts and by selective feeding), (3) sedimentary structures (by destroying bedding through bioturbation;
Recognition of Substrate Types
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Fig. 12.8. Inferred relationships among the major environmental factors of the Great Bahama Bank determining the relations between substrate and bottom-dwelling organisms. A: All ecologic factors depend on basin geometry and basin energy. Substrate types are predominantly controlled by water turbulence, circulation and salinity. After Newell et al. 1959. B: Relationships between feeding types, substrate and water energy in oolitic shoals. Increasing substrate mobility has two major ecologic consequences: (1) Increase in bottom current velocity results in an increase of substrate mobility, which in turn causes winnowing of organic matter from the sediment and prevents deposition of organic detritus. Consequently, the substrate contains only a relatively low percentage of organic matter. The low organic content limits the number of depositfeeding individuals and taxa. (2) Shifting bottoms inhibit colonization of the substrate by many marine plants. This limits the taxonomic and numerical abundance of the vegetal-matter consuming herbivores. The scarceness of organic nutrients in the substrate and the poorly developed flora on the substrate are reflected by a bottom fauna characterized by low-diversity suspension feeders represented by vagile forms (e.g. mollusks) on highly mobile substrates (such as oolite shoals), or sessile forms (e.g. sponges) on moderately mobile substrates (such as fine-grained sands). After Purdy (1964).
or installation of lamination through organic mats), and (4) water content and fluid movement within intergranular pores as well as sediment chemistry (thus influencing early diagenesis; see Meadows 1986). The activity of benthic organisms can change • grain size, sorting and shape, usually by digestive activity, • texture and fabric of the sediment, by mechanical displacement of the sediment, • sedimentary structures (bedding, lamination, topography), • water content and fluid movement in pores, by respiration and excretion, • compaction, • shear stress and stability (softness and hardness), and • sediment chemistry. Substrate and feeding types Basically four feeding types can be recognized: (1) suspension feeders feeding on microorganisms and organic detritus suspended in water, (2) deposit feeders feeding on the same material in or on the sediment, (3) carnivores feeding on other animals, and (4) herbivores, feeding on plant material. Omnivores employ two or more of the listed feeding types.
Bottom-dwelling organisms are deposit feeders and suspension feeders as well as carnivores and herbivores. The distribution of the various feeding types depends on the distribution of the requisite food source. Deposit feeders, which feed on the organic debris and detritus accumulated on or within the substrate, occur predominantly in fine-grained sediments. Suspension-feeders, which filter or strain suspended organic matter or microscopic organisms out of the water are common in medium- to coarse-grained silty and sandy sediments. Suspension feeders depend on the amount of food in the surrounding water; deposit feeders depend on the amount of food in and on the sediment. The distribution of nutrients at sea bottom and in the water column is controlled by differences in water energy. Fig. 12.8 summarizes the basic relationships between water energy, nutrient source and organisms with different feeding strategies.
12.1.3.2 Recognizing Substrate Types Thin-section criteria that allow primary substrate types to be differentiated are listed in Box 12.5. The main criteria of soft bottoms are ichnofabric and the composition of bottom and sediment dwellers. Hard bottoms can be recognized from borings and encrustations.
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Substrate Stability
Fig. 12.9. Recognition of substrate stability by benthic assemblages. The sample exhibits a skeletal grainstone with bryozoans (B) and crinoids (CR). This assemblage is typical for many Mississippian open-marine shelf limestones. The grain-supported texture indicates deposition of skeletal sand forming a firmground substrate. Most grains are not significantly abraded. The skeletal grains are of different sizes. Bryozoans represent two morphotypes (fenestrate forms - FB, and ramose cryptostomid forms - CB). Most cryptostomid bryozoans are well preserved, indicating only limited reworking. These bryozoans are regarded as autochthonous components of the assemblage. Fenestrate bryozoans grew as loose sheets over the substrate. The preservation of crinoids as isolated elements showing no signs of strong reworking and the dominance of sessile organisms unable to tolerate shifting substrates suggest stable substrate conditions. Similar patterns were described by Feldman et al. (1993). Early Carboniferous: Calgary, Canadian Rocky Mountains.
Plate 103 Soft Bottoms and Hard Substrates The sea bottom differs in consistency and stability of the sediment. Non-cemented carbonate muds and sands form soft bottoms, lithified carbonate sediments form hard bottoms. The state of consistency and stability prior to lithification can be inferred from paleontological and textural criteria. 1 Soft bottoms and hard substrates demonstrated by a poorly sorted skeletal floatstone. The fine-grained micritic matrix is a strongly burrowed bioclastic wackestone with shell debris. Larger fossils are radiolitid rudist bivalves (R), corals (C) and siliceous sponges (S). Rudists and corals are encrusted by coralline red algae (A). Criteria indicating substrate conditions are burrows and borings. Burrows (B) within the wackestone matrix are common but hard to recognize at a first view. Look at the typical, here only faintly visible, subrounded outline (white arrows), the slightly darker color and the close setting of the burrows. These criteria characterize a soft silty looseground substrate (Goldring 1995) with slight stabilization enabling the settlement of serpulid worms (SER). Corals and rudist shells form local hard substrates as indicated by borings and biogenic encrustation. The borings (BO) in the rudists have affected both the inner aragonitic and the outer calcitic shell layers. The borings marked by black arrows tell a different story. They transect the sedimentary fabric and are filled with gray crystal silt. These criteria show that the borers did not attack individual fossils but the whole rock and that this attack took place in a meteorically influenced post-lithification stage. The sample comes from a small reef mound consisting of floatstones and rudstones, and local coral and rudist boundstones. Late Cretaceous: Haidach near Brandenberg, Tyrol, Austria. Courtesy of D. Sanders (Innsbruck).
Plate 103: Soft Bottoms and Hard Substrates
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Box 12.5. Criteria for recognizing carbonate substrate types (for trace fossils see Goldring 1991, 1995; Bromley 1996). Biota • Distribution of deposit feeders and suspension feeders. • The ratio of epifauna and endofauna provides information on substrate stability. The mode of feeding is controlled by grain size: Fine-grained soft substrates are dominated by endobenthos; increase of substrate stability with increasing grain size results in an increase in epibiontic elements, deposit feeders decrease in coarse-grained sediments due to the early decomposition of organic matter. • Mode of life (e.g. cemented versus burrowing or boring; provides information on the consistency of the substrate). Trace fossils • Burrowing, bioturbation and ichnofabric pointing to soft bottoms (Sect. 5.1.4; Fig. 5.4; Pl. 103). Burrows are formed in muds, sands and firmgrounds (Pl. 19/ 1, 3, 7). • Stages in the preservation of trace fossils provide information of different stages of hardening. • Tiering. Animals producing trace fossils live at different levels (tiers) in the sediment and colonize newly deposited sediment at different rate in different ways. Tiering patterns offer information on substrate conditions, sedimentation rates and oxygenation levels (Bromley and Ekdale 1986; Wetzel 1991; Uchman 1995). • Abundant macro- and microborings on bed surfaces pointing to hardgrounds and firmgrounds (Sect. 9.3). Microfacies criteria • Depositional texture types (lime mudstones versus grainstones). • Biofabrics and current orientation of grains (Sect. 5.1.2) pointing to looseground conditions. • Biolamination point to firmground conditions (Sect. 5.1.3). • Mineralized crusts on bed surfaces point to hardgrounds (Sect. 5.2.4.1 and Sect. 9.2; Pl. 23/2, 3, 7). • Disturbances of sedimentary fabrics and structures. • Differences in grain size, sorting and packing.
Bioturbation and substrate types Because fine-grained sediment contains more organic matter than coarse-grained sediment, burrowing and bioturbation fabrics are common in micritic limestones (Sect. 5.1.4). The differentiation of soft-, looseand firmgrounds is primarily based on the preservation of burrows (Goldring 1995). An example is shown in Pl. 103.
Light
Substrate types of cold-water carbonates Carbonate sands formed in temperate and cold-waters of shelf platforms are characterized by the predominance of calcitic skeletal grains that are permanently reworked (Sect. 2.4.4.3). Cements are usually restricted to intraskeletal pores. Interskeletal cementation is weak due to the scarceness of primary aragonitic grains that could provide carbonate for cementation after dissolution. Sands remain unlithified for long periods, therefore many substrates are loosegrounds or firmgrounds. Basics: Carbonate substrates Bottjer, D.J., Ausich, W.I. (1986): Phanerozoic development of tiering in soft substrate suspension-feeding communities. – Paleobiology, 12, 400-420 Brett, C.E. (1988): Paleoecology and evolution of marine hard substrate communities: an overview. – Palaios, 3, 374-378 Copper, P. (1992): Organisms and carbonate substrates in marine environments. – Geoscience Canada, 19, 97-112 Ekdale, A.A. (1985): Paleoecology of the marine endobenthos. – Palaeogeography, Palaeoclimatology, Palaeoecology, 50, 63-81 Fabricius, F. (1968): Calcareous sea bottoms of the Rhaetian and Lower Jurassic sea from the west part of the Northern Calcareous Alps. – In: Müller, G., Friedman, G. (eds.): Recent developments in carbonate sedimentology in central Europe. – 240-249, Berlin (Springer) Feldman, H.R., Brown, M.A., Archer, A.W. (1993): Benthic assemblages as indicators of sediment stability: evidence from grainstones of the Harrodsburg and Salem limestones (Mississippian, Indiana). – In: Keith, B.D., Zuppann, C.W. (eds.): Mississippian oolites and modern analogs. – American Association of Petroleum Geologists, Studies in Geology, 35, 115-128 Goldring, R. (1995): Organisms and the substrate response and effect. – In: Bosence, D.W.J., Allison, P.A. (eds.): Marine paleoenvironmental analysis from fossils. – Geological Society, London, Special Publications, 83, 151180 Purdy, E.G. (1964): Sediments as substrates. – In: Imbrie, J., Newell, N.D. (eds.): Approaches to paleoecology. – 238271, New York (Wiley) Further reading: K172
12.1.4 Light Large-scale reductions in the incident light of benthic calcareous red algae have been suggested as causes for the mass extinction at the end of the Cretaceous (Aguirre et al. 2000), The functioning of the oceanic ecosystem is dependent for energy to a high degree on the photosynthetic activity of phytoplankton confined to a thin zone of lighted surface water. Light-dependent photosynthesis is also of fundamental importance for marine benthic plants (algae and sea grass), and for the life of many benthic invertebrates harboring photosynthetic endosymbionts.
Zonation of Light Intensity
Solar incident light at the water surface is closely dependent on latitude. Once the light has entered the water, its rate of penetration depends on the extinction coefficient of the water. Since this varies, subsurface illumination shows little correlation with a simple latitudinal pattern. To broadly generalized, light penetration decreases from the open ocean to the coasts, and from the tropics toward the poles. The amount of light depends on the angle at which the light strikes the water surface. A maximum angle producing a maximum penetration of light occurs in the tropics where light conditions are optimal all year long and photosynthesis works in all seasons. Approaching the poles, the amount of light penetrating the water surface and photosynthesis rates differ significantly with the season. About 50% of the light penetrating clear water is absorbed in the upper tens of meters. The depth to which light penetrates depends on the absorption of light by the water, wavelength, transparency, reflection from the surface of the water, reflection from suspended particles, season of the year and latitude. Large numbers of suspended particles in the water, e.g. in coastal waters, severly reduce light penetration; photosynthesis is possible only within a range of a few meters. In clear tropical waters, where interfering particles are rare, light intensity may be sufficient for photosynthesis down to 100–150 m. Sunlight is composed of radiation in a spectrum of wavelengths. As these wavelengths enter the seawater, violet and red components are quickly absorbed within the first few meters. Green and blue components are absorbed less rapidly and hence penetrate more deeply, although the intensity of light decreases. These differences control the distribution of photosynthetic organisms.
12.1.4.1 Zonation and Light Conditions Distribution patterns of light-dependent organisms are used in ecological and bathymetric zonations of the oceans (Liebau 1984; Lüning 1985). One common method distinguishes the illuminated photic and the non-illuminated aphotic zone. Other concepts differentiate three zones: • In the euphotic zone or photic zone there is sufficient penetration of light to support photosynthesis. Its lower boundary is the water depth at which the visible light is reduced to 1% of the light touching the water surface. This boundary corresponds approximately to the maximum compensation depth of algae and the lower limit of the chlorocline. The euphotic zone com-
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prises the supratidal, the intertidal and parts of the subtidal zones. The zone can be further subdivided into a shallow and a deeper euphotic zone using differences in light intensity and wavelength recorded by specific assemblages of endolithic cyanobacteria and algae (Budd and Perkins 1980; Glaub 1994) • The dysphotic zone is the twilight zone characterized by dim light and little photosynthesis. The boundary between the dysphotic zone and the aphotic zone is defined by the existence of obligate photoautotroph multicellular organisms. This boundary corresponds to the lower limit of the rhodocline, which receives 0.01 to 0.001% of the surface light and corresponds to the occurrence of crustose coralline algae. The zone is sometimes regarded as the lower part of the euphotic zone. Twilight is common in shadowy places, e.g. submarine caves and reefs. • In the aphotic zone there is not enough light for photosynthesis. Endolithic microborers are restricted to chemoheterotroph taxa. The zone comprises parts of the deep shelf and the deep-sea bottoms. These zones can be further differentiated into vertically arranged units whose boundaries are defined by the distributional limits for specific organisms (Liebau 1984): The limit of the uppermost unit, the chlorocline, is characterized by the last occurrence of green algae. The hermacline is defined by the lower limit of reef-building organisms with photosymbiotic algae. The rhodocline can be recognized by the lower limit of encrusting calcareous red algae. The lower limit of the chrysocline is defined by the last occurrence of autotroph single-celled algae, particularly diatoms. The deep-lying ophthalmocline zone is typified by blind arthropods, particularly ostracods. The depths of these boundaries depend on the optical quality of the water and the latitudinal position (see Sect. 12.3).
12.1.4.2 Recognition of Photic and Aphotic Conditions The following criteria are useful for estimating paleolight conditions. Microfacies criteria • The formation of some carbonate grain types is influenced by light conditions. Most cortoids (Sect. 4.2.3) and porostromate skeletal oncoids (cyanoids; Sect. 4.2.4.1) originate in the well-lighted euphotic zone.
Recognizing Oxygen
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• Aggregate grains (Sect. 4.2.7) are abundant in the upper euphotic zone. Algae and cyanobacteria • Occurrence of calcareous algae. Depth distribution patterns of green algae are much narrower than those of red algae, which can utilize minimal light intensities and specific wave lengths. Corallinacean red algae, therefore, may also occur in deeper waters down to more than 200 meters. • Dasyclad algae are restricted to a few meters in depth; the depth range of udoteacean green algae varies between a few tens of centimeters to about 80 meters. Since the fragments of calcareous green algae are often transported, some caution is needed when using the occurrence of dasyclads as an indicator of the upper euphotic zone. • Because light conditions control the growth rates of encrusting coralline algae, the thickness of the calcified algal body can be used in evaluating relative water depths (Sect. 10.2.1.2). • The taxonomic composition of red algal oncoids (rhodoids) reflects differences in light conditions. Rhodoids offer a possibility of distinguishing shallower and deeper parts of the euphotic zone. • Cyanobacteria are aerobic phototrophs. They occur in illuminated as well as in shadowy habitats and can better tolerate changes in light, salinity, water temperature and other ecological factors than algae (Sect. 9.1). Microborers • Microorganisms boring in shells or hard lithified substrates are excellent indicators of light conditions and water depths (Fig. 9.16). Boring cyanobacteria, algae and fungi are used to subdivide the euphotic and dysphotic zones. Organisms with photosymbionts • Benthic larger foraminifera with green and red algal symbionts are restricted to shallow euphotic areas (Sect. 10.2.2.1). Differences in light attenuation by the water column are reflected by different wall structures (Hohenegger 1999). Planktonic foraminifera with algal endosymbionts occur in water depths down to 50 and 100 m. • Various invertebrates (e.g. corals, some bivalves) are characterized by photosymbiotic algae. Calcification and distribution of zooxanthellate reef corals is greatly determined by endosymbiotic algae. The symbiosis of ancient corals and algae is inferred from the morphological criteria of the corals (e.g. integration and size of corallites) and from stable isotope data.
Morphological adaptions • Morphological features that may be interpreted as adaptions to light conditions occur in various invertebrates (e.g. eye reduction in ostracods and trilobites). Basics: Light Aguirre, J., Riding, R., Braga, J.C. (2000): Late Cretaceous incident light reduction: evidence from benthic algae. – Lethaia, 33, 205-213 Budd, D.A., Perkins, R.D. (1980): Bathymetric zonation and paleoecological significance in Puerto Rican shelf and slope sediments. – Journal of Sedimentary Petrology, 50, 881-553 Glaub, I. (1994): Mikrobohrspuren in ausgewählten Ablagerungsräumen des europäischen Jura und der Unterkreide (Klassifikation und Palökologie). – Courier Forschungsinstitut Senckenberg, 174, 1-324 Jerlov, N.G. (1970): Light. – In: Kinne, O. (ed.): Marine ecology. Vol. I (Environmental factors), part 1. – 95-102, London (Wiley) Liebau, A. (1980): Paläobathymetrie und Ökofaktoren: Flachmeerzonierungen. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 160, 173-216 Liebau, A. (1984): Grundlagen der Ökobathymetrie. – Paläontologische Kursbücher, 2, 149-184 Lüning, K. (1985): Meeresbotanik. Verbreitung, Ökophysiologie und Nutzung der marinen Makroalgen. – 375 pp., Stuttgart (Thieme) Riedl, R. (1966): Biologie der Meereshöhlen. – 636 pp., Hamburg (Parey) Further reading: K168
12.1.5 Oxygen Oxygen is a bio-limiting element for marine metazoa and one of the most important factors influencing species diversity and abundance. The absence of oxygen is generally considered to be essential for inhibiting microbial decay and forming organic-rich sediments (Allison et al. 1995). Oxygen-controlled environments occur in tidal flats, open to poorly enclosed shelf seas, upwelling areas, filled basins, and hydrothermal vent and cold seep areas (Oschmann 1995). All these environments show cyclic oxygen fluctuations. Estimates of dissolved-oxygen levels in ancient oceans contribute towards an understanding of the history of ocean circulation and paleoclimate. Paleo-oxygenation levels are based on fabrics (intensity of bioturbation), biota (foraminifera, bivalves) and geochemical criteria (organic carbon/sulfur ratio, sulfur isotopes, rare element concentrations). The oxygen content is generally high in the surface waters of the oceans. A shallow oxygen maximum (SOM) occurs in the upper ocean layers (Hayward
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Oxygen in the Water Column
1994), developing largely from photosynthetic oxygen production. The SOM forms on two different time scales, episodic (by physical events in which nutricline water is transported into the mixed layer and then ventilated to the atmosphere, producing a phytoplankton bloom) and gradual (developing from an input of ‘new’ nutrients transported from the depths or/and by sinking particles). Oxygen-minimum zone: Decomposition of organic matter produced in the surface waters causes a decrease in oxygen in many intermediate waters. The depth of that layer ranges in the Pacific Ocean and in the Indian Ocean from 100 to 1500 m, but oxygen values less than 0.5 ml/l water occur only between 500 and 1000 m. Benthic environments are only affected where the minimum layer intersects with shelf and slope environments. In ancient times, the oxygen minimum layer was of variable thickness and intensity. Reduced or expanded oxygen minimum zones are caused by variations in primary production and density stratification within the water body, which in turn control the advection and diffusion of oxygen. Expanded oxygen minimum zones have been reported from the Cretaceous.
12.1.5.1 Terminology and Classification Declining oxygen gradients cause a decrease in benthic faunal diversity until the oxygen-free anaerobic zone is reached (Rhoads and Morse 1971; Byers 1977; Savrda and Bottjer 1987). Ranking by diversity reveals paleo-oxygenation gradients (Wignall 1990). Levels of dissolved free oxygen are classified by the content of free oxygen in the water expressed as ml/l, and indicated by the terms • oxic: >1ml/l O2 • dysoxic: 1.0 to 0.2 ml/l O2 • suboxic: 0 to 0.2 ml/l O2 • anoxic: <0.2 ml/l O2, commonly no free O2. The term euxinic refers to anoxic conditions but with free H2S in the water column. Anoxic sediments and anoxic bottom waters are produced where there is a deficiency in oxygen due to very high organic productivity and a lack of oxygen replenishment to the water or the sediment, e.g. in the case of water stratification or stagnation (Southam et al. 1982). Extended periods of anoxia exclude virtually all higher invertebrates from the benthic environment. Most anoxic sediments are siliciclastic sediments represented by black shales. This term denotes a fine-
grained, dark-colored and compacted sediment, usually with distinct lamination and commonly more than 0.5% organic matter. The dark gray color is caused by finely dispersed iron sulfides and organic compounds. Argillaceous black shales were common in the Proterozoic, the Early Paleozoic, and during the Jurassic and Cretaceous (Kukal 1990). Controls on the mode of life and the biofacies by oxygen are described by the terms • aerobic: normal benthic fauna, presence of free oxygen, no oxygen restriction, • dysaerobic: impoverished benthic fauna stressed by low bottom-water oxygen values • anaerobic: no benthic fauna, due to the lack of free oxygen. These zones fit depth-related environments in which oxygen values remain rather constant. The term poikiloaerobic (Oschmann 1991) refers to annual and seasonal fluctuations of low-oxygen values in shallow shelf and epeiric environments that are reflected by low diversity and an opportunistic benthic fauna. The term exaerobic (Savrda and Bottjer 1987) refers to a biofacies near the anaerobic-dysaerobic boundary characterized by specialized benthic shelly organisms harboring chemosymbiotic bacteria. Oxygen-restricted biofacies (ORB) and inferred paleo-oxygenation levels can be derived from the number of species and the sediment fabric (Wignall and Hallam 1991). Examples were summarized by Allison et al. (1995). ORB 1 and ORB 2 are equivalent to the anaerobic biofacies and characterized by the total absence of benthic organisms and a fissile sediment fabric. ORB 3 and ORB 4 contain only a few benthic species which are very rare in ORB 3 or abundant on some bedding planes (ORB 4). Sediment fabric is planar laminated but bioturbation may occur. ORB 4 corresponds to the poikiloaerobic and episodically disaerobic facies. ORB 5 and ORB 6 are characterized by bioturbation. ORB 3 to ORB 6 correspond to the dysaerobic biofacies.
12.1.5.2 Recognizing Paleo-Oxygenation Generally, low organic content, light rock colors, existence of diverse and highly-diverse benthic faunas and trace fossils, and evidence of hydraulic activity (e.g. cross-stratification and high-energy textures (Sect. 12.1.1) are taken as proxies of oxic environments. High organic content, dark rock colors, bituminous smell,
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Box 12.6. Criteria of paleo-oxygenation levels (based on Allison et al. 1995; Goldring 1991; Oschmann 1988, 1991; Rhoads and Morse 1971; Sageman et al. 1991; Savrda and Bottjer 1986, 1989). Biota • Presence or absence of benthic organisms. Complete absence or lack of benthic organisms with shells. • Absence of deposit feeders may indicate oxygen-deficient conditions. • Trace fossil assemblages. Abundance, type and size of trace fossils. • Diversity of benthic organisms and trace fossils. Number of species. • A benthic fauna of very low diversity, or lack of any benthic elements point to oxygen-deficient conditions. • Specific morphological and physiological adaptions of benthic organisms (e.g. mollusks, brachiopods) to oxygen-deficient and H2S conditions. • Morphology, test size, wall thickness and pore patterns are used to differentiate a benthic foraminiferaloxygen index (Kaiho 1994). The index describes dissolved-oxygen levels in the modern ocean. • Exceptional preservation of soft parts of fossils potentially indicating anoxic conditions. Fabric • Lamination. • Fine-grained, microlaminated sediments with medium to high amount of organic matter indicate oxygen-deficient environments. • Alternation of millimeter-laminated sediments to variously burrowed sediments may point to dysoxic conditions. • Abundance and intensity of bioturbation. Size and orientation of burrows. • A shift from bioturbation to discrete burrows indicates a gradual decrease in the oxygen level. Geochemistry • Preservation and composition (organic compound ratios, e.g. pristane/phytane) of organic matter (Jones 1987). • Total organic carbon (Pedersen and Calvert 1990) and hydrogen index (Tyson and Pearson 1991). • Organic carbon/sulfur ratios (Gautier 1986; Kajiwara and Kaiho 1992). • Uranium enrichment (Wignall and Myers 1988). • Rare element anomalies (e.g. cerium; Wright et al. 1987) and concentrations (Anderson et al. 1989). • Pyritization of iron minerals (Raiswell et al. 1988).
fine lamination, high lateral persistence, absence or rareness of benthic organisms and trace fossils are used to suggest anoxic or dysoxic conditions. These features occur in argillaceous sediments as well as in limestones (e.g. Tyszka 1994). Box 12.6 lists criteria used to interpret ancient oxygen-poor or oxygen-free sea bottom environments.
Oxygen in the Sediment
12.1.5.3 Case Study: Black Shale Development on a Carbonate Platform A specific Late Triassic facies type (Seefeld facies) of the Northern Calcareous Alps is characterized by organic-rich laminated marlstones and limestones intercalated in a sequence consisting of dolomitic wackestones and mudstones of the Hauptdolomit facies (Hopf et al. 2001). This facies represents a tidal, partly restricted depositional environment. Differences in paleo-oxygenation conditions are indicated by biotic composition and diversity, preservation of fossils, microfacies, bioturbation and microlamination (Fig. 12.10). Aerobic conditions are indicated by a gastropod-ostracod-foraminifera community and a monotypic bivalve assemblage. Small burrowing bivalves associated with ostracods reflect dysaerobic conditions. Anoxic conditions are characterized by the absence of benthic biota, the presence of microlamination, and the excellent preservation of a fish fauna within the bitumen-rich beds. Organic geochemical data (pristane/ phytane ratio) provide evidence of cyclically occurring anoxic conditions in morphological depressions within the platform, and a stagnant, stratified water body. Basics: Oxygen Allison, P.A., Wignall, P.B., Brett, C.E. (1995): Paleo-oxygenation: effects and recognition. – In: Bosence, D.W.J., Allison, P.A. (eds.): Marine paleoenvironmental analysis from fossils. – Geological Society London, Special Publications, 83, 97-112 Berner, R.A. (1984): Sedimentary pyrite formation: an update. – Geochim. Cosmochim. Acta, 48, 605-615, Oxford Ekdale, A.A., Mason, T.R. (1988): Characteristic trace-fossils associations in oxygen-poor sedimentary environments. – Geology, 16, 720-723 Etter, W. (1995): Benthic diversity patterns in oxygenation gradients: an example from the Middle Jurassic of Switzerland. – Lethaia, 28, 259-270 Hopf, H., Thiel, V., Reitner, J. (2001): An example of black shale development on a carbonate platform (Late Triassic, Seefeld, Austria). – Facies, 45, 203-210 Jones, R.W. (1987): Organic facies. – Advanced Petrology and Geochemistry, 2, 1-90 Kukal, Z. (1990): Recent and ancient marine anoxic environments. – Casopis pro mineralogii a geologii, 35, 287-300 Oschmann, W. (1991): Anaerobic-poikiloaerobic-aerobic: a new zonation for modern and ancient neritic redox facies. – In: Einsele, G., Riecken, W., Seilacher, W. (eds.): Events and cycles in stratigraphy. – 565-571, Berlin (Springer) Oschmann, W. (1993): Environmental oxygen fluctuations and the adaptive response of marine benthic organisms. – Journal of the Geological Society London, 150, 187-191 Oschmann, W. (1994): Adaptive pathways of benthic organisms in marine oxygen-controlled environments. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 191, 393-444
Paleo-Oxygenation: Example
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Fig. 12.10. Estimating paleo-oxygenation. The fabric of the black laminated lime mudstone is characterized by couplets of densely spaced dark and bright laminae (A), alternating with light layers composed of silt-sized grains (B). The lack of bioturbation and fossils in the microlaminated sets reflects anaerobic conditions (Oxygen-Restricted Biofacies – ORB 2) that alternated with dysaerobic layers containing a few ostracod shells (arrow; ORB 4). A younger, thin black, more or less horizontal fissure enriched in bitumen is intercalated. Late Triassic (Norian, Seefeld facies): Wiestal quarry, Hallein, Austria. Raiswell, E., Buckley, F., Berner, R.A., Anderson, T.F. (1988): Degree of pyritization of iron as a paleoenvironmental indicator of bottom-water oxygenation. – Journal of Sedimentary Petrology, 58, 812-819 Rhoads, D.C., Morse, J.W. (1971): Evolutionary and ecologic significance of oxygen-deficient marine basins. – Lethaia, 4, 413-428 Savrda, C.W., Bottjer, D.J. (1991): Oxygen-related biofacies in marine strata: an overview and update. – In: Tyson, R.V., Pearson, T. (eds., 1991): Modern and ancient continental shelf anoxia. – Geological Society of London, Special Publication, 58, 201-219 Southam, J.R., Peterson, W.H., Brass, G.W. (1982): Dynamics of anoxia. – Palaeogeography, Palaeoclimatology, Palaeoecology, 40, 183-198 Tyson, R.V., Pearson, T. (eds., 1991): Modern and ancient continental shelf anoxia. – Geological Society of London, Special Publication, 58, 470 pp. Wignall, P.B. (1994): Black shales. – Geology and Geophysics Monograph Series, 30, 136 pp., Oxford (Oxford University Press) Wignall, P.B., Hallam, A. (1991): Biofacies, stratigraphic distribution and depositional models of British onshore Jurassic black shales. – In: Tyson, R.V., Pearson, T.H. (eds.): Modern and ancient continental shelf anoxia. – Geological Society London, Special Publication, 58, 291-309 Further reading: K166
12.1.6 Seawater Temperature Sea-water temperatures and the prevailing paleoclimate exerted a major control on the depositional patterns, diagenetic pathways and reservoir styles of carbonates (Sun and Esteban 1994). Regional and global changes of water temperatures are regarded as essential in diversity reduction and extinction of organisms. The reconstruction of paleo-water temperatures is largely based on biotic and geochemical proxies. Temperature is one of the most important factors in governing life processes and distribution of organisms. Vital life processes in the oceans function normally only within defined ranges of temperature (approximately 2 °C to 40 °C). Seawater temperature varies laterally with changes in latitudes and vertically with depths. The depth of the most rapid temperature decline below the warm surface water is called the thermocline. This zone is a persistent feature in the tropics, but is absent in the polar zone. The temperatures of tropical surface waters range between about 20 and 30 °C throughout the year. In temperate surface waters warm temperatures are restricted to summer. Water temperatures below the thermocline are rather constant.
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12.1.6.1 Seawater Temperature: Biotic Proxies The interpretation of paleotemperature conditions is predominantly based on paleontological features that are only sometimes recorded in thin sections. The only thin-section criteria exhibiting significant temperature controls are skeletal grain associations (see Sect. 12.2) and ooids (Sect. 4.2.5). The occurrence of abundant ooids is often used as a proxy for warm water environments of tropical settings (Opdyke and Wilkinson 1990). The main assumption is that many carbonate ooids form in warm waters of high salinities (Lees 1975). Common paleontological criteria used in the evaluation of seawater temperature are • Distribution patterns of benthic associations. Warmwater and cool-water associations exhibit marked taxonomic differences. • Transfer of the temperature range tolerated by extant taxa to ancient taxa with close taxonomic relationships. This method can be successfully applied to Quaternary and sometimes also to Tertiary taxa, but becomes increasingly more hazardous for Mesozoic and Paleozoic taxa. • Transfer of temperature ranges and abundance patterns of defined recent organism groups to fossil groups. Many modern marine invertebrates are limited to warm waters (e.g. larger foraminifera, hermatypic corals); others occur predominantly in temperate and cold waters (e.g. brachiopods). This approach has many pitfalls. • Diversity of benthic organisms expressed by species number and abundance. The general rule, that diversity is high in warm water and low in cold water, has many exceptions. • Evaluation of organisms exhibiting strong temperature controls. Commonly used groups are algae, foraminifera, corals, mollusks and ostracods. • Benthic foraminifera. Many agglutinated foraminifera (Textulariina) prefer cold-temperate and brackish waters. Miliolina are abundant in warm waters of normal and high salinity. Rotaliina occur in waters of intermediate temperatures and normal salinity. The ‘temperature index’, using the relative proportions of warm water and cold water genera of benthic foraminifera, is valuable in evaluating paleotemperatures of shelf environments back to the Mid-Cretaceous. • Planktonic foraminifera. The sea surface temperature (SST) can be estimated using the transfer function of planktonic foraminifera and other calcareous plankton (see Murray 1995). The success of the method depends on the quality of preservation of foraminiferal
Seawater Temperature
assemblages. Warm-water planktonic foraminifera tend to be more susceptible to solution than cold water foraminifera; this may lead to an enrichment in cold water forms (Miao et al. 1994). The method is successfully applied to Cenozoic paleoenvironments. • Reef-building corals: Warm sea-surface temperature is a key factor controlling the distribution of modern scleractinian corals and coral reefs in the tropics and subtropics owing to physiological limitations to metabolic activity (calcification) and growth (larval survival) and ecological limitations (e.g. competitive interactions between corals and soft algae). These relationships involve only zooxanthellate corals. Scleractinian corals without endosymbiotic algae also construct reefs in cold waters. The distribution patterns of Mesozoic and also Early Tertiary coral reefs exhibit many exceptions as compared with modern temperature-controlled patterns (see papers in Kiessling et al. 2002). The question of climate controls on reefs built by organisms other than corals is strongly debated (Kiessling 2001). • Skeletal mineralogy. The mineral composition of invertebrate skeletons exhibits temperature-dependent relations (Dodd and Stanton 1990). The relation between the aragonite/calcite ratio and temperature is not always obvious, however, because of the effects of other environmental factors, e.g. salinity, and because the mineralogy-temperature effect can differ from species to species. • Morphological criteria of skeletal elements, e.g. the thickness and size of shells. • Rhythmic growth patterns of bivalve shells or other invertebrates, e.g. banding of corals, generally associated with seasonality. Thickness, density and spacing of increments as well as trace element data and stable isotopes reflect seasonal fluctuations of water temperature. Examples are known from bivalves, corals and corallinacean red algae. However, temperature fluctuations are just one factor in the formation of growth increments. • Functional criteria of skeletal elements, e.g. the number and length of spines, size of pores, and coiling directions of planktonic foraminifera.
12.1.6.2 Geochemical Proxies of Seawater Temperatures Carbonate oxygen isotope palaeothermometry The most widely used geochemical technique to determine paleotemperatures is based on the ratio of oxygen isotopes in calcite and aragonite. The method is applied to modern organisms and fossils as well as to carbonate sediments and limestones. The fraction-
Seawater Temperature
ation of oxygen isotopes is temperature dependent. The ratio of the two most common isotopes (16O and 18O) in calcareous fossils can, in principle, be used to reconstruct the temperatures of ancient oceans. Foraminifera are commonly analyzed in order to determine seasurface, deep-water and bottom-water temperatures (Berger 1979). Mollusk and brachiopod shells as well as carbonate cements have also been successfully used. Difficulties of the technique include problems related to the isotopic composition of ancient oceans, the possibility of non-equilibrium fractionation in organically precipitated calcites (specially in macrofossils), and the diagenetic alteration of the δ18O values of carbonate shells (Corfield 1995). Further information in Arthur et al. (1983), Wefer and Berger (1991) and Hoefs (1997). Trace elements The concentration of trace elements in fossils depends on the physical chemistry of the skeletal formation process, the physiology of the organisms, various environmental parameters, and diagenetic processes. Many studies of minor element concentration and elemental ratios in fossils and carbonate sediments deal with magnesium and strontium. These elements are used to estimate water temperature ranges as well as other environmental constraints. Magnesium concentration correlates positively with water temperature in many invertebrates. The correlation is especially apparent in groups characterized by skeletons made by High-Mg calcite (e.g. corallinacean red algae) but also occurs in many other invertebrates. The Sr concentration of inorganic calcite and aragonite is positively correlated to increasing temperature. The strontium concentration of biogenic aragonite and calcite shows positive as well as negative correlations with water temperature. The Sr/Ca ratio in coral skeletons reflects sea-surface temperatures (Beck et al. 1992). Magnesium and strontium concentrations can not be used universally as paleothermometers, because the correlation between temperature and rare element concentration may fit a major group (e.g. all brachiopods) or only genera or species of a group (e.g. of mollusks). See Moore (1989) and Dickson (1990) for further information. Basics: Seawater Temperature Beck, J.W., Edwards, R.L., Ito, E., Taylor, F.W., Recy, J., Rougiere, F., Joannot, P., Henin, S. (1992): Sea-surface temperature from coral skeletal strontium/calcium ratios. – Science, 257, 644-647 Francis, J.E., Smith, M.P. (eds., 2002): Paleoclimate reconstructions using fossils and lithological indicators. – Palaeogeography, Palaeoclimatology, Palaeoecology, 182, 1-143
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Kiessling, W. (2001): Paleoclimatic significance of Phanerozoic reefs. – Geology, 29,751-754 Murray, J.W. (1995): Microfossil indicators of ocean water masses, circulation and climate. – In: Bosence, D.W., Allison, P.A. (eds.): Marine paleoenvironmental analysis from fossils. – Geological Society of London, Special Publications, 83, 245-264 Ribaud-Laurenti, A., Hamelin, B., Montaggioni, L., Cardinal, D. (2001): Diagenesis and its impact on Sr/Ca ratio in Holocene Acropora corals. – International Journal of Earth Sciences, 90, 438-451 Rosenberg, G.D. (1990): The ‘vital effect’ on skeletal trace element content as exemplified by magnesium. – In: Carter, J.G. (ed.): Skeletal biomineralization patterns, processes and evolutionary trends. Volume I. – 567-577, New York (Van Nostrand) Strauch, F. (1972): Zur Klimabindung mariner Organismen und ihre geologisch-paläontologische Bedeutuung. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 140, 82-127 Sun, S.Q., Esteban, M. (1994): Paleoclimatic controls on sedimentation, diagenesis and reservoir quality: lessons from Miocene carbonates. – American Association of Petroleum Geologists, Bulletin, 76, 519-543 Wefer, G., Berger, W.H. (1991): Isotope paleontology: growth and composition of extant calcareous species. – Marine Geology, 100, 207-248 Further reading: K155, K156 (trace elements), K158, K159 (stable isotopes), K171 (paleotemperature evaluation).
12.1.7 Salinity Understanding salinity controls on biota and microfacies is important specifically for carbonates that have been formed in marginal-marine and nonmarine environments. Salinity is a measure of dissolved salts in seawater. It is defined as the weight in grams of dissolved salts in 1000 g of seawater and commonly expressed in terms of per mille (‰) which is equivalent to parts per thousand (ppm) of dissolved solids in the water. Common cations are Na, Mg, Ca, K, and Sr. Terminology: Salinity conditions of ancient environments are often described in rather general terms those as normal marine water, brackish water or freshwater. Normal marine conditions are characterized by a salinity in the range of 33 to 38‰, average 35‰. Brackish water is an indefinite term for waters with a salinity between that of normal seawater and that of normal freshwater (see Fig. 12.11). Freshwater contains only small quantities of dissolved salts. Salinity decreases in areas of mixing with fresh water and varies strongly in estuaries and intertidal settings. Salinity is low in semiclosed sea basins (e.g. Baltic Sea). Salinities that are substantially greater than that of normal seawater (>40‰) are called hypersaline. They occur in arid
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Paleosalinity freshwater 0.5‰
5‰
limnic brackish oligohaline
brackish 10‰ brackish mesohaline
18‰
marine 30‰
marine brackish restricted marine brachyhaline
normal marine euhaline
40‰ restricted marine hypersaline
Cyanobacteria .................................++++++++++++++++++++++++++++++++++++++++++++++++++++++++++++++ Dasyclad green algae ...................................................................................................................................+++++++ Udoteacean green algae ...............................................................................................................................+++++++ Charophycean algae .............................++++++++++++++++++++++++++++++++ Corallinacean red algae ...................................................................................................................+++++++++++ Agglutinated benthic foraminifera .......................................................+++++++++++++++++++++++++++++++++++++ Calcareous benthic foraminifera ...............................................................................++++++++++++++++++++++ Radiolaria .................................................................................................................................................+++++++ Coralline demosponges ................................................................................................................................+++++ Hexactinellid sponges ...............................................................................................................................+++++++ Corals .......................................................................................................................................................+++++++ Bryozoans .......................................................................................................................+++++++++++++++++++ Brachiopods .........................................................................................................................++++++++++++++++++ Serpulids ...........................................................................................+++++++++++++++++++++++++++++++++++ Gastropods .....................................+++++++++++++++++++++++++++++++++++++++++++++++++++++++++++ Bivalves ..........................................+++++++++++++++++++++++++++++++++++++++++++++++++++++++++++ Cephalopods ......................................................................................................................................++++++++++ Ostracods ........................................+++++++++++++++++++++++++++++++++++++++++++++++++++++++++++ Balanid crustaceans ..................................................................................+++++++++++++++++++++++++++++ Echinoderms ......................................................................................................................................++++++++++
Fig. 12.11. Approximate distribution of some recent organisms relative to salinity. The figure does not include rare exceptions (e.g. freshwater sponges). The subdivisions follows the Venice System (Oertli 1963).
evaporitic settings (e.g. parts of the Red Sea; Persian Gulf), in salinas and restricted lagoons as well as in semiclosed basins and bays. The term schizohaline describes strongly fluctuating salinity conditions which may be caused by episodic flushing of hypersaline environments by freshwater (Folk and Siedlecka 1974; Pl. 50/6). Estimations of paleosalinity are based predominantly on biotic, microfacies and geochemical data.
12.1.7.1 Biotic and Microfacies Proxies of Paleosalinity Paleontological data are heavily used in reconstructing paleosalinities and discriminating normal marine, brackish and freshwater environments, although salinity is rarely the only parameter influencing faunal composition and diversity. But if salinity is the dominant factor, salinity gradients can be determined using benthic associations, particularly of mollusks (Fürsich and Werner 1986). Some difficulties arise in differentiating hypersaline and brackish conditions using paleontological data alone. The capability to regulate the osmotic pressure may allow organisms to thrive both in hypersaline and brackish waters and biota as well as microfacies types may exhibit considerable similarities (e.g. Fürsich et al. 1995).
Organisms are stenohaline or euryhaline. Stenohaline organisms do not or are only able to tolerate a narrow range of salinity. Euryhaline organisms tolerate wide fluctuations in salinity. The differentiation of euryhaline and stenohaline organisms on a higher taxonomic level is a widely used approach in estimating paleosalinities. Stenohaline are planktonic foraminifera, radiolarians, corals, most bryozoans, brachiopods, cephalopods, and with a few exceptions echinoderms. Common euryhaline are cyanobacteria, sponges, many bivalves, gastropods, and ostracods. Important paleontological criteria that can be used in estimating paleosalinities are • Salinity ranges. Comparison of the salinity dependence of modern and ancient organisms. Fig. 12.11 shows the salinity ranges of some groups. The distribution of major taxonomic units within these groups may be more specific than the range of the whole group as demonstrated by the foraminifera. Textulariina are abundant in brackish waters, Miliolina occur in waters of normal to hypersaline salinity and Rotaliina are common in waters of normal salinity. • Faunal composition. The lack or scarcity of stenohaline invertebrates is characteristic of environments with fluctuating salinity conditions. • Extended microbial mats and stromatolites (e.g. Persian Gulf; Shark Bay, Australia).
Paleosalinity: Geochemical Proxies
• Ostracods are most valuable indicators of paleosalinities (Pl. 48/2; Pl. 93/1; Pl. 130/2). They occur with facies-diagnostic species and assemblages in freshwater, brackish water and marine environments. • Charophycean algae (Pl. 65/1, 3; Pl. 130/4) are used as indicators of freshwater and brackish water environments (Sect. 10.2.1.8). Alternations of limestones with charophyta, ostracod limestones and intercalations of evaporites may reflect seasonally varying salinities. • Faunal diversity (species richness and evenness) can reflect high-stress conditions. A diversity minimum exists around a salinity of 5‰. Diversity decreases away from the species minimum toward both the freshwater end and the marine end of the salinity range (Remane 1958). Similarly, the number of individuals often decreases from normal marine environments to both brackish or hypersaline restricted environments. See Croghan (1983) for critical remarks. • Shell size and thickness: The thickness and size of freshwater and brackish shells is often smaller than that of marine shells. Decreasing and increasing salinities are often related to a decrease in the maximum size. • Taphonomic features. Burrowers and grazers normally destroy microbial mats in marine waters of normal salinity. These organisms are absent in high salinity areas. Microbial mats formed in hypersaline environments therefore have a greater chance for preservation. Microfacies data reflecting paleosalinity deviating from that of the normal marine seawater include specific grain types: • Radial-fibrous ooids are common, but not restricted to freshwater, brackish and hypersaline environments (Sect. 4.2.5; Fig. 4.20; Fig. 4.26). • Distorted ooids and half-moon ooids (Fig. 4.23). • Pisoids/vadoids (Sect. 4.2.6; Box 4.19; Pl. 14). • Microbial peloids are abundant in hypersaline environments. • Freshwater- and brackish oncoids (Sect. 4.2.4.1) differ from marine oncoids by specific nuclei and lamination, and by better preservation (Box 4.10; Pl. 12/7; Pl. 131/5). The case study displayed in Pl. 104 demonstrates the controls of alkalinity and salinity on carbonate formation and lithification.
12.1.7.2 Geochemical Proxies of Paleosalinity Common criteria used in recognizing differences in the salinity of waters and diagenetic fluids are stable isotopes and trace elements.
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Stable isotopes Differences in the stable carbon and oxygen composition from either fossil or rock samples between marine and nonmarine carbonates (Keith and Weber 1964) can be used to reconstruct paleosalinities (Mook 1971; Dodd and Stanton 1990; Yin 1991; Joachimski 1991). The results, usually based on δ13C data, should however be checked by paleontological criteria (Yin et al. 1995, Mátyás et al. 1996). Trace elements Trace elements in skeletons, both of fossils and in carbonate rocks provide a possibility for estimating paleosalinities: The favorite fossils used in paleosalinity interpretations are ostracods because these organisms occur in continental, estuarine, marine and hypersaline waters, because their Low-Mg calcite valves are relatively resistant to diagenetic changes influencing the geochemical data, and because the relationships between the uptake of elements and chemical conditions of the aquatic environment are well studied (Bodergat 1983). Contrary to mollusk shells, which represent a whole life growth pattern, ostracod tests – corresponding to molts – may record water chemistry during just a very short interval of a few days only. In tropical seawater, the concentrations of Mg, Sr and Na increase with increasing salinity. These elements and their ratios are commonly studied in the context of paleosalinity (e.g. Dodd 1967; Chivas et al. 1986). Sodium in carbonate rocks has been used as a possible index to the salinity of diagenetic solutions (Land and Hoops 1973) as well as a paleosalinity indicator (Veizer et al. 1977; Ogorolec and Rothe 1979; Sass and Bein 1988). Changes in salinity during the deposition of subtidal to supratidal limestone/dolomite cycles are reflected by decreasing Sr and increasing Na, Fe and Mn contents (Buggisch et al. 1994). Since rare element concentrations and the Mg/Ca, Sr/Mg and the Sr/Ca ratios of fossil shells and carbonate rocks depend on many other factors than salinity, their implication for paleosalinity interpretations is much debated (Dodd and Stanton 1990; Rosenberg 1990; Carpenter and Lohmann 1992; Holmes et al. 1992; Rao 1996). Boron: Another tool studied in determining paleosalinities is the boron content of argillaceous sediments (Harder 1970; Perry 1972). Spivack et al. (1987) have shown that the bulk of boron in clay minerals is not in isotopic equilibrium with seawater and can, therefore
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Satonda Model
Plate 104 Salinity and Alkalinity Controls: The Satonda Model Satonda is a small volcanic Indonesian island north of Sumbawa. The island is unique in that the crater lake is filled with seawater with increased alkalinity. Because lake and ocean waters cannot communicate, the lake water acquired higher alkalinity, pH and carbonate mineral saturation than in the open ocean. These conditions exclude most marine macrobiota but favor reef-like structures composed of red algae, serpulids and foraminifera, and biogenic mats built by in situ calcifying cyanobacteria and green algae. Morphology and microstructure of these deposits are similar to certain types of microbialites that are common in the Late Precambrian and Early Phanerozoic oceans. The conditions exemplified by the Satonda case study have been used to establish a new model focused on the role of alkalinity for calcification processes in early Phanerozoic oceans (Kempe and Kazmierczak 1990). The formation of carbonates was controlled by changes in salinity and alkalinity (Kempe and Kazmierczak 1993). Salinity changes are reflected by the composition of the ostracod fauna (Schudack and Reitner 1996). The crater lake was originally filled with freshwater. Shortly after the lake was flooded with seawater. During this phase marine organisms colonized the lake starting with serpulids. Large input of organic matter from the crater walls caused intense oxygen consumption. As a result, the bottom water of the lake became anaerobic and sulfate reduction led to the precipitation of carbonates. The lake became stratified into water bodies of different salinity and high alkalinity. A significantly increased carbonate mineral supersaturation caused a in situ calcification of cyanobacterial microbialites forming microstromatolites (–> 1) and to the lithification of red algal crusts (–> 7). 1 Macroscopic view of microstromatolites exhibiting subcolumnar structures. The stromatolites were formed by cyanobacteria in association with non-calcifying green algae Cladophora in a water depth of 13 m and waters with a salinity of about 30 ‰. 2 In-situ calcification of cyanobacteria. The SEM view of the surface of a coccoid cyanobacterial mat shows the capsular character of gelatinous sheats surrounding Pleurocapsales cells embedded within a microgranular Mg-calcite matrix precipitated in vivo by the mat itself. The lower part of the photograph corresponds to a dark lamina of the microstromatolite, the upper part to a fibrous aragonitic lamina (compare –> 4). 3 The internal fabrics of the microstromatolites are laminated (this figure) or cystose (–> 6). The light aragonite laminae alternating with dark Mg-calcite cyanobacterial laminae overgrow filaments of non-calcifying siphonocladalean green algae (empty molds). 4 The vertical section of the microstromatolitic structure shows sparry (light) laminae consisting of fibrous aragonite and microcrystalline (dark) laminae composed of High-Mg calcite. The differences in the mineralogical composition are interpreted by seasonally changing supersaturation in the lake. During the wet season the supersaturation is lowered due to rain allowing the growth of the mats. The carbonate supersaturation rises during the dry season resulting in an in vivo precipitation of High-Mg calcite at the surface mat. Decay of the mat below the permineralized surface triggered by sulfate-reducing bacteria connected with increase in alkalinity and of the Mg/Ca ratio causes the post mortem precipitation of aragonite instead of Mg-calcite. Similar alternations of thin microgranular and thicker fibrous laminae are common features of ancient fabrics interpreted as microbialites. 5 The upper part of the calcareous crust consists of intergrown corallinacean red algae and variously calcified pleurocapsalean cyanobacteria. The lower part consists of squamariacean red algae. 6 The cystose microbialite structure reminds on a common Palaeozoic fossil (Wetheredella, see Pl. 51/9) whose systematic position is not settled (Kazmierczak and Kempe 1992). 7 Parts of the reef framework consist of encrusting squamariacean red algae (Peyssonelia) and accumulations of dark peloids of fecal and cyanobacterial origin. Present-day Squamariaceae grow as flattened crusts or petaloid forms on soft substrates (e.g. circumlittoral soft-bottoms of the Mediterranean). The internal structure is characterized by closely packed basal cells from which vertical rows of cells arise. Note the attached fans of aragonite (arrows). Fossil squamariaceans show in thin sections a yellow-brown tint in contrast to corallinacean red algae (Pl. 54/1, 3) exhibiting a dark grey tint. 8 Green algal meadows with Cladophora occur in the upper parts of the lake down to 16 m. Red algal crusts are also found in deeper positions. The photograph shows longitudinal and cross sections of algal filaments (arrows) and irregularly shaped open voids. Compare the picture with the Tertiary algal structures shown in Pl. 130/1. –> 1–8: Courtesy of S. Kempe (Darmstadt)
Plate 104: Salinity and Alkalinity Controls: The Satonda Model
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not be used in salinity analysis. In contrast, the abundance and isotopic composition of boron incorporated in biogenic carbonate skeletons may be useful in tracing changing paleoenvironmental conditions (Vengosh et al. 1991).
12.1.7.3 Microfacies Proxies for Hypersaline and Evaporitic Conditions High salinities are common in arid shallow-marine environments, e.g. evaporite-carbonate shorelines or lagoons, and can be inferred for ancient sediments by lithologic and mineralogical criteria (e.g. evaporites). Box 12.7 summarizes criteria signaling hypersaline and saline conditions that can also be recognized in microfacies studies. Basics: Salinity Buggisch, W., Noé, S., Krumm, S. (1994): Geochemische und fazielle Muster in peritidalen Kalk/Dolomit-Zyklen der oberen Bellerophon-Schichten (Oberperm) in den Südalpen. – Abhandlungen der Geologischen Bundesanstalt Wien, 50, 69-87 Chivas, A.R., De Deckker, P., Shelley, J.M.G. (1986): Magnesium content of non-marine ostracod shells: a new palaeosalinometer and palaeothermometer. – Palaeogeography, Palaeoclimatology, Palaeoecology, 54, 43-61 Folk, R.L., Siedlecka, A. (1974): The ‘schizohaline’ environment: its sedimentary and diagenetic fabrics as exemplified by Late Paleozoic rocks of Bear Island, Svalbard. – Sedimentary Geology, 11, 1-15 Friedman,G.M., Shukla, V. (1980): Significance of authigenic quartz euhedra after sulfates: example from the Lockport Formation (Middle Silurian of New York). – Journal of Sedimentary Petrology, 50, 1299-1304 Fürsich, F.T. (1994): Palaeoecology and evolution of Mesozoic salinity-controlled benthic macroinvertebrate associations. – Lethaia, 26, 327-346 Grimm, W.D. (1964): Ausfällung von Kieselsäure in salinar beeinflußten Sedimenten. – Zeitschrift der deutschen geologischen Gesellschaft, 1962, 589-619 Jacobs, L., Swennen, R., Van Orsmael, J., Notebaert, L., Vieaene, W. (1982): Occurrences of pseudomorphs after evaporitic minerals in the Dinantian carbonate rocks of the eastern parts of Belgium. – Bulletin de la Société belge de Géologie, 91, 105-123 Kempe, R., Kazmierczak, J. (1993): Satonda Crater Lake, Indonesia: hydrogeochemistry and bicarbonates. – Facies, 28, 1-32 Kempe, S., Kazmierczak, J. (1994): The role of alkalinity in the evolution of ocean chemistry, organization of living systems and biocalcification processes. – Bulletin de l’Institut océanographique, Monaco, no. spéc., 13, 61-117 Mátyás, J., Burns, S.J., Müller, P., Magyar, I. (1996): What can stable isotopes say about salinity? An example from the Late Miocene Pannonian lake. – Palaios, 11,31-39 Mook, W.G. (1971): Paleotemperatures and chlorinities from stable carbon and oxygen isotopes in shell carbonate. –
Hypersaline and Evaporitic Conditions
Box 12.7. Microfacies criteria pointing to hypersaline and saline depositional and/or diagenetic conditions. • Interbedded lime mudstones and layers with anhydrite or gypsum crystals (Pl. 109/9). Common in supratidal sabkha environments. • Microfolds in fine-grained dolomites resulting from dehydration of primary gypsum or hydration of anhydrite (Pl. 125/4). • Collapse breccias may result from the dissolution of evaporites interbedded with limestones. The breccia beds are characterized by stratiform geometries with welldefined lower boundaries. • Silicified or calcitized nodules formed by replacement of anhydrite nodules (Hennebert and Hance 1980). Common in carbonates formed in intertidal environments or on supratidal sabkha flats. • Half-moon ooids (Fig. 4/23) may owe their shape to a collapse of insoluble parts following the dissolution of envelopes composed of anhydrite or gypsum. The grains can indicate vanished evaporites. • Pseudomorphs of evaporite minerals in limestones (Paszkowski and Szydlak 1986). • Palisade calcites. Millimeter- to centimeter-sized, conical calcite rosettes (Pl. 125/2) or clusters representing pseudomorphs after gypsum formed in restricted environments (Swennen et al. 1981). • Length-slow chalcedony after sulfate evaporite minerals (Pittman and Folk 1971). Length-slow means that the slower polarized light component produced in double refraction is the one vibrating parallel to the crystal length. Formed when silica replaces a sulfate mineral or when it replaces carbonate minerals in an evaporite milieu. May originate early after deposition of the sediment or after lithification. • Authigenic idiomorphic and double-terminated euhedral quartz crystals (Pl. 109/5, 6) are common in micritic limestones formed in hypersaline environments. The carbonates are often associated with evaporites and salt deposits (Grimm 1962, 1964). These quartz crystals are formed both during early and late diagenetic phases (Richter 1971). • Authigenic euhedral feldspar. The association of albite-bearing carbonates and evaporites and saline fluid inclusions in albite are considered as tracers of paleobrine-carbonate reactions. These reactions may take place shortly after the deposition of carbonate mud in hypersaline environments, but also post-depositionally in deep burial and incipient metamorphic settings (Misik 1994; Spötl et al. 1999). • Low-diversity but abundant skeletal grains. Limestones with accumulations of ostracods (Pl. 130/2, 3), or of monospecific bivalves and gastropods may have been formed in environments with significant departures from normal marine salinities.
Palaeogeography, Palaeoclimatology, Palaeoecology, 9, 245-263 Oertli, H.J. (1964): The Venice system for the classification of marine waters according to salinity. – Pubblicazioni della Stazione Zoologica di Napoli, 33, Supplement, p. 611
Productivity and Nutrients
Pittman, J.S., Folk, R.L. (1971): Length-slow chalcedony after sulphate evaporite minerals in sedimentary rocks. – Natural Physical Science, 230, 64-65 Rye, D.M., Sommer, M.A. (1980): Reconstructing paleotemperature and paleosalinity regimes with oxygen isotopes. – In: Rhoads, D.C., Lutz, A.R. (eds.): Skeletal growth of aquatic organisms. Biological records of environmental change. – 169-202, New York (Plenum Press) Spötl, C., Longstaffe, F.J., Ramseyer, K., Rüdinger, B. (1999): Authigenic albite in carbonate rocks - a tracer of deepburial brine migration? – Sedimentology, 46, 649-666 Yin, J. (1991): Stable carbon and oxygen isotopes in Jurassic shells as palaeosalinity indicators. – Neues Jahrbuch für Geologie und Paläontologie, Monatshefte, 1991, 163-176 Yin, J., Fürsich, F., Werner, W. (1995): Reconstruction of palaeosalinity using carbon isotopes and benthic associations: a comparison. – Geologische Rundschau, 84, 223236 Further reading: K155, K156 (trace elements), K158, K159 (stable isotopes), K170 (paleosalinity criteria).
12.1.8 Productivity and Nutrients Nutrients are important controls on the rate of carbonate production, bioerosion and the formation and destruction of carbonate platforms and reefs. Productivity collapse is regarded as a prime factor in major extinction events (e.g. Wood 1993; Caplan et al. 1996). Productivity is the rate of production of organic material by biological processes. Primary productivity refers to the production of energy-rich organic compounds from inorganic materials predominantly by photosynthesis, but also by chemosynthetic bacteria, and is expressed in terms of grams of carbon fixed per unit area or volume of seawater per time interval. Prime factors affecting primary productivity are light, nutrients and vertical mixing of water, which is different in temperate, tropical and polar seas. Secondary productivity designates the production of organic material by animals utilizing preexisting plant and animal organic matter. Major texts concerning marine productivity are listed under ‘Basics’. Brasier (1995a, 1995b) gave a concise summary of current thinking on the controls and effects of nutrient levels on organisms and environments. 12.1.8.1 Nutrients Bio-limiting nutrient elements are phosphorus, nitrogen, iron and silicon. In the presence of light and optimal temperatures these elements control the primary production by photosynthesis. Nutrient supply results from terrestrial influx or organic matter, marine transgression or upwelling. Optimal conditions for primary
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production are largely confined to coastal regions of the continental shelf due to the input of nutrients together with weathered materials from rocks on land and upwelling of cold nutrient-rich waters from the deep. Sediments associated with upwelling are rich in marine organic carbon, biologically produced silicon and phosphorus. In shallow depths plants photosynthesize faster than they respire and more organic material is produced than consumed. At the compensation depths photosynthesis and respiration are equal. Below these depths respiration exceeds photosynthesis and respiration by animals as well as inorganic oxidation continues. Destruction of organic matter frees nutrient elements that go back into solution in seawater. This results in a low concentrations of nutrients at the surface and increasing concentration below the compensation depth. This basic pattern is changed by vertical circulation. Some primary production occurs below the compensation depth in deep waters by organisms using chemical energy. Spectacular examples are bacteria that obtain their energy by oxidizing sulfide in hot vents. These bacteria form the basis of an entire ecosystem. The geographic variation in nutrient element concentration is influenced by global circulation patterns. Areas of divergent currents usually exhibit high nutrient element concentrations and high productivity. The waters can also be enriched in nutrient elements by upwelling water bodies. Areas of convergence in the center of circulation gyres usually have low nutrient concentrations. Nutrients may also be provided by geothermal endoupwelling (Rougerie and Wauthy 1986, 1993; Rougerie et al. 1992): In this thermo-convective process in Pacific atolls and barrier reefs remnant geothermal heat from the volcanic foundation drives a continual upward flow of deep oceanic waters (rich in nutrients) through the overlying porous reef framework. The nutrients at the reef surface support the high metabolism of the living community. The importance of this mechanism has been also stressed for ancient reefs (Rougerie and Fagerstrom 1994). Surface waters may be eutrophic or oligotrophic. Eutrophic conditions are characterized by high and oscillating levels of nutrient supply, and high net productivity. Good examples are the upwelling areas off the west coast of Africa. Oligotrophic conditions are characterized by low and stable nutrient supply and lower net productivity. Examples are the coral reefs of the Caribbean, and the Indian and Pacific Ocean. Mesotrophic conditions are characterized by increasing nutrient supply favoring suspension feeders. Factors which
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can raise or lower nutrient levels and primary production, and encourage or reduce eutrophic conditions were summarized by Föllmi et al. (1994) and Brasier (1995).
12.1.8.2 Estimating Paleonutrient Levels Several factors triggering higher or lower nutrient levels are recorded by facies, paleontological and geochemical criteria (Box 12.8).
12.1.8.3 Effects of Nutrient Excess on Reef and Platform Carbonates The primary carbonate sediment producers of coral reef communities and of shallow carbonate platforms are
Nutrient Level Criteria
highly adapted to nutrient-deficient environments. Changes in nutrient concentration are regarded as a major control on the demise of coral reefs and drowning of platforms (Schlager 1981; Hallock and Schlager 1986; Föllmi et al. 1994; Dupraz and Strasser 2002) and on the biotic composition of reefs (Birkeland 1988). The input of phosphates and nitrates stimulates the growth of plankton, which reduces water transparency, limiting depth ranges of zooxanthellate corals and calcareous green algae and thereby reducing carbonate production. Higher nutrient concentrations and plankton densities stimulate the growth of fast-growing soft algae and ahermatypic suspension-feeding benthic organisms, and increase bioerosion by microborers. Many ancient drowned platforms and reefs exhibit evidence of nondeposition, bioerosion and reduced redox potential, pointing to excess nutrient availability
Box 12.8. Criteria of carbonates reflecting possible differences in nutrient levels. Note that some of the facies criteria can also be caused by other environmental constraints. Based on Brasier (1995a, 1995b), Wood (1999) and other sources. Facies criteria • Basic types of autochthonous limestones. The dominant reef biota and the resulting carbonate types are related to nutrient levels. Bindstones and framestones originate in low-nutrient environments; mud mounds formed by finegrained micrite with weakly calcified heterotrophs, and bafflestones originate in areas of higher and high nutrient levels (Fig. 8.1; Wood 1993). • Microbial mats. Abundant mats may point to eutrophic conditions due to increased nitrate fixation by cyanobacteria and other microbes (see Sect. 9.2.). • Bioerosion. Levels of bioerosion are high in eutrophic areas, and relatively low in oligotrophic environments.An increase in nutrient supply (C, N, P and trace metals such as Fe and Mo) in oligotrophic environments influences the population of bioeroders and may inhibit reef growth (Peterhänsel and Pratt 2001). The relations between bioerosion and micro-encrusters reflect changes from oligotrophic to mesotrophic and eutrophic conditions. • Bioturbation. Intensive burrowing of fine-grained carbonates may indicate eutrophic conditions caused by the transport of nutrient elements and organic matter from deeper sediment layers to the sediment-water interface. • Oncoids and cortoids. Common and autochthonous cyanobacterial oncoids and cortoids formed by microborers indicate higher nutrient levels. • Terrigenous grains. Increased fluvial runoff (e.g. during regression) will deliver nutrients and lead to higher nutrient levels than a reduced input of terrigenous material into shallow-marine carbonate environments (Erlich et al. 1990). Biota • Trophic composition of paleocommunities. A dominance of deposit-feeders indicates a preferred accumulation of nutrients within the sediment. A dominance of suspension feeders points to food particles largely suspended in the water column.
• Benthic nutrient types. Suspension feeders are common in eutrophic environments, symbiotic forms and grazers are more common in oligotrophic environments. • Biotic diversity is often low in eutrophic and higher in oligotrophic environments. • Density and abundance of sessile benthic fossils may reflect differences in nutrient levels. An example are Mesozoic siliceous sponges exhibiting high densities in low nutrient areas and low densities in high nutrient environments. • Changes in ecologic successions and guilds in reefs. A vertical succession from frame-building to baffling and from high-growing to low-growing sessile organisms may indicate an increase in nutrients. • Planktonic organisms. Accumulations of radiolarians, diatoms, and organic-walled phytoplankton as well as the occurrence of cold-water planktonic foraminifera may indicate upwelling and high bio-productivity. Mineralogy • Phosphatic fossils. Accumulations of phosphatic grains. Formation of phosphorites (see Sect. 13.1.2.4). • Glauconite formation. Phosphogenesis associated with minimal carbonate accumulation may indicate eutrophic conditions (Föllmi et al. 1994). Geochemistry • Organic matter and organic carbon in carbonate rocks. Intercalation of bedded carbonates and laminated organicrich sediments and fluctuations in organic carbon may reflect fluctuations in bio-productivity and nutrients or they reflect preservation differences (Parrish 1995). • Geochemical criteria. Ba/Ca ratios as well as positive δ18O and more negative δ13C values in planktonic shells of upwelling areas are valuable indicators of nutrient conditions (see Sect. 13.2). • Phosphate grains and phosphorites. Both criteria are used in the context of the discussion on bio-productivity (Föllmi 1996; see Sect. 13.1.2.4).
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during drowning. Nutrient excess has various causes. Reef carbonates that are overlain by shales may be victims of nutrients in terrestrial runoff. Rapid transgression that flooded subaerially exposed platforms and ramps may have supported the influx of soils and nutrients to areas of reef growth. Other mechanisms for drowning by excess nutrients include rapid pulses of sea-level fluctuations and local or regional upwelling. Studies of modern coral reefs indicate that an increase in nutrient input and the related increase in productivity lead to changes in dominant species, in an increase in fast-growing and smaller taxa, and in greater competition for space. Corals may be less abundant than algae, sponges and bryozoans. This change can be recorded in ancient reefs by distinct differences in successions and growth types. Note that the relations between reef builders and nutrients summarized by the Hallock and Schlager model can be used to explain changes in ancient reefs formed by zooxanthellate corals, but can not be applied to reefs formed by siliceous sponges, bryozoans or microbes.
mise of coral reefs and carbonate platforms. – Palaios, 1, 389-398 Hottinger, L. (1987): Conditions for generating carbonate platforms. – Memorie della Società Geologica Italiana, 40, 265-271 Peterhänsel, A., Pratt, B.R. (2001): Nutrient-triggered bioerosion on a giant carbonate platform making the postextinction Famennian benthic community. – Geology, 29, 1079-1082 Rougerie, F., Fagerstrom, J.A. (1994): Cretaceous history of Pacific basin guyot reefs: a reappraisal based on geothermal endo-upwelling. – Palaeogeography, Palaeoclimatology, Palaeoecology, 112, 239-260 Summerhayes, C.P., Prell, W.L., Emeis. K. (eds., 1992): Upwelling systems: evolution since the early Miocene. – Geological Society, London, Special Publication, 64 Thiede, J., Suess, E. (eds., 1983): Coastal upwelling. Its sediment record. Part B. – NATO Conference Series, IV, 10b, New York (Plenum) Whelan, W.K., Farrington, J.W. (eds., 1992): Organic matter: productivity, accumulation and preservation in recent and ancient sediments. – 550 pp., New York (Columbia University Press) Wood, R. (1993): Nutrients, predation and the history of reefbuilding. – Palaios, 8, 526-543 Further reading: K212
Basics: Productivity and nutrients Berger, W.H., Smetacek, V.S., Wefer, G. (eds., 1989): Productivity of the ocean: present and past. – 470 pp., Chichester (Wiley) Birkeland, C. (1988): Second-order ecological effects of nutrient input into coral communities. – Galaxea, 7, 91-100 Bombardière, L., Gorin, G.E. (2000): Stratigraphical and lateral distribution of sedimentary organic matter in Upper Jurassic carbonates of SE France. – Sedimentary Geology, 132, 177-203 Brasier, M.D. (1995a): Fossil indicators of nutrient levels. 1. Eutrophication and climate change. – In: Bosence, D.W.J., Allison, P.A. (eds.): Marine Palaeoenvironmental analysis from fossils. – Geological Society, London, Special Publications, 83, 113-132 Brasier, M.D. (1995b): Fossil indicators of nutrient levels. 2. Evolution and extinction in relation to oligotrophy. – In: Bosence, D.W.J., Allison, P.A. (eds.): Marine palaeoenvironmental analysis from fossils. – Geological Society, London, Special Publications, 83, 133-150 Caplan, M.L., Bustin, R.M., Grimm, K.A. (1996): Demise of a Devonian-Carboniferous carbonate ramp by eutrophication. – Geology, 34, 715-718 Dupraz, C., Strasser, A. (2002): Nutritional modes in coralmicrobialite reefs (Jurassic, Oxfordian, Switzerland): evolution of trophic structure as response to an environmental change. – Palaios, 17, 449-471 Föllmi, K.B., Weissert, H., Bisping, M., Funk, H. (1994): Phosphogenesis, carbon-isotope stratigraphy, and carbonate-platform evolution along the Lower Cretaceous northern Tethyan margin. – Geological Society of America, Bulletin, 106, 729-746 Hallock, P. (1988): The role of nutrient availability in bioerosion: consequences to carbonate buildups. – Palaeogeography, Palaeoclimatology, Palaeoecology, 63, 275-291 Hallock, P., Schlager, W. (1986): Nutrient excess and the de-
12.2 Estimating Paleoclimatic Conditions: Grain Association Analysis Shelf carbonates from warm- and cool-water regimes can be recognized by identifying the dominant skeletal constituents. Grain association analysis is one of the most promising tools of microfacies studies, but the results should be verified on the basis of accumulated evidence (James 1997). Cold climatic settings, for example, are indicated by specific grain associations, but are more clearly proven by ice-rafted dropstones or calcite pseudomorphs after ikaite, a metastable hexahydrate of calcium carbonate (see Sect. 2.4.1.5). Grain association patterns have a high potential in understanding climate change and paleolatitude shifts of carbonate ramps and platforms. The following text deals with the grain composition patterns of modern and ancient shelf carbonates. Examples of facies models deduced from these patterns are described in Chap. 14. 12.2.1 Concepts Modern shelf carbonates are formed in tropical and nontropical zones. The criteria of these sediments were already discussed in Sect. 2.4.4. Tropical carbonate sedimentation at low latitudes takes place in warm-water environments. Non-tropical sedimentation occurs at mid- and high latitudes in warm-temperate and cool-
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temperate, and in polar cold-water environments (Fig. 2.9). The warm-temperate zone running roughly parallel to the equator at latitudes of approximately 30° S and 30° N forms the link between the tropical zone and the cool-temperate zone. Basic differences between tropical and non-tropical shelf carbonates relate to biota and diagenetic patterns as well as to the composition of grain associations. Grain association analysis Grain associations (sometimes also called grain assemblages) are characterized by the dominating skeletal and non-skeletal grains. Ooids, aggregate grains and peloids are abundant constituents of warm-water carbonates, but are absent in cool-water carbonates except for the peloids. The abundance and association of skeletal grains are controlled by the dependence of benthic organisms on environmental factors discussed in the preceding sections. The composition of skeletal grain associations is influenced by many factors including light, water temperature, salinity, and the rate of carbonate production versus the rate of siliciclastic input. Lees and Buller (1972) and Lees (1975) tied the composition of modern low- and high-latitude shelf carbonates to temperature and salinity. The authors introduced the term chlorozoan skeletal association for tropical carbonates composed of calcareous green algae (Chlorophyta) and hermatypic corals (Zoantharia). The name chloralgal refers to tropical carbonates dominated by chlorophycean algae. The term foramol skeletal association was introduced for non-tropical carbonates consisting predominantly of foraminifera and mollusks. The term has proven to be less useful because of the wide compositional spectrum of non-tropical cool-water sediments. As a result, other expressions have come in use, partly based on the investigation of recent sediments (e.g. bryomol; Nelson 1988), partly on studies of Tertiary (Hayton et al. 1995), Cretaceous (Carannante et al. 1988) and Late Paleozoic (Beauchamp 1994; Beauchamp and Desrochers 1997) shelf limestones. The original concept favoring water temperature and, to a lesser content, salinity as the main controls on the global distribution of chlorozoan (chloralgal) and foramol associations draws the boundary between both associations at the 15 °C to 18 °C or more generally at the 20 °C annual isotherm, which parallels the tropical zone in most parts of the world, except in areas affected by Coriolis-driven cold-water currents. In these areas, upwelling of nutrient-rich cold water favors the dominance of autotrophs (e.g. calcareous red algae) over heterotrophs. A significant impoverishment in the foramol association at the 10 °C and 5 °C isotherms leads to differentiation of temperate carbonates into
Photozoan and Heterozoan Associations
warm-temperate and cold-temperate carbonates (Brookfield 1988). The predominantly siliceous polar hyalosponge sediments lie beyond the 5 °C isotherm, often in areas with semipermanent or permanent ice cover (Henrich et al. 1992). Photozoan and Heterozoan Associations James (1997) accentuating the light dependence/independence of benthic carbonate-producing organisms introduced two major association categories: • The Photozoan Association is characterized by an association of benthic carbonate grains including (1) skeletons of light-dependent organisms (algae; invertebrates, e.g. scleractinian corals and foraminifera, containing photosymbionts), and/or (2) non-skeletal grains (ooids, peloids and others), and (3) skeletons from the Heterozoan Association. The term replaces the term chlorozoan association and covers the chloralgal as well as many association types based on ancient limestones (see Box 12.9). It reflects the existence of shallow, warm-water, benthic calcareous communities and their resulting sediments, today mostly confined to tropical and subtropical settings. • The Heterozoan Association consists of benthic carbonate particles produced by (1) organisms that are light-independent, and plus or minus (2) calcareous red algae. The Heterozoan Association comprises many grain associations established for Cenozoic shelf carbonates (Box 12.9). It occurs from the poles to the equator and from the intertidal zone to the deep ocean. Cool-water carbonates are always heterozoan. But the Heterozoan Association does not necessarily indicate cool-water sediments! Heterozoan associations predominate where the supply of other sediments (photozoan, pelagic or siliciclastic) is limited by environmental constraints. In warm water the Heterozoan Association characterizes environments below the photic zone. But heterozoan grain associations can also be produced in the photic zone as demonstrated by the Mediterranean Sea, where heterozoan-type organisms living on sea grass form carbonates in warm-temperate settings (Betzler et al. 1997). The photozoan and heterozoan concept is useful in tracing major and prolonged paleoclimate trends and/ or paleolatitudinal shifts. A rigorous use of the concept, however, can mask differences better indicated by the individual grain associations listed in Box 12.9. Grain association types Box 12.9 summarizes the criteria of grain association types of modern and ancient shelf carbonates. All
Grain Association Types
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Box 12.9. Grain association types of warm-water and cool-water shelf carbonates. The list records the name, author and age of the material used, followed by the abundant and (separated by a hyphen) minor constituents, and the observed or inferred setting. H (heterozoan) and P (photozoan) at the end of the paragraph indicate the position within the major Grain Associations distinguished by James (1997). Percentage composition corresponds to average values for grains not for bulk samples. • Barnamol (Hayton et al. 1995; Tertiary). Barnacles • Chlorosponge (Beauchamp 1994; Permian). Green al(~60%), bivalve mollusks (~20%). – Bryozoans, echino- gae, coralline calcareous sponges. – Benthic foraminifera, derms, benthic foraminifera. Non-tropical coastal and open- oncoids. Warm-water. Low latitudes. P. • Chlorozoan (Lees and Buller 1972; modern). Calcaremarine shelf environments. H. • Bimol (Hayton et al. 1995; Tertiary). Bivalve mollusks ous green algae, zooxanthellate corals. – Benthic foramin(infaunal and/or epifaunal; > 60%, up to 80%). – Bryozo- ifera, articulated coralline red algae, mollusks, non-skelans, benthic foraminifera, echinoderms, calcareous red al- etal grains. Tropical. Warm-water. P. • Echinofor (Hayton et al. 1995; Tertiary). Echinoderms gae, barnacles. Non-tropical, but also tropical. H. • Bryomol (Nelson 1988; modern and Tertiary). Bryo- (about 40%; mainly echinoids), benthic foraminifera zoans (> 50%), bivalve mollusks (infaunal and epifaunal). (~35%; foraminifera may outspace the echinoderms). – – Benthic foraminifera, echinoderms, calcareous red al- Bryozoans, planktonic foraminifera. Non-tropical. H. gae, barnacles. Abundant in non-tropical, mostly cool- and • Foramol (Lees and Buller 1972; modern). Benthic foracold-water environments, but also occurring in subtropi- minifera, mollusks. – Bryozoans. The term is now somecal and some tropical environments. High-energy open what outdated and replaced by new terms, including the shelves and ramps. Analysis of bryozoan growth form as- rhodalgal and molechfor associations. – Non-tropical. Temsists in distinguishing shallow and deeper settings (see over- perate and cool-water. H. • Hyalosponge (Beauchamp 1994; Permian). Siliceous view by Smith 1995). H. See Pl. 106/2, 3. • Bryonoderm (Beauchamp 1994; Beauchamp and sponge spicules. The name is taken from modern examples Desrochers 1997; Permian). Bryozoans, echinoderms (usu- (Henrich et al. 1992). Deep-water sediments (carbonates ally crinoids). – Brachiopods; rare peloids. Regarded as and cherts) at low latitudes, or shallow to deeper-water the equivalent of modern bryomol associations. Low-lati- settings at high latitudes. H. tude deeper-water shelves to high-latitude, deep- to shal- • Molechfor (Carannante et al. 1988; Tertiary). Mollusks, low-water shelves. Shallow and deep cool-water carbon- benthic foraminifera. Non-tropical, warm-temperate enviates are biotically alike, but differ in depositional texture ronments. H. (grainstone versus packstone). Cool-water and cold-water. • Nannofor (Hayton et al. 1995; Tertiary). Planktonic foraminifera (~80%), nannofossils (coccolithophorids). – H. See Pl. 106/1. • Bryonoderm extended (Beauchamp 1994; Permian). Echinoderm spines, sponge spicules, benthic foraminifera, Bryozoans, echinoderms. – The main composition corre- bryozoans. Not a typical shelf assemblage. Usually in offsponds to that of the bryonoderm association, but addi- shelf deeper waters, but also in shallow, partly enclosed tional biota elements (solitary and colonial corals, benthic basins. smaller foraminifera, larger foraminifera, brachiopods) may • Rhodalgal (Carannante et al. 1988; Tertiary). Calcareous crustose red algae (> 80%; commonly rhodoids), bryobe common. Non-tropical and tropical environments. H. • Chloralgal (Lees 1975; modern). Calcareous green al- zoans, benthic foraminifera, barnacles, bivalves, echinogae. – Benthic foraminifera, branching coralline algae, mol- derms. – Common in non-tropical warm-temperate settings. lusks. Tropical and warm-temperate zones. Controlled by High-energetic photic zone. Transitional between chlorozoan-dominated tropical areas and cold-temperate ones. salinity and water temperature. P. See Pl. 105/2. • Chloroforam (Beauchamp 1994; Permian).– Calcare- Also in tropical zones where environmental conditions preous green algae (dasyclads, udoteaceans), benthic foramin- vent the development of a chlorozoan facies. H. ifera (encrusting and free forms, smaller and larger fora- • Rhodechfor (Hayton et al. 1995; Tertiary). Calcareous minifera), non-skeletal grains (ooids, aggregate grains) and red algae (up to 80%), echinoderms (mostly echinoids), benthic foraminifera, bryozoans. – Barnacles, bivalves. – oncoids. Warm-water. Low latitudes. P. See Pl. 105/1. Non-tropical temperate settings. H.
associations consider the most abundant skeletal constituents in wackestones, packstones, and grainstones. Some definitions are satisfied with the qualitative aspect, but most of them use quantitative data derived from point-counter measurements. Hayton et al. (1995) subjected the skeletal abundance data to cluster analyses that indicate the dominant groups and their variations. Grain association analysis is an essential element in categorizing Standard Microfacies Types (Sect. 14.3) and interpreting Standard Facies Models (Sect. 14.3.1).
12.2.2 Practical Advice, Examples and State of Current Information 12.2.2.1 Distinguishing Grain Association Types Attributing unknown samples to the associations summarized in Box 12.9 requires the following steps: • Check criteria that point to autochthonous grain assemblages or to mixing and resedimentation of environmentally different grains.
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• Identify skeletal grains in thin sections. • Determine their abundance, i.e. major and minor constituents, using semi-quantitative or point-counter techniques. • Compare the results with the grain association types listed in Box 12.9. Note that these types are rather loosely defined and that sometimes ‘minor’ skeletal grains will appear at a first rank position. • Assign your grain associations to statistically based assemblages, using cluster analysis or more sophisticated methods (see Sect. 6.3). • For Cenozoic carbonates consider larger foraminifera for differentiating warm-temperate and tropical zones (Betzler et al. 1997; Halfar et al. 2000). Of particular value in determining the skeletal grain assemblage of Cenozoic shelf carbonates is the composite triangular diagram developed by Hayton et al. (1995).
12.2.2.2 Examples Plates 105, 106 and 107 display Paleozoic, Cretaceous and Tertiary grain association types used to evaluate paleoclimatic conditions. The examples consider tropical and non-tropical shelf carbonates. Non-tropical reef
Warm-Water Grain Associations
carbonates will be discussed in Sect. 16.2.5.2. The samples were selected from case studies whose paleolatitudinal position (derived from the paleomaps in Kiessling et al. 2002) is in fair accordance with the attribution of the sample to tropical, temperate or coldwater zones.
12.2.2.3 State of the Art The composition of Cenozoic warm-water carbonates is well known, although more studies are needed based on thorough quantitative data. Nearly all the studies of ancient temperate and cool-water carbonates deal with Ordovician, Late Paleozoic, Late Cretaceous and Tertiary examples (Box 12.10). Ordovician and Late Paleozoic case studies reflect the influence of glaciation on climate-dependent cold-water carbonates formed in periods of deglaciation. Regional climatic variations and the occurrence of tropical cold-water grain associations in warm-water environments are discussed in papers on Cretaceous shelf carbonates (e.g. Gischler et al. 1994). Tertiary examples from southern Australia and New Zealand are of particular interest, because of the continuous record of temperate sedimentation in these regions.
Plate 105 Skeletal Grain Association Analysis: Tropical Warm-Water Shelf Carbonates Modern low-latitude tropical and mid- to high-latitude non-tropical shelf carbonates differ in the overall composition of abundant and dominating skeletal grains. Common constituents in tropical warm-water settings are calcareous green algae, benthic foraminifera, mollusks and hermatypic corals. This gross composition can also be recognized in Mesozoic carbonates. Corals, larger foraminifera and some bivalves harbor endosymbionts that require light and are therefore limited to water depths of only a few meters. Non-skeletal particles are ooids, peloids and aggregate grains. The plate displays common warm-water grain assemblages of the Photozoan Association type. 1 Chloroforam grain association. This association is typified by the abundance of calcareous green algae and benthic foraminifera, often accompanied by non-skeletal grains. Characteristic chloroforam associations occur as early as the Paleozoic and are common in the warm waters of tropical platforms and shallow ramps. The sample is from the inner part of a Cretaceous extended ramp whose paleoposition was approximately 10° S. Algae are dasyclads (D; Cylindroporella and others) and udoteaceans (U; Arabicodium). Foraminifera are represented by miliolids (M) and biserial textulariids (T). Other grains include cortoids and peloids (left corner). Note early marine and shallow burial interskeletal cements, a common feature of warm-water carbonates. Middle Cretaceous: Subsurface, Arabia. 2 Chloralgal grain association characterized by the abundance of calcareous green algae. This association has been known since the Cambrian. Modern examples occur in protected areas of tropical and subtropical platforms. The plate-like fragments are Epimastopora Pia, an alga commonly assigned to the dasyclads. Some dasyclads are partly encrusted by foraminifera. Additional biota in modern chloralgal associations are benthic foraminifera and mollusks. Late Paleozoic associations commonly exhibit brachiopods (BR) instead of mollusks. The letter O indicates rare ostracods. The sample comes from the inner part of a locally restricted carbonate platform with a paleolatitude of ~10° S. SMF 18-DASY. Early Permian: Carnic Alps, Austria.
Plate 105: Warm-Water Grain Associations
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Positive aspects of grain association analysis • Compositional analysis assists in recognizing longterm climate changes and short-term climate fluctuations. Long-term changes were studied by Beauchamp (1994) in the Canadian Sverdrup Basin. Benthic grain associations, their biota and mineralogical criteria indicate a significant cooling for the Permian of the Canadian Arctic. Other examples are known from the Late Cretaceous of Sardinia (Carannante et al. 1995). Shortterm variations of climate and sea level (cool lowstands and warm highstands) are documented by successions
Aspects of Grain Associations
of repeated intervals of bryonoderm and chloroforam carbonates in Late Pennsylvanian cyclothems (Samankassou 2002), and bryomol and chlorozoan carbonates in the Upper Miocene ramps in southern Spain (Brachert et al. 1997). • The evaluation of grain type associations is important in understanding the drowning of carbonate platforms. Foramol platforms are characterized by low sedimentation rates. Rising sea level will cause drowning of these platforms, whereas tropical chlorozoan platforms may persist (Simone and Carannante 1988).
Plate 106 Skeletal Grain Association Analysis: Cold-Water and Temperate Shelf Carbonates Bryozoans are often the most abundant carbonate grains in non-tropical carbonates. They contribute to characteristic skeletal grain associations (e.g. bryomol association or bryonoderm association). Modern bryozoans are common in cool-water environments ranging from temperate to subpolar settings, but also live in tropical settings and reefs (Sect.10.2.3.4). Bryozoan sands originate in water depths of several tens of meters. The site of active carbonate production and the accumulation of sands and muds is the deep shelf down to several hundreds of meters. Paleoenvironmental inferences to be drawn from bryozoan sediments of different ages must be handled with care because Paleozoic and post-Paleozoic bryozoan-rich carbonate deposits have strikingly different global distribution patterns. Paleozoic deposits spread through a wide range of latitudes, but were more common in the tropics except in times of significant cooling (Ordovician, –> 1; Late Paleozoic, –> 2). Post-Paleozoic and particularly Cenozoic deposits (–> 3) accumulated mainly in belts north and south of the equatorial zone. The plate displays examples of the Heterozoan Association. 1 Bryonoderm grain association: Characterized by abundant trepostomid bryozoans (B) associated with echinoderm fragments. Most echinoderms are cystoid plates (C). In this example the different size of the echinoderm grains is not caused by transport, but reflects primarily different sizes of morphologic elements of cystoids (Dullo 1992). Heterozoan sediments consisting of bryozoans and echinoderms were widely distributed in platforms and ramps during the mid- to Late Ordovician cooling episodes related to continental glaciations. Bryonoderm associations occur in cool-temperate and cold waters of low-latitude deeper-water shelves and high-latitude, deep- to shallow-water shelves. The sample comes from small bryozoan-cystoid mounds formed in paleolatitudes of > 60° S, most probably in cold-water environments. Late Ordovician (Wolaya Formation): Valentintörl, Carnic Alps, Austria. 2 Bryomol grain association in cold-water carbonates of Gondwana sediments. Modern bryomol associations are typified by a high percentage of bryozoans and mollusk shells. Pre-Cenozoic bryomol associations are characterized by abundant bryozoans associated with brachiopods. The skeletal packstone consists of bryozoans (B) sometimes growing on mollusk shells, brachiopods (BR), a few echinoderm fragments (E) and rare bivalve shells (not in the photograph). The sediment is strongly burrowed. Burrows are not compacted. Early cementation is missing, possibly because of the lack of aragonite particles providing the carbonate for cement formation. During the Permian, Tasmania lay at about 70° S in subpolar regions. The sea was periodically covered by sea-level ice. Terrigenous tillites pass laterally into marine carbonates that are associated with dropstones (Rao 1981). A cold-water environment is reflected by the bryozoan-dominated heterozoan grain association, the prevailing original calcite mineralogy of the bioclasts, the low-diversity fauna, and the abundance of bored brachiopods and echinoderms. Early Permian (Berriedale limestone): Southeastern Tasmania, Australia. 3 Bryomol grain association. The Tertiary sample exhibits the high abundance of bryozoans (B) and the common association of bryozoans with balanides (barnacles, BA) and corallinacean red algae (RA) characteristic of many Cenozoic bryomol carbonates. The limestone displayed in the thin section is an example of warm-temperate carbonates (Nebelsick 1989). This specific bryozoan-rich microfacies is limited to environments with low terrigenous input and low sedimentation rates. Bryozoans are represented by nodular encrusting celleporiform growth forms (CB) and thinner encrusting membraniform growth forms (MB). Most skeletal grains consist of calcite and High-Mg calcite. Aragonitic grains (gastropods, cheilostome bryozoans) have been leached, resulting in high moldic and open vuggy porosity (irregular white areas, e.g. at the top). In contrast to the sample shown in –> 2, early cementation was possible due to the dissolution of aragonitic skeletal grains. Paleolatitude was approximately 45° N. Late Tertiary (Eggenburgian, Early Miocene): Zogelsdorf, NW Vienna, Austria. –> 2: Sample provided by C.P. Rao (Hobart, Tasmania)
Plate 106: Cold-Water and Temperate Grain Associations
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Non-Tropical Grain Associations
Box 12.10. Selected studies on ancient non-tropical carbonates. Most studies use grain-association patterns to differentiate non-tropical and tropical carbonates. Papers that can serve as a guide in comparing samples with established grain association types are marked by an asterisk. Tertiary: *Bernecker and Weidlich 1990; Bernecker et al. 1997; Betzler 1995, 1997; Betzler et al. 1993, 1995, *1996, *1997, *2000; *Boreen and James 1992, 1995; *Brachert et al. 1993, *2001; *Braga et al. 2001; Burgess and Anderson 1983; Carannante et al. 1996; Clarke et al. 1996; *Cunningham and Collins 2002; *Franseen et al. 1997; *Friebe 1994; *Gammon et al. 2000; Gordon et al. 1994; *Hayton et al. 1995; *Holdgate and Gallagher 1997; James 1991, *1997; James and Bone 1992a, 1992b, *1994, *2000; Kamp et al. 1988; *Krautworst 1999; Lauriat-Rage et al. 1993; *Lukasik et al. 2000; MacGregor 1983; McGowran et al. 1997; *Mutti et al. 1999; Nebelsick *1989, *1992, 2000; Nelson 1978, 1999; *Nelson and James 2000; Nelson et al. *1988; Nicolaides and Wallace 1997; *Randazzo et al. 1999; Scudeler Bacelle and Renato 1988;
Unexpected grain associations and a few warnings: Many carbonate environments display a specific grain association related primarily to water temperature and associated biota. Nevertheless, unexpected grain associations are rather common in environments, where the prevailing temperature would point to a different association.
Shubber et al. 1997; Simone and Carannante 1988; Spjeldnaes and Moisette 1997; Vecsei and Sanders 1999 Cretaceous: Carannante and Simone 1987; Carannante et al. *1988, 1995, *1997; Gischler et al. 1994; *Simone and Carannante 1988; *Surlyk 1997 Carboniferous and Permian: *Beauchamp 1994; *Beauchamp and Desrochers 1997; *Brandley and Krause 1997; Draper 1988; *Ezaki et al. 1994; Fan Yingnian 1988; Hüneke et al. 2001; Martindale and Boreen 1997; *Rao 1981, 1988a, 1988b, *1991; Rao and Green 1982; *Samankassou 2002, *Stemmerik 1997, 2001; Weidlich 2002 Ordovician: *Brookfield 1988; *Dullo 1992; *Lavoie 1995; *Pope and Read 1997; Rao 1990, *1991; Rao and Wang 1990; Grimwood et al. 1999; Webby 2002
Several points must be taken into consideration when using grain associations as paleoclimatic proxies. Cool-water associations may be controlled by factors other than just water temperature. (1) Salinity control: The foramol sediments of the western Florida shelf situated approximately 25° N in tropical to subtropical warm waters can be explained
Plate 107 Skeletal Grain Association of Warm-Temperate and Tropical Warm-Water Shelf Carbonates Warm-temperate carbonates originate in a zone between the tropical warm-water zone with mean bottom water temperatures > 22 °C and the mid- to high-latitude cool-temperate zone with mean temperatures from about 10 to 18 °C. The first zone is aragonite/High-Mg calcite-dominated and characterized by calcareous green algae and coral reefs, benthic foraminifera, and strong early cementation. The second zone is calcite-dominated and characterized by abundant bryozoans, bivalves, barnacles, echinids and foraminifera, and only weak cementation. For the Cenozoic, the boundary between the tropical zone and the warm-temperate zone is commonly drawn at the occurrence/absence of coral reefs and distribution patterns of larger foraminifera genera (Hallock 1986: tropical zone; Halfar et al. 2000: warm-temperate zone. See Sect. 10.2.2). The boundary between the warm-temperate and the cold-temperate zone is indicated by a marked decrease in biotic diversity and the complete absence of green algae and ooids. 1
Bryozoa-Foraminifera association. Many Miocene shallow-marine carbonates are characterized by an assemblage of bryozoans, benthic foraminifera and corallinacean red algae. The thin section exhibits a bryozoan packstone rich in celleporiform and branching zoaria. Bryozoans (B) are competing with coralline red algae (RA). Other fossils are rotaliid larger foraminifera (Heterostegina; H) and serpulid worms (S; Ditrupa). Heterosteginid foraminifera occur both in tropical and warm-temperate environments. The Heterozoan Association with corallinacean red algae and serpulids points to warm-temperate conditions. This is in agreement with the approximate paleolatitude of 45° N. Late Tertiary (Middle Miocene): Holy Cross Mountains, central Poland. 2 Foraminifera-Bryozoa association. The floatstone exhibits larger foraminifera considered as light-dependent because of inferred algal endosymbionts. Orbitoid foraminifera are Orbitoides apiculata Schlumberger (O), Lepidorbitoides socialis (Leymerie) - L, and Siderolites (SI); alveolinids are Nummufallotia cretacea (Schlumberger) - N. Other skeletal grains are bryozoans (B) and serpulids (Ditrupa; S). Black grains are nodular corallinacean red algae (RA). The sample comes from carbonates formed in subtidal areas between rudist reefs. The biotic composition of the Heterozoan Association as well as the paleolatitudinal position (›30° N) point to warm-water conditions in the transition zone between tropical and warm-temperate regions. Late Cretaceous (Late Maastrichtian): Sierra de Urbasa, Province Navarra, northern Spain. –> 1: Studencki 1988; 2: Schwentke and Wiedmann 1985
Plate 107: Warm-Temperate and Tropical Skeletal Grain Associations
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by high or variable salinity controlling the biotic spectrum and producing a ‘foramol’ association in warmwater environments (Hallock and Schlager 1986). (2) Nutrient control: A nutrient-surplus caused by local upwelling can prevent the development of a characteristic chlorozoan biota with corals and calcareous green algae (Hallock and Schlager 1986). An ancient example is the foramol facies of the Late Cretaceous Lacazina limestone of northern Spain (Gischler et al. 1994). (3) Depth control: One of the major problems in the use of skeletal grain associations as indicators of paleoclimatic or paleolatitudinal conditions is the fact that water temperature changes observed from low to high latitudes can also be observed with increasing water depths (Brookfield 1988; Carannante et al. 1988). At low latitudes the 15 °C, 10 °C and 5 °C isotherms used as boundaries between tropical, warm-temperate and cold-temperate shelf carbonates lie at increasing water depths beneath the sea surface. As a consequence, deep-water shelf carbonates formed at low latitudes may be similar in grain association to shallow-water carbonates forming at high latitudes (Beauchamp 1994). (4) Transport control and mixing of grain types: Any evaluation of grain associations of non-tropical shelf carbonates should take into account the possibility of up- and downslope transport and resedimentation of grains on open high-energy shelves or ramps which are often affected by currents. Particular caution is needed for grainstones and packstones formed in temperate and cool-water shelves whose composition might reflect a mixture of grains originating from bathymetrically and climatically different sources or may include relict grains of different age. Basics: Grain association analysis Beauchamp, B. (1994): Permian climatic cooling in the Canadian Arctic. – Geological Society of America, Special Paper, 288, 299-246 Beauchamp, B., Desrochers, A. (1997): Permian warm-to very cold-water carbonates and cherts in Northwest Pangea. – In: James, N.P., Clarke, A.D. (eds.): Cool-water carbonates. – SEPM Special Publications, 56, 327-347 Betzler, C., Brachert, T., Nebelsick, T.C. (1997): The warm temperate carbonate province. A review of facies, zonations, and deliminations. – Courier Forschungsinstitut Senckenberg, 201, 83-99 Brachert, T.C., Hultzsch, N., Knoerich, A.C., Krautworst, U.M.R., Stückrad, O.M. (2001): Climatic signatures in shallow-water carbonates: high-resolution stratigraphic markers in structurally controlled carbonate buildups (Late Miocene, Southern Spain). – Palaeogeography, Palaeoclimatology, Palaeoecology, 175, 211-237 Brookfield, M.E. (1988): A mid-Ordovician temperate carbonate shelf: the Black River and Trenton Limestone Groups of southern Ontario, Canada. – Sedimentary Geology, 60, 137-153
Paleowater Depth
Carannante, G., Cherchi, A., Simone, L. (1995): Chlorozoan versus foramol lithofacies in Upper Cretaceous rudist limestones. – Palaeogeography, Palaeoclimatology, Palaeoecology, 119, 137-154 Carannante, G., Esteban, M., Milliman, J.D., Simone, L. (1988): Carbonate lithofacies as paleolatitude indicators: problems and limitations. – Sedimentary Geology, 60, 333346 Halfar, J., Godinez-Orta, L., Ingle, J.C. (2000): Microfacies analysis of recent carbonate environments in the Southern Gulf of California, Mexico - a model for warm-temperate to subtropical carbonate formation. – Palaios, 15, 323-342 Hayton, S., Nelson, C.S., Hood, S.D. (1995): A skeletal assemblage classification system for non-tropical carbonate deposits based on New Zealand Cenozoic limestones. – Sedimentary Geology, 100, 123-141 Hüneke, H., Joachimski, M., Buggisch, W., Lützner, H. (2001): Marine carbonate facies in response to climate and nutrient level: The Upper Carboniferous and Permian of Central Spitsbergen (Svalbard). – Facies, 45, 93-136 James, N.P. (1997): The cool-water carbonate depositional realm. – In: James, N.P., Clarke, A.D. (eds.): Cool-water carbonates. – SEPM Special Publications, 56,1-20 Lees, A. (1975): Possible influences of salinity and temperature on modern shelf carbonate sedimentation. – Marine Geology, 19, 159-198 Lees, A., Buller, A.T. (1972): Modern temperate-water and warm-water shelf carbonate sediments contrasted. – Marine Geology, 13, M67-M73 Nelson, C.S. (1988): An introductionary perspective on nontropical shelf carbonates. – Sedimentary Geology, 60, 3-12 Randazzo, A.F., Müller, P., Lelkes. G., Juhász, E., Hámor, T. (1999): Cool-water limestones of the Pannonian basinal system, Middle Miocene, Hungary. – Journal of Sedimentary Research, A69, 283-293 Schlanger, S.O. (1975): The geographic boundary between the coral-algal and the bryozoan-algal facies: a paleolatitude indicator. – 9th International Congress of Sedimentology, Nice, 187-193 Simone, L., Carannante, G. (1988): The fate of foramol (‘temperate-type’) carbonate platforms. – Sedimentary Geology, 60, 347-354 Further reading: K021 (cold-water carbonates), K171 (water temperature)
12.3 Assessing Water Depths In biological and ecological textbooks water depth is discussed in terms of environmental constraints affecting the organisms or zonation patterns of benthic communities. A separate chapter on water depths is generally missing, as opposed to geological textbooks that try to answer the geologist’s favorite question: ‘How deep was the water?’ The simple answer ‘shallow’ or ‘deep’ is no longer sufficient for sequence stratigraphy and should be replaced by more precise statements on the water depth within the shallow or deep depositional settings. A reliable sequence stratigraphy depends
Recognizing Water Depth
greatly upon the reliability of water-depth indicators and needs bathymetrically well-characterized reference beds that reflect a defined sea-level position. The interpretation of cyclic sedimentation patterns requires paleobathymetric pinpoints. Bathymetric subdivision of benthic marine settings The morphology of the sea bottom and physically or biologically defined boundaries are used to differentiate environments situated at various water depths (Sect. 2.3.1). Physically defined boundaries are given by the tidal range, the fair-weather wave base and the storm wave base. These levels are characterized by changes in sediment composition and texture. Biological boundaries are given by environmental constraints that govern the depth distribution and zonation of benthic organisms. The most important limiting ecological factors are light (Sect. 12.1.4.1) and seawater temperature (Sect. 12.1.6). The absolute depth of these boundaries is exceedingly variable. Hydrostatic pressure may be a major control in depth distribution of benthic organisms when other parameters do not vary.
12.3.1 Hints to Paleowater Depths Paleowater depths can be deduced from various evidence, including paleontological and microfacies criteria, mineralogical and geochemistry data, sedimentological features, and stratigraphic and geomorphological aspects. Bathymetry and shallow-marine zonation patterns were summarized by Liebau (1980), various methodological approaches to paleobathymetric interpretations by Hallam (1967) and Luterbacher (1984). The following text lists some criteria that assist in the assessment of bathymetric zones. Most features characterize ‘shallow’ environments comprising intertidal and subtidal shelves. Specific criteria of ‘deep-water’ carbonates are discussed in Sect. 15.8. Biological evidence The evaluation of some of the criteria listed below is an integrated part of microfacies analysis and SEM studies (calcareous algae, larger foraminifera, microborers, micro-encrusters). Other criteria require extensive palecologically-oriented investigations. • Calcareous algae: Depth distribution patterns of green and red algae are one of the best indicators of ancient depositional depths because of their strong light dependence. Untransported dasyclad green algae point to very shallow depths (tens of centimeters) down to only a few meters (Sect. 10.2.1.7). They record photic bottom conditions and occur within the tidal zone and
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the upper subtidal zone. Depth range of udoteacean green algae is broader and extends from a few tens of centimeters to some tens of meters (Sect. 10.2.1.5). The lower limit of the chlorocline (defined by the range of modern non-dasyclad green algae) is about 100 m. In the context of the facies interpretation of limestones, this boundary is sometimes used to separate upper and lower subtidal depositional settings. Corallinacean red algae, common in Cretaceous and Cenozoic limestones, occur both in shallow and deeper waters, in tropical settings between 20 and 80 m, in cool-water settings down to more than 200 m (Sect. 10.2.1.2). Squamariacean red algae, known since the Early Cretaceous, occur at water depths down to about 50 m. • Stromatolites: The formation of laminated microbial deposits is not as strongly light-dependent as the depth distribution of calcareous green and red algae. Ancient stromatolites were formed in depth ranges of meters to tens of meters. Growth forms, the composition and lamination of stromatolites is used to differentiate intertidal/subtidal and deeper subtidal environments (Sect. 9.1.4). • Benthic foraminiferal biofacies is useful for estimating paleobathymetry in Cenozoic and Mesozoic subtidal shelf carbonates (Sect. 10.2.2). The estimates made are based on biofacies models reflecting distribution in the modern ocean. The bathymetric arrangement of biofacies is regarded as a response to oceanic watermass structure. Some caution, however, is necessary because ancient taxa may have had broader bathymetric ranges. • Larger foraminifera: Many larger foraminifera can be taken as good evidence of shallow photic environments because of their endosymbiotic associations (Hottinger 1988). • Ratio of foraminiferal groups: Various diagrams describing the proportion of benthic foraminifera (calcareous versus agglutinated groups; ratio of agglutinated, calcareous and planktonic groups; plankton/ benthos ratio) are successfully used in estimating paleodepth conditions in shelf and deep-sea environments. • Ostracods: The group occurs from very shallow waters to bathyal regions (Sect. 10.2.4.7). Many benthic groups have developed morphological adaptions to life on the deep and very deep aphotic ocean bottom, allowing bathymetric zones to be distinguished (Benson 1984). • Zooxanthellate corals: Most modern scleractinian corals with endosymbiotic primary producers occur in shallow depths of a few tens of meters in water generally less than 100 m (Sect. 10.2.3.3). Fossil scleractinian reef corals and coral reefs are often regarded as evidence of well-lighted shallow environments. Considering the strong control of reef corals by factors other
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than light and symbiosis, these interpretations argue for support by additional criteria (e.g. growth form patterns, community structure). Zooxanthellate symbiosis may have been independently established several times in scleractinian coral groups. The existence of zooxanthellate symbiosis in Paleozoic rugose corals is strongly debated. SEPM Special Publication 72 (Kiessling et al. 2002) contains much data on paleolatitudinal distribution and the occurrence of Phanerozoic reef corals in shallow or non-shallow environments. These data should be checked. • Distributional patterns: Water depth on its own does not directly control distributional patterns, but bathymetric inferences may exist from environmental variables (e.g. temperature, salinity, substrate) that limit distribution with increasing depth. Benthic organisms exhibit distinct distributional patterns from the shoreline over the shelf to the shelf break. Gradient analyses of fossil benthic assemblages (e.g. mollusks, brachiopods) provide indications of relative water depths. • Zonation patterns of microborers: Endolithic microborers are excellent paleobathymetric indicators for euphotic and dysphotic depths of shelves, reefs and ramps and also allow aphotic regions to be recognized (Sect. 9.3.4; Fig. 9.16). Zonation patterns of ancient endoliths show a high coincidence between paleodepths derived from the composition of the endolithic associations and the paleodepths accounting for the paleogeographical situation and facies distribution, independent of the age of the examples. Comparative associations occurring in comparative paleowater depths are known from the Early Paleozoic to the recent. • Micro-encrusters: Distributional patterns of millimeter- to centimeter-sized encrustations resulting from the growth of various heterotrophic and autotrophic organisms offer a good chance for recognizing bathymetrically different environments of carbonate ramps above the fair-weather wave base, between the wave base and the storm wave base, and below the storm wave base (Sect. 9.2.3; Fig. 9.10). • Trace-fossil communities are strongly bathymetrically controlled due to their behavioral response to a bathymetric gradient in the food supply and oxygen. Paleobathymetric estimations based solely on trace fossil assemblages can have pitfalls, because factors other than nutrients (e.g. oxygenation and sedimentation rates) are also major controls on the composition of the communities (see Frey et al. 1990 for a discussion). Microfacies evidence Specific grain types and fabrics can be used to differentiate very shallow, shallow, deeper and deep depositional settings. Fine-grained micritic matrix originates
Recognizing Water Depth
everywhere. If you use the prevailing grain types of a limestone as bathymetric indicators, you should consider that the place of origin of sedimentary particles and the depth of the final resting place can differ. • Ooids: Accumulations of calcareous tangentiallystructured ooids originate today in very shallow and shallow tropical waters down to a few meters of water depth (down to about 15 m). See Sect. 4.2.5. • Aggregate grains: Depth distribution is similar to that of ooids. Many grapestones are formed in very shallow depths of only a few tens of centimeters. • Cortoids: The existence of micrite envelopes around skeletal grains is not an unmistakable indication of shallow-water deposition (Sect. 4.2.3), but ancient grainstones consisting of a high percentage of cortoids occurring in association with calcareous green algae and mollusk grains can be taken as proof of shallow subtidal depths in euphotic and dysphotic environments within a range of some tens of centimeter to a few meters. • Oncoids: Marine carbonate oncoids are formed in the shallow euphotic and dysphotic zone, both in intertidal and subtidal environments (Sect. 4.2.4.1). Cyanoids point to water depths within the meter range. Differences in biotic composition (algae, microbes, foraminifera), lamination and nuclei offer possibilities of distinguishing oncoids formed in shallow or deeper settings (Box 4.10). Ferruginous or phosphatic non-carbonate oncoids are candidates for ‘deepwater’ but not all candidates pass the test. • Rhodoids occur from intertidal to deep subtidal environments (Sect. 4.2.4.2). They originate in shallow and deep settings and often characterize the deeper part of the photic zone and the dysphotic zone below the wave base. The depth range of modern tropical rhodoids comprises tens of meters, often around 80 m, but may be extended to more than 180 m. Macroids are common at depths of several tens of meters. Non-tropical rhodoids are formed at various depths. Temperate-cool water rhodoids occur between about 5 and 150 m with a maximum between 20-40 m; polar cold-water rhodoids are known from depths of about 60 to 100 m. Ancient rhodoids bear a high potential as paleodepth indicators, if the biotic composition and the growth strategies are thoroughly analyzed. Depositional fabrics • Open-space and fenestral fabrics are excellent paleodepth indicators, but require detailed studies of the subcategories and the fabric types (Sect. 5.1.5). Birdseyes are good clues to water depths within a centimeter range; stromatactis-type structures have a wider
Water Depth on a Carbonate Ramp
bathymetric range most probably comprising both photic and dysphotic settings. The common occurrence of Stromatactis mounds on deep ramps is regarded as clue for water depths of more than 100 m. • Bioturbation fabrics reflect differences in nutrients, oxygen and substrate that may differ at different water depths. Abundance and burrowing types are used to distinguish shallow- and deep-water environments (Sect. 5.1.4). Mineralogical evidence • Carbonate minerals: High hydrostatic pressure, low water temperature and high partial pressure of CO2 lead to dissolution patterns of aragonite and calcite that characterize deep-water levels (Sect. 2.4.5.6). • Authigenic minerals: Syndepositional diagenetic minerals, e.g. glauconite, chamosite or phosphorite are often used in defining water depths although this approach is strongly disputed (Amorosi 1992).
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Sedimentological evidence • Sedimentary structures: Physical depth boundaries, e.g. wave base and storm wave base are reflected by specific sedimentary structures as well as microfacies criteria (Sect. 12.1.2.1; Demicco and Hardie 1994). Stratigraphic evidence • Estimates of absolute paleowater depth values are rare, but may be calculated using stratigraphic evidence if the original paleorelief between the platform and the basin is preserved and compaction is considered (Kenter 1990).
12.3.2 Case Study: Assessing the Water Depth of a Carbonate Ramp Fig. 12.12 and Pl. 108 demonstrate the potential of calcareous algae and other biota in recognizing rela-
Fig. 12.12. Depth zonation along a Late Viséan ramp profile in the Béchar Basin, Western Algeria. The depth zonation is based on the distribution of algal assemblages and other biota and the recurrence of distinct associations in an analogue outcrop study of microfacies, paleontological data and limestone texture. Note the different distribution of green and red algae. Most bedded grainstones and packstones overlying the mound facies originated in the euphotic zone corresponding to water depths within the meter- to about 20 m range. Most mounds consisting of bafflestones, wackestones and mudstones were formed in the aphotic zone characterized by the absence of algae. The numbers refer to bathymetric zones defined by qualitative and quantitative biotic composition. Oolitic and bioclastic carbonates formed in zone 1 exhibit large-scale cross stratification. Zone 2 was dominated by crinoids which were frequently reworked by storms. Zone 3 is characterized by brachiopod shell banks above the coral thicket of zone 4. The upper part of zone 5 shows evidence of reworking. The zones 5 to 7 are parts of the mound facies. The ramp is differentiated not only by biota but also by the dominant sediment: Oolitic and bioclastic sands prevail on the upper ramp, crinoid banks on the mid ramp, and massive dome-shaped mounds on the deep ramp. The reservoir potential occurs in ooid and crinoid grainstones overlying the mounds, formed in the bathymetric zones 1 and 2 (Madi et al. 2000). Width of the bars indicate relative abundance of major fossil groups. FWB - fair-weather wave base, SWB - storm wave base. After Madi et al. (1996).
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tive water depths. The western part of the Algerian Béchar Basin contains Mississippian carbonates that form superposed shoaling-upward sequences consisting of sponge-bryozoan-crinoid reef mounds and bedded shelf carbonates deposited on a ramp. Recognizing paleobathymetric zones is the basis for a facies model that assists in the hydrocarbon exploration of the up to 1200 m thick sequence (Bourque et al. 1995; Madi et al. 1996). Basics: Paleowater depths Bathurst, R.G.C. (1967): Depth indicators in sedimentary carbonates. – Marine Geology, 5, 447-471
Water Depth on a Carbonate Ramp
Benson, R.H. (1984): Estimating greater paleodepths with ostracodes, especially in past thermospherics oceans. – Palaeogeography, Palaeoclimatology, Palaeoecology, 48, 61-106 Cisne, J.I., Gildner, R.F. (1988): Measurement of sea level change in epeiric seas: The Middle Ordovician Transgression in the North American Midcontinent. – SEPM Special Publication, 42, 217-225 Ekdale, A.A. (1988): Pitfalls of paleobathymetric interpretations based on on trace fossils assemblages. – Palaios, 3, 464-472 Geister, J. (1984): Die paläobathymetrische Verwertbarkeit der scleractinen Korallen. – Paläontologische Kursbücher, 2, 46-95 Hallam, A. (ed., 1967): Depth indicators in marine sedimen-
Plate 108 Recognizing Paleowater Depths The Early Carboniferous carbonate ramp of the Béchar Basin in western Algeria offers the possibility of checking the value of algae and other microfossils as indicators of bathymetric zones and in distinguishing different parts of a homoclinal ramp. Because of their sensitivity to the various spectra of visible light, the distribution of small encrusting green and red algae can be used to determine depth-related zones. The absence of green algae is interpreted as indicating deposition within the aphotic zone, whereas the presence of red algae combined with the absence of green algae is thought to indicate the lower part of the photic zone with little light (dysphotic). The abundance of dasyclad green algae is indicative of the upper part of the photic zone. The example reflects the profile of many other Early Carboniferous ramps with low-energy mud-rich ‘Waulsortian’ mounds in the lower part (see Pl. 144), crinoid banks in the central part and high-energy calcareous sands in the upper part (Wright 1986; Jeffery and Stanton 1996). The plate displays facies-diagnostic algae and foraminifera characterizing the shallow photic and dysphotic parts of the Viséan ramp. Not shown are facies-diagnostic biota of the mound facies (cf. Fig. 12.12). Ungdarellacean and stacheinacean red alga are characteristic constituents of Carboniferous platforms and ramps (see Sect. 16.3.2.1; Box 10.2). The fossils shown on the plate occur in bedded skeletal grainstones and packstones deposited at shallow well-illuminated water depths (–> 1 and 9) and in somewhat deeper, less illuminated environments below the fair-weather wave base (–> 2–8, 10). 1 2
Algal-foraminiferal grainstone with the dasyclad alga Koninckopora. Upper ramp. Photic bathymetric zone 1. Foraminiferal grainstone with Saccaminopsis. The uniserial test consists of globular and ovate chambers. The Saccaminopsis microfacies is common in Early Carboniferous shelf carbonates. Massive skeletal grainstone forming a cap on the top of the mounds. Dysphotic bathymetric zone 5. 3-4 Green alga Exvotarisella Elliott. The thallus corresponds to a cylindrical tube. The central part is subdivided by irregular cavities formed by thick pseudosepta (PS). The alga is a common element of Late Viséan-Namurian shelf carbonates. Lower ramp. Dysphotic bathymetric zone 5. 5 Green alga Issinella characterized by an erect cylindrical shape and a fibro-radial wall, perforated by thin unramified branches. The alga is known from the Middle Devonian to the Late Viséan. Lower ramp. Upper part of the dysphotic bathymetric zone 5. 6 Red algal grainstone with Fasciella encrusting bryozoans (B). Characteristic criteria of the alga are irregular concentric layers of elongated, undulating cells and a yellowish hyaline structure. Known from the Viséan to the Bashkirian. Fasciella is common in high-energy environments. Lower ramp. Dysphotic bathymetric zone 4. 7-8 Red alga Ungdarella Maslov with microbial encrustations (ME) characterized by ramified growth, rectangular and subquadratic cells and yellowish-hyaline structure. Known from the latest Viséan to the Early Permian. Massive crinoid mound facies. Lower ramp. Dysphotic bathymetric zone 4. 9 Red alga Epistacheoides Petryk and Mamet. The encrusting form resembles Ungdarella, but differs in the vesicular central and the rectangular peripheral cells. This alga and similar taxa are common constituents of Viséan and Middle Carboniferous carbonate ramps. Upper ramp. Photic bathymetric zone 1. 10 Red alga Stacheia Brady, a thin encrusting form consisting of one or two rows of cells. Wall structure yellowish-hyaline. Known from the Viséan and the Middle Carboniferous. Massive skeletal grainstone forming a cap on top of the mounds. Lower ramp. Dysphotic bathymetric zone 4. –> 1–10 from Madi et al. 1996
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tary environments. – Marine Geology, 5, 327-567 Hottinger, L. (1984): Zur Tiefenverbreitung von Großforaminiferen – Paläontologische Kursbücher, 2, 140-147 Kenter, J.A.M. (1990): Carbonate platform flanks: slope angle and sediment fabric. – Sedimentology, 37, 777-794 Liebau, A. (1980): Paläobathymetrie und Ökofaktoren: Flachmeer-Zonierungen. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 160, 173-216 Luterbacher, H.P. (ed., 1984): Paläobathymetrie: Konzepte, Methoden, Möglichkeiten, Probleme. – Paläontologische Kursbücher, 2, 225 pp. Luterbacher, H.P., Van Hinte, J.E. (eds., 1984): Deep-water paleobathymetry. – Palaeogeography, Palaeoclimatology, Palaeoecology, Special Issue, 48 Madi, A., Bourque, P.A., Mamet, B.I. (1996): Depth-related ecological zonation of a Carboniferous carbonate ramp: Upper Viséan of Béchar Basin, Western Algeria. – Facies, 35,59-80 Oberhänsli, H. (1984): Stabile Isotopen: Hilfsmittel für die Paläobathymetrie. – Paläontologische Kursbücher, 2, 96103 Porrenga, D.H. (1967): Glauconite and chamosite as depth indicators in marine sediments. – Marine Geology, 5, 495501 Seilacher, A. (1984): Bathymetrie von Spurenfossilien. – Paläontologische Kursbücher, 2, 104-123 Vogel, K., Balog, S.-J., Bundschuh, M., Gektidis, M., Glaub, I., Kutschinna, J., Radtke, G. (1999): Bathymetrical studies in fossil reefs, with microendoliths as paleoecological indicators. – Profil, 16, 181-191 Further reading: K169
12.4 Looking for Seismic Events Earthquakes contribute to the destruction of carbonate sequences, soft sediment deformation and dewatering, and the redeposition of marine sediment (Shiki and Cita 2000). Seismic shocks provoke limited internal fracturing of semi-lithified limestone beds, resulting in seismites characterized by a progressively higher degree of disruption and brecciation from the bottom to the top of the bed (Seilacher 1984; Plaziat et al. 1990; Martire 1996). Earthquakes can cause sediment deformation, create water escape structures, sedimentary dikes, sedimentary breccias, recumbent folds, ball-andpillow structures, and convolute bedding. But it must be noted that many of these features also originate in shallow-marine high-energy settings without earthquake-driven causes. Criteria used in evaluating ancient seismic activity in carbonate rocks are intraclast and breccia beds, and micro-scale dewatering structures. Intraclast limestones: Laterally continuous limestone beds composed of intraclast wackestones and packstones, and associated with deformation structures, intraformational conglomerates and breccias may point
Influence of Earthquakes
to earthquake activity. Interesting case studies interprete Late Devonian intraclast carbonate parabreccias consisting of protointraclasts as seismites (Spalleta and Vai 1984), and extended stratigraphic carbonate breccia beds in the Triassic Muschelkalk as seismites (Rüffer 1996). Unusual intraclastic limestones in Early Triassic carbonates formed after the end-Permian mass extinction have been explained as the result of tsunami waves related to seismic catastrophes (Wignall and Twitchett 1999). Small-scaled dewatering structures possibly triggered by submarine earthquakes have been described from micritic limestones, and may be recorded by • patchily distributed very small spar-filled voids within microbioclastic wackestones or mudstones. These structures have been interpreted as result of pore-water escape during partial sediment consolidation, • spar- or sediment-filled flame structures (Plaziat et al. 1990), • vertical, crinkled spar-filled micro-cracks, width about 1 mm or less, • close-set parallel vertical micro-faults. Most examples described deal with carbonates formed on epeiric ramps, at platform-edges or on upper slopes. Basics: Paleoseismic events Ettensohn, F.R., Rost, N., Brett, C.E. (eds., 2002): Ancient seismites. – Geological Society of America, Special Paper, 359, 190 pp. Kullberg, J.C.,Oloriz, F., Marques, B., Caetano, P.S., Rocha, R.B. (2001): Flat-pebble conglomerates: a local marker for Early Jurassic seismicity related to syn-rift tectonics in the Sesimbracea (Lusitanian, Portugal). – Sedimentary Geology, 139, 49-70 Plaziat, J.-C., Purser, B.H., Philobbos, E. (1990): Seismic deformation structures (seismites) in the syn-rift sediments of the NW Red Sea (Egypt). – Bull. Soc. Géol. France, ser. 8, 6, 419-434, Paris Pratt, B.R. (2002): Tepees in peritidal carbonates: origin via earthquake-induced deformation, with examples from the Middle Cambrian of western Canada. – Sedimentary Geology, 153, 57-64 Rüffer, T. (1996): Seismite im Unteren Muschelkalk westlich von Halle (Saale). – Hallesches Jahrbuch für Geowissenschaften, B18, 119-130 Seilacher, A. (1984): Sedimentary structures tentatively attributed to seismic events. – Marine Geology, 55, 1-12 Shiki, T., Cita, D.S. (2000): Sedimentary features of seismites, seismo-turbidites and tsunamiites - an introduction. – Sedimentary Geology, 135, 7-9 Spalleta, C., Vai, B. (1984): Upper Devonian intraclast parabreccias interpreted as seismites. – Marine Geology, 55, 133-144 Further reading: K086
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13 Integrated Facies Analysis
To understand the formation and diagenesis of carbonate rocks mineralogical and geochemical data derived from the study of non-carbonate constituents, trace elements, and stable isotopes must be integrated. Other indicators for depositional and diagenetic conditions are organic matter and organic carbon. The following shows how these data can be successfully combined with data based on microfacies analyses. The first section deals with acid-insoluble residues in carbonate rocks and with authigenic minerals (oxides, silicates, sulfides, sulfates and phosphorites) formed after the deposition of the sediment. The second section discusses the value of trace elements in tracing the depositional and diagenetic history of limestones, and the importance of stable isotopes studied in the context of microfacies analyses. The third section gives a brief overview of organic matter in carbonate rocks.
13.1 Non-Carbonate Constituents Non-carbonate constituents of limestones and dolomites include clastic terrigenous minerals as well as authigenic minerals. The study of clay minerals in acidinsoluble residues has a high potential for recognizing changes in erosion, climate changes and sea level and estimating the influx of siliciclastic material into shallow and deep-marine carbonate basins. Authigenic minerals grow after sediment deposition during diagenesis. They describe the course of diagenesis and are proxies for varying chemical and physical conditions.
13.1.1 Insoluble Residues (IR): Clay Minerals and Detrital Quartz The input of terrigenous material into carbonate environments takes place by eolian transport, fluvial transport (e.g. during flash-floods from wadis), or by erosion of underlying rocks (e.g. in tidal zones).
Common clay minerals in mudstones and also in limestones are those of the smectite group, kaolinite, illite and mixed-layer illite/smectites and chlorite. Kaolinite and other clay minerals are often used as paleoclimatic proxies, but the post-depositional diagenetic alteration of clays related to the depth of burial and the pore-water geochemistry must be thoroughly considered (see review by Ruffell et al. 2002). As a consequence of different transport behavior clay minerals are deposited in the ocean at different distances from the coast. The relative abundance of kaolinite as a detrital mineral may reflect proximity to the sediment source and deposition in relatively nearshore settings, whereas smectite may be deposited in deeper more offshore settings. The relative proportions of illite/kaolinite to smectite in marine deposits have been used to infer shallowing/regressive and deepening/transgressive episodes. Techniques: The use of IR as paleoenvironmental indicators requires data on the amount of non-carbonate constituents and the mineralogical composition. The non-carbonate fraction of carbonate rocks is commonly extracted by acid digestion with a variety of acids and chelating agents (Müller 1967), or by the use of acidic ion exchange resins (French et al. 1984). All these treatments may attack or even dissolve clay minerals. Weak organic acids may be more useful than diluted mineral acids. Clay and silt fractions are separated by sedimentation methods (see McManus 1988). Insoluble residues in carbonate rocks: All the clay minerals occurring in clays or sandstones are also known from limestones. A diagenetic origin of kaolinite found in limestones can be neglected because of the protective effect of the carbonate matrix against the formation of an acid milieu. The most common marine clay mineral is illite, which may be subjected to multiple redeposition. A diagenetic origin of illites found in limestones can not be excluded. Mixed-layer minerals are often the result of the transformation of illite to smectites in soils and subsequent transport to marine
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_13, © Springer-Verlag Berlin Heidelberg 2010
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basins. Detrital quartz grains within the silt fraction should be specifically studied with respect to crystal habit and surface features (Krinsley and Doornkamp 1973). Significance. The mineralogical composition of insoluble residues of carbonate rocks is commonly explained as the result of terrigenous input. However, the possibility of transformation or neoformation of minerals in the course of burial diagenesis must be kept in mind. The amount of clay in limestones varies strongly within a range of less than 1% and more than 10%. Clay content is often higher in micritic rocks than in high-energy grain-supported limestones that are winnowed with respect to fine-grained constituents. These differences are controlled by hydrodynamic conditions and depositional settings. Combined investigations of insoluble residues and microfacies have been successfully applied to facies analysis (Box 13.1) as well as to applied studies of carbonate rocks (see Chap. 17). Basics: Insoluble residues in limestones Adatte, T., Rumley, G. (1984): Microfaciès, minéralogie, stratigraphie et évolution des milieux de depôts de la plateforme berriaso-valanginienne des régions de Sainte-Croix (VD), Cressier et du Landeron (NE). – Bulletin de la Societé Neuchâteloise des Sciences Naturelles, 107, 221-239 Bausch, W. (1980): Tonmineralprovinzen in Malmkalken. – Erlanger Forschungen, Reihe B, 8, 78 pp. Bausch, W. (1996): Noncarbonates as controlling factor in reef growth and as a tool in reef stratigraphy (with examples from the Upper Jurassic of southern Germany). – Göttinger Arbeiten zur Geologie und Paläontologie, Sonderband, 2, 203-205 Bolliger, W., Burri, P. (1967): Versuch einer Zeitkorrelation zwischen Plattformkarbonaten und tiefermarinen Sedimenten mit Hilfe von Quarz-Feldspat-Schüttungen (mittlerer Malm des Schweizer Jura). – Eclogae geologicae Helvetiae, 60, 491-507 Cater, J.M.L. (1984): An application of scanning electron microscopy of quartz sand surface textures to the environmental diagnosis of Neogene carbonate sediments, Finestrat Basin, south-east Spain. – Sedimentology, 31, 717-731 Chamley, H. (1989): Clay sedimentology. – 623 pp., Berlin (Springer) Chamley, H., Proust, N.-J., Mansy, J.-L., Boulvain, F. (1997): Diagenetic and paleogeographic significance of clay, carbonate and other sedimentary componentes in the Middle Devonian limestones of the western Ardennes, France. – Palaeogeography, Palaeoclimatology, Palaeoecology, 129, 369-385 Dürkoop, A., Richter, D.K., Stritzke, R. (1986): Fazies, Alter und Korrelation der triadischen Rotkalke von Epidauros, Adhami und Hydra /Griechenland). – Facies, 24, 105-150 Piller, W., Mansour, A.M. (1994): Origin and transport mechanisms of non-carbonate sediments in carbonate-dominated
Insoluble Residues
Box 13.1. Use of insoluble residues (IR) of limestones in facies analysis. • Differentiating carbonate platforms and ramps. Landlocked platforms are more commonly affected by siliciclastic input than open-marine oceanic platforms whose carbonates are generally very poor in non-carbonate residues. Different parts of carbonate platforms can be characterized by different mineralogical composition of the insoluble residues (Adatte and Rumley 1984). • Distinguishing depositional settings. Diagrams of IR content versus the silt/clay ratios of limestones can separate limestones formed in different basins (Bausch 1987). IR contents increase in platform-slope-basin transverses from platform-edge reefs to near-reef basinal facies and to off-reef basinal sediments (Dürkoop et al. 1986). • Assessing sedimentation patterns. Bed-per-bed studies of limestone sequences reveal quantitative differences in the composition of IR that may reflect cyclic depositional patterns and allow regressive and transgressive phases to be defined. • Quantifying sedimentation rates. The proportion of carbonate and non-carbonate within limestone beds and bed thickness are useful for describing depositional rates. • Recognizing hidden discontinuities in limestone sequences. Abrupt and significant changes in IR contents and marked differences in the mineralogical composition and in grain sizes from bed to bed may indicate depositional breaks. • Understanding source areas and hinterland erosion. Changes in erosion and the topography of the source areas are reflected by clay mineral composition, the illite/kaolinite proportions, grain sizes and IR amounts (e.g. Weaver and Stevenson 1971). • Understanding conditions in transitional terrestrialmarine areas. Clay-mineralogical assemblages of marls at the top of shallowing-upward sequences reflect fluctuating freshwater and marine conditions affecting shallow-water carbonates (Deconinck and Strasser 1987). • Reconstructing transport mechanisms. SEM surface features of quartz grains are helpful in reconstructing source area and transport mechanisms of terrigenous input into carbonate environments (Cater 1984; Piller and Mansour 1994). • Correlating platform and basin carbonates: Compositional patterns of IR are used for stratigraphic correlation of shallow- to deep-marine carbonates on a regional scale (Gygi and Persoz 1986; Bausch 1996; Schweizer 1996). Changes in the amount of quartz and feldspars studied in IR and in thin sections of carbonates offer a possibility for recognizing time lines and time units (Bolliger and Burri 1967). • Approaching an understanding of major paleoclimatic differences. The abundance and composition of specific clay minerals in association with other proxies are used as key indicators for humid versus arid climates. The criteria used to define humid climate are the abundance of kaolinite, Th-enriched clays, siliciclastics, positive δ13C values and wood. Arid climate may be indicated by the abundance of illite, U and K-rich clays, evaporites and carbonates, and negative δ13C values (Ruffell et al. 2002).
Authigenic Minerals
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environments (Northern Safaga Bay, Red Sea, Egypt). – Abhandlungen der Geologischen Bundesanstalt Wien, 50, 369-379 Ruffell, A., McKinnley, J.M., Worden, R.H. (2002): Comparison of clay mineral stratigraphy to other proxy paleoclimate indicators in the Mesozoic of NW Europe. – Philosophical Transactions of the Royal Society London, A360, 675-693 Singer, A. (1984): The paleoclimatic interpretation of clays in sediments - a review. – Earth Science Review, 21, 251293 Velde, B. (1992): Introduction to clay minerals. Chemistry, origins, uses and environmental significance. – 208 pp., New York (Chapman and Hall) Further reading: K068, K153
13.1.2 Authigenic Minerals Authigenic minerals grow after sediment deposition during diagenesis. They include both replacive minerals and cements. Replacive minerals grow in the place of pre-existing minerals; they are formed from circulating pore fluids or they are alteration products of primary detrital grains. In contrast, cements are crystals growing within existing pore spaces. Excellent thinsection photographs of authigenic minerals can be found in Scholle (1979). 13.1.2.1 Silicification of Carbonates, Authigenic Feldspar and Glauconite Silicification and idiomorphic quartz A wide variety of sediments is commonly affected by silicification in deep-marine as well as in shallowmarine limestones. The major sources of silica in sediments are: (1) siliceous tests and skeletal elements of organisms, (2) river input of solutions from the weathering of continents in semi-arid climates (Laschet 1984), and (3) silica supplied in solution by hydrothermal volcanic systems. Silicification of carbonate rocks involves replacement of carbonate by silica as well as precipitation of pore-filling silica cement (e.g. Noble and Van Stempvoort 1989). Studies of silicified carbonates usually focus on the description of fabrics (e.g. Wilson 1966), the extent of silicification (minor, selective, pervasive), and the relative timing of silicification with respect to diagenetic events affecting the carbonate host rock (see Hesse 1990b for an overview; Pl.109/7, 8). Silicification fabrics may take the form of euhedral quartz (commonly >0.5 mm), drusy megaquartz (0.5 to 10s mm), microquartz (<10 µm) or chalcedony (10s to 100s mm, fibrous silica crystals forming radial or spherulitic structures).
Fig. 13.1. Authigenic idiomorphic quartz crystals in a lime mudstone formed in a hypersaline environment. Note the well developed euhedral crystal shape. Compare Pl. 109/5 for a SEM photograph. Late Cretaceous: Subsurface. Ras al Khaimah, United Arabian Emirates.
Selective or pronounced silicification of fossils or other carbonate grains is a common feature (Maliva and Siever 1988; Herrig 1993; Whittle and Alsharhan 1994; Martín Penela 1995). In thin sections silicification is recorded by idiomorphic quartz crystals or as chert appearing as a colorless microcrystalline aggregate similar in appearance to micrite or microspar, but with much lower birefringence. The term chert refers to a rock composed largely or entirely of fine-grained quartz. Cherts are formed in various environments including marine and freshwater (hypersaline lakes) and at subaerial exposure surfaces (Knauth 1979). Marine environments can be differentiated using microfacies criteria (Affolter 1997). Many chert nodules exhibit ghost structures after depositional textures and biota. In marine settings cherts occur as nodules varying in size from a few centimeters to about 30 cm, or in layers, often interbedded with shales or limestones. Most chert nodules were formed early during diagenesis. They are common in pelagic chalks. In shallow-marine areas (ramps, platforms) the input of silica from land is an important source (Laschet 1984). In shallow-marine settings chert nodules are sometimes related to mixing zones. The silica for nodules and bedded cherts occurring in pelagic deep-marine limestones is usually supplied by dissolution of sponge spicules and radiolarians. Recent pure siliceous ooze of the South Atlantic may be an example of a diagenetic environment with early chert formation (Bohrmann et al. 1994). Idiomorphic quartz crystals (Fig. 13.1; Pl. 109/5, 6) are often found in carbonates affected by saline or hypersaline pore waters (see Box 12.7). This situation is
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not limited to lagoons or restricted bays in arid climates, but can also originate from salinar solutions derived from salt diapirs underlying platforms and reefs (Reitner 1986). Authigenic quartz and feldspars are usually restricted to micritic limestones. Authigenic feldspars Idiomorphic crystals, euhedral outline, carbonate inclusions and specific optical criteria are characteristic for authigenic feldspars. Their mineralogy appears to be determined by the composition of the host sediment. The most common replacement feldspar is albite. Only a few reports of authigenic K-feldspars are known (e.g. Buyce and Friedman 1975). Euhedral albite crystals occur isolated or in association with authigenic quartz. The feldspar crystals are double terminated and many of them have long, welldeveloped prismatic crystal faces. The size of the crys-
Authigenic Minerals
tals varies greatly (long diameter about 30 to 300 µm). Diagenetic processes which could mobilize silica from within carbonate sequences are pressure solution of silicates, replacement of silicates by carbonates, and clay mineral diagenesis connected with the conversion of smectite clay minerals to illite. Glauconite This hydrous potassium-iron-alumo-silicate mineral today forms exclusively in marine environments and is found in carbonate as well as in siliciclastic sediments. In modern oceans glauconite occurs between 50 and 500 m and is abundant in mid-shelf to upper slope settings at depths between 50 and 300 m. Criteria. The mineral is typically found as rounded aggregates, as pellets consisting of very fine-grained scaled particles, as authigenic mineral within skeletal grains (e.g. echinoderms, foraminifera) or as surface
Plate 109 Common Authigenic Minerals in Limestones The plate displays some examples of authigenic non-carbonate minerals that are common in limestones: Pyrite (–> 1–4), quartz (–> 5, 6) and gypsum/anhydrite (–> 9). Aspects of silicification are shown in –> 7 and 8. 1 Pyrite: Framboidal pyrite in residual clay seams of stylolites. The rounded aggregates consist of many tiny euhedral crystals. The size distribution of framboidal pyrite is a measure of redox conditions within sediments (Wilkin et al. 1996). Compare Fig. 13.1. SEM photograph. Late Cretaceous: Subsurface, Ras al Khaimah, United Arab Emirates. 2 Pyrite: Well-crystallized pyrite in residual clay seams of stylolites. Compare Fig. 13.2. SEM photograph. Same locality as –> 1. 3 Pyritization of fossils. Monaxon sponge spicula. Parts of the siliceous spicula is replaced by pyrite, other parts by calcite. Late Jurassic: Northern Alps. Roman mosaic stone, Kraiburg am Inn, Bavaria, Germany. 4 Pyrite. Scattered pyrite crystals in a dolomitic mudstone. Same locality as –> 1. 5 Silicification. Authigenic quartz crystals in a lagoonal lime mudstone. Scattered ‘hexagonal’ replacement quartz crystals are not uncommon in fine-grained micritic limestones, and have been regarded as indications of hypersaline fluids (see Box 12.7). SEM photograph. Same locality as –> 1. 6 Silicification. The authigenic quartz crystals exhibit inclusions (arrows) of biogenic structures pointing to an early diagenetic replacement of the limestone. Middle/Upper Jurassic: Argelita north of Valencia, Spain. 7 Timing of silicification. The reef-building fossil Zondarella communis Keller and Flügel (stromatoporoid or stromatolite) is characterized by dome-shaped growth layers. The layers consist of densely-spaced discontinuous laminae (arrows) separated by micritic layers that have been replaced by blocky calcite (C) and euhedral dolomite crystals (D). Silicification (S) took place subsequent to dolomitization within solution voids. Silicification (chertification) of carbonate rocks, subsequent to dolomitization is a common feature (Whittle and Alsharhan 1994). Replacement of calcite by chert may occur at a much more rapid rate than that of dolomite (indicated by floating dolomite within chertified matrix). Ordovician (San Juan Formation, Late Arenigian): Las Lajas, Precordillera, western Argentina. 8 Silicification of rotaliid foraminifera: The silicification of the aragonitic shells took place subsequent to the redeposition of the shallow-marine bioclasts in deeper parts of a ramp, probably under shallow burial conditions. Early Tertiary: Southern Galala Basin, Egypt. 9 Evaporite-dolomite couplets formed within a sabkha environment. Alternating layers consisting of fine-grained dolomite (black) and gypsum. Requirements for the formation of supratidal sabkhas: (1) hot and arid climate, (2) sea-marginal location with adequate supply of saline water to the sabkha to develop evaporites, (3) supply of terrigenous mud, (4) protection from winds, and (5) low tidal range, allowing the groundwater level to be maintained within the zone of capillary rise. Crossed nicols. Early Tertiary (Eocene): Paris Basin, France. –> 8: Courtesy of J. Kuss (Bremen)
Plate 109: Common Authigenic Minerals in Limestones
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coatings on hard substrates. Rounded glauconite grains reflect the original shape of fecal pellets or rounding by transport. The bulk of glauconite pellets may be clay debris from the sea floor, that passed through the digestive tract of burrowing organisms. In thin sections glauconite grains are easily identifiable by (a) green, blue-green, yellowish-green or brownish-green colors both in ordinary and polarized light. (b) speckled greenish birefringence and (c) a very fine granular texture. Significance. Glauconite is usually regarded as an indicator of marine environment, relatively shallow deposition and slow sedimentation. The mineral is often concentrated at discontinuity surfaces indicating depositional breaks (Pl. 22/1, 2). Different mineralogical compositions of glauconites, distinguished by diffractometric parameters, appear to be related to specific sequence system tracts (Amorosi 1993). Any use of glauconite as an environmental proxy should consider three points: • Differentiation of authigenic and detrital glauconite grains. Glauconites are formed coevally or non-coevally with sedimentation, and intrabasinal or extrabasinal. Amorosi (1993) distinguished four major groups: (1) Intrabasinal authigenic glauconites formed in situ that have not undergone transport. (2) Intrabasinal detrital glauconite grains transported from submerged structural highs or derived from in situ reworking processes. (3) Extrabasinal detrital glauconite grains eroded from older stratigraphic levels. Glauconite grains occurring in association with terrigenous quartz may point to this possibility (Pl. 49/1; Pl. 75/6, 8). (4) Perigenic authigenic glauconites dispersed in an unconsolidated sediment on the sea floor, that have undergone transport from the area where they were formed by tides, storms or turbidity currents. • Glauconite grains are particularly common in ramp and platform carbonates of Cambro-Ordovician and Cretaceous age. This concentration points to specific and possibly non-actualistic conditions controlling the formation of the mineral. • Differences between modern and ancient environments. The occurrence of glauconite cannot be used a priori as an environmental indicator of either mid-shelf and deeper water and/or a slow rate of sedimentation because of striking differences between modern and ancient glauconite-bearing limestones. Chafetz and Reid (2000) showed that Cambro-Ordovician glauconite-rich strata were formed under shallow-water to tidal conditions.
Pyrite in Limestones
13.1.2.2 Sulfides: Pyrite Opaque in transmitted light, pyrite is most readily identified by reflecting light owing to its characteristic yellow-gold appearance. Authigenic pyrites in limestones are usually developed in the form of cubic euhedral crystals (Fig. 13.2; Pl. 109/2). Pyrite attracts sedimentologists, for the mineral is a valuable indicator of chemical processes (Wilkin et al. 1996) and diagenetic stages (Hudson 1992). Fossils preserved in pyrite are attractive for their beauty and for morphological details revealed by pyritization (Pl. 109/3; Canfield and Raiswell 1991). Most of the pyrite in sedimentary rocks is of diagenetic origin, although detrital and synsedimentary pyrite occurs, too. Authigenic pyrite commonly forms under reducing conditions replacing organic material or in close proximity to organic material. Pyrite is formed in normal marine, euxinic, and freshwater environments. Three principal factors limit ultimately how much pyrite is formed in sediments (Berner 1985): (1) the amount and reactivity toward bacterial sulfate reduction of organic matter supplied to the sediments; (2) the amount and reactivity toward H2S of detrital minerals supplied to the sediment; and (3) the availability of dissolved sulfate. The primary factor limiting pyrite formation in normal marine sediments, e.g. those deposited in oxygenated bottom waters, is organic matter. In euxinic environments pyrite can form at locations where no organic matter is deposited, because H2S is present everywhere in the bottom waters. Pyrite formation in euxinic basins is limited by the amount and reactivity of detritic iron minerals. Very little pyrite is formed in freshwater lake and swamp sediments
Fig. 13.2. Authigenic pyrite crystals grown within layers of cyanobacterial oncoids rich in organic matter. Note the welldeveloped crystal shape. SEM photograph. Late Cretaceous: Subsurface, Ras al Khaimah, United Arabian Emirates.
EvaporiteMinerals
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seal rocks and assist in the understanding of climatically controlled cyclic sedimentation patterns. Abundant evaporite minerals are chlorides (halite) and sulfates (gypsum and anhydrite). Gypsum only occurs near the earth’s surface, whereas anhydrite is formed at the surface and in the subsurface where it replaces gypsum. Gypsum may replace anhydrite on the uplift of evaporite sequences or after removal of overburden sediments. Fig. 13.3. Framboidal pyrite consisting of tiny euhedral crystals, formed within residual clays of stylolites. The clay layers are distinctly displaced by the growing pyrite. Compare Pl. 109/1. SEM photograph. Cretaceous: Subsurface, Ras al Khaimah, United Arabian Emirates.
because of the rarely dissolved sulfate, which is a hundred times lower than in sea water. Microscopic spheroidal clusters of interlocking crystals, termed framboidal pyrite because of the resemblance to raspberries, are known from shelf carbonates and deep-marine limestones (Fabricius 1961; Honjo et al. 1965). They have been explained as petrified sulfur bacteria; these are crystal aggregates whose formation was triggered by bacterially controlled processes, or as inorganic crystallization out of iron sulfide gels. Framboidal pyrite occurs scattered or as rounded clusters within fine-grained micritic limestones (Pl. 109/1) and marls and residual clays (Fig. 13.3) as well as in the form of pyrite fillings within fossils (e.g. radiolarians) and synsedimentary cavities. Early diagenetic geopetal fillings are potential tools for recognizing top and bottom structures and tectonic overturning. Reworked pyrite (nodules, tubes and fossil steinkerns) is believed to occur widely through the Phanerozoic, where currents eroded muddy shelves or slopes in anaerobic to minimally dysaerobic settings and exhumed pyrite grains from the underlying beds (Baird and Brett 1986). 13.1.2.3 Sulfates: Evaporite Minerals Authigenic evaporite minerals are common in limestones and dolomites. Evaporites are often interbedded with dolomite, limestone and fine-grained siliciclastic shales, e.g. thin red beds. Gypsum crystals can be identified by characteristic lath-like crystal shapes, weak birefringence and low reliefs. Anhydrite is distinguished from gypsum by a higher relief and stronger birefringence. For thin sections of evaporites see Garrison et al. (1978), Scholle (1979) and Adams et al. (1984). Evaporites are sedimentary deposits composed of minerals precipitated from brines concentrated by extensive or total evaporation of water. Evaporites form
Depositional settings of evaporites Evaporites originate in a number of environments: • Continental (subaerial and subaqueous) settings are characterized by rapid variations in basinal flooding and mineralogically and geochemically enormously varied deposits (e.g. playa deposits) originating from the waters of inland desert basins. • Supratidal (marine sabkha). These evaporite deposits form by evaporation in marginal-marine arid environments (e.g. Persian Gulf and Red Sea). They are characterized by either/or both siliciclastic and marine carbonate matrices and affected by meteoric and subaerial conditions. Distinct facies belts from seaward to landward or from the base of the sedimentary sequence upward are common (Alsharhan and Kendall 1994). A normal sabkha deposit consists of marine microbially controlled laminated carbonate or siliciclastic muds, nodular sulfates, desiccation crusts and windblown sands. The sulfates are usually gypsum or anhydrite and occur in microbial beds as well as in beds characterized by nodular or enterolithic masses. The finegrained carbonates are commonly dolomitized. Well preserved gypsum and anhydrite can be distinguished in thin sections by their crystal morphology (Schreiber et al. 1982). • Intertidal to shallow subaqueous settings (hypersaline sea). Evaporites occur as isolated crystals within microbial mats, or build gypsum beds. The morphology of the gypsum crystals and the mats is commonly well preserved as described from ancient examples. • Deep subaqueous settings below the wave base and photic zone. These evaporites are characterized by a long-distance continuity of laminar beds consisting of gypsum, anhydrite or halite, sometimes by co-occurrence with turbidites. Marine evaporites originate from ocean waters in barred basins, shallow epeiric seas, or in deep-marine settings. For case studies of ancient evaporites see Handford et al. (1982) and Rouchy et al. (2001). Criteria Primary sedimentary structures of evaporites include lamination, cross bedding, graded bedding, ripple marks and mud cracks. A laminated anhydrite or gypsum se-
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Phosphates
observed on the scale of outcrops and cores as well as in microfacies samples. The studies must consider the lateral and vertical development of depositional patterns, depositional cycles, e.g. shallowing-upward sequences, and the rapid changes in physico-chemical environmental controls. Common small-scale microfacies include stratification, occurrence of massive sulfates with gypsum crystals, intercalation of thin marine carbonates, intervals with reworked gypsum grains, and claystones with nodular gypsum features (Aigner and Bachmann 1989; Balzer 1997).
Fig. 13.4. Calcite pseudomorphs after idiomorphic euhedral evaporite crystals formed in a near-coastal hypersaline lagoonal environment. Dark micritic grains correspond partly to black pebbles. Late Jurassic (Kimmeridgian): Immouzzer, western High Atlas, Morocco. After Hüssner (1985).
quence consists of alternating laminae of white calcium sulfate mineral and gray lamellae rich in dolomite and organic carbon. Another common fabric is the alternating occurrence of dolomite layers and layers of gypsum crystals. Diagenetic structures are abundant due to the extreme solubility of evaporite minerals. A common diagenetic structure is chicken-wire anhydrite consisting of elongated white anhydrite in a dark microcrystalline anhydrite matrix. Nodular gypsum or anhydrite structures are common in sabkha environments, but should be used with caution for environmental interpretations. Enterolithic folding is a structure characterized by an intralayer crumbling (Pl. 125/4). Evaporite-solution breccias have a well-defined base and may extend upward through many beds of different lithology (Sect. 5.3.3.3). Most primary and diagenetic structures, e.g. lamination patterns, chicken-wire and nodular structures as well as the replacement of anhydrite after gypsum or calcite after evaporite minerals (Fig. 13.4) can be studied in thin sections. Facies and microfacies studies of evaporites The different environments in which evaporite formation takes place are characterized by a great number of sedimentary structures that are controlled by the concentration of various ions, temperature and organic components present, and reflected by microfacies types. Each variation produces differences in the style and rates of sedimentation, no matter which evaporite minerals are present. Facies analyses of evaporite rocks and evaporite-carbonate sequences require a differentiation between primary facies characteristics and diagenetic overprints (Schreiber et al. 1982). Both features can be
Other thin-section criteria are: Depositional and diagenetic fabric • lamination types, • relation of layers (couplets, successive, unrelated, disrupted), • microfolds, • microbial relicts, • quiet-water ooids, peloids or siliciclastic grains enclosed within evaporite layers, • arenitic or ruditic resedimented gypsum. Crystal fabric • morphology and growing habits of gypsum crystals e.g. single, twins (swallow tail), stellate clusters (Pl. 125/2, or in massive beds. For classifications of evaporites based on structural and textural criteria see Maiklem et al. (1969), Langbein (1979) and Balzer (1997). 13.1.2.4 Phosphates Phosphatic sediments form in different environments – not only beneath upwelling systems, but also atop seamounts, guyots and plateaus, on oceanic islands, and on some shelves and in epeiric environments. The formation of phosphorites includes two basic processes: (1) Physicochemical or biochemical precipitation of carbonate fluorapatite within the sediment, at the sediment/water interface and/or during diagenesis. Microbial contribution is important (Krajewski et al. 1994). (2) Hydraulic and biological reworking of primary phosphorites by which phosphatic grains are concentrated to form granular phosphorite. Non-transported phosphorites may occur in any or all of the principal phosphorite settings; many of them are associated with stratigraphic condensation, depositional breaks and/or unconformities. Granular phosphorites developed on shelves and in epeiric seas and appear to dominate phosphorite occurrences in the Phanerozoic record (Glenn et al. 1994). Phosphatic grains include peloids, compound grains,
Phosphatization
oncoids, and intraclasts. Primary phosphatic bioclasts (vertebrate bones, fish scales and teeth) are common. Phosphatization of carbonates Phosphatic particles also form by replacement of carbonate substrates, skeletal and non-skeletal carbonate particles (Fig. 4.16) as well as microbial structures, e.g. microstromatolites. These particles are characterized by marginal or complete impregnation/replacement showing various degrees of preservation of the primary structures (Swett and Crowder 1982; Rodgers 1994). Extensive interparticle, intraparticle cementation and rim cementation by cryptocrystalline carbonate fluorapatite, starting from the sediment-water interface and penetrating pelagic carbonates up to a few centimeters takes place during interruptions in sedimentation. Intraparticle cements occur within skeletal pores and boring pores. Skeletal material may be totally transformed to structureless carbonate fluorapatite peloids. Growth of gradually thickened phosphate coatings (rim cement) around a single or, more commonly, several phosphatized skeletal grains or lithoclasts may produce large coated composite grains. Interparticle cements and grains appear brownish and isotropic between crossed nicols, due to the presence of organic matter and iron oxides (Pl. 110). Basics: Authigenic minerals in carbonate rocks Silicification of carbonates, authigenic quartz and feldspar DeMaster, D.J. (1981): The supply and accumulation of silica in the marine environment. – Geochimica et Cosmochimica Acta, 45, 1715-1732 Hesse, R. (1990a): Origin of chert: diagenesis of biogenic siliceous sediments. – In: McIllreath, I.A., Morrow, D.W. (eds.): Diagenesis. – Geoscience Canada, Reprint Series, 4, 227-252 Hesse, R. (1990b): Silica diagenesis: origin of inorganic and replacement cherts. – In: McIllreath, I.A., Morrow, D.W. (eds.): Diagenesis. – Geoscience Canada, Reprint Series, 4, 253-276 Kastner M., Siever, R. (1979): Low temperate feldspars in sedimentary rocks. – American Journal of Science, 279, 435-479 Laschet, C. (1984): On the origin of cherts. – Facies, 10, 257-289 Maliva, R., Siever, R. (1988): Mechanisms and controls of silicification of fossils in limestones. – Journal of Geology, 96, 387-398 Martín Penela, A.J. (1995): Silicification of carbonate clasts in a marine environment (Upper Miocene, Vera Basin, SE Spain). – Sedimentary Geology, 97, 21-32 Molenaar, N., de Jong, A.F.M. (1987): Authigenic quartz and albite in Devonian limestones: origin and significance. – Sedimentology, 34, 623-640 Wilson, R.G.C. (1966): Silica diagenesis in Upper Jurassic limestones of southern England. – Journal of Sedimentary Petrology, 36,1036-1049 Further reading: K064, K091, K092
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Glauconite Amorosi, A. (1993): Use of glauconites for stratigraphic correlation: a review and case studies. – Giornale di Geologìa, 55, 117-137 Chafetz, H.S., Reid, A. (2000): Syndepositional shallow-water precipitation of glauconite minerals. – Sedimentary Geology, 136, 29-42 Odin, G.S., Matter, A. (1981): De glauconiarum origine. – Sedimentology, 28, 611-641 Triplehorn, D.E. (1966): Morphology, internal structure, and origin of glauconite pellets. – Sedimentology, 6, 247-266 Further reading: K093 Pyrite Baird, G.C., Brett, C.E. (1986): Erosion of an anaerobic seafloor: significance of reworked pyrite deposits from the Devonian of New York State. – Palaeogeography, Palaeoclimatology, Palaeoecology, 57,157-193 Berner, R.A. (1985): Sulphate reduction, organic matter decomposition and pyrite formation. – Philosophical Transactions of the Royal Society of London, A315, 25-38 Canfield, D.R., Raiswell, R. (1991): Pyrite formation and fossil preservation. – In: Allison, P.A., Briggs, D.E.G. (eds.): Taphonomy: releasing the data locked in the fossil record. – 337-387, New York (Plenum) Hudson, J.D. (1982): Pyrite in ammonite-bearing shales from the Jurassic of England and Germany. – Sedimentology, 29, 639-667 Schallreuther, R. (1984): Framboidal pyrite in deep-sea sediments. – Initial Reports of the Deep Sea Drilling Project, 75, 875-891 Schieber, J. (2002): Sedimentary pyrite: a window into the microbial past. – Geology, 30, 531-534 Wilkin, R.T., Barnes, H.L., Brantley, S.L. (1996): The size distribution of framboidal pyrite in modern sediments: an indicator of redox conditions. – Geochimica et Cosmochimica Acta, 60, 3897-3912 Further reading: K067 Sulfates: Evaporites Aigner, T., Bachmann, G.H. (1989): Dynamic stratigraphy of an evaporite-to-red bed sequence, Gipskeuper (Triassic), southwest German Basin. – Sedimentary Geology, 62, 5-25 Balzer, D. (1997): Mikrofazies-Analyse von Ca-Sulfatgesteinen des Zechstein. – Geologisches Jahrbuch, D, 106, 3-99 Handford, C.R., Loucks, R.G., Davies, G.R. (1982): Depositional and diagenetic spectra of evaporites - a core workshop. – SEPM Core Workshop, 3, 395 pp. Renault, R.W. (ed., 1989): Sedimentology and diagenesis of evaporites. – Sedimentary Geology, 64,207Richter-Bernburg, G. (1985): Zechstein-Anhydrite. Fazies und Genese. – Geologisches Jahrbuch, A85, 1-82 Rouchy, J.M., Taberner, C., Peryt, D. (eds., 2001): Sedimentary and diagenetic transitions between carbonates and evaporites. – Sedimentary Geology, 140, 1-189 Sarg, J.F. (2001): The sequence stratigraphy, sedimentology and economic importance of evaporite-carbonate transitions. – Sedimentary Geology, 140, 9-34 Shearman, D.J. (1991): Evaporite sediments and rocks: the calcium sulphate and halite facies. – 304 pp., New York (Van Nostrand)
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Warren, J. (1999): Evaporites. Their evolution and economics. – 448 pp., Oxford (Blackwell) Further reading: K065 Phosphates and phosphorites Baturin, G.N. (1982): Stages of phosphorite formation. – Developments in Sedimentology, 33, 344 pp. Bentor, Y.K. (ed., 1980): Marine phosphorites. – Society of Economic Paleontologists and Mineralogists, Special Publications, 29, 259 pp. Föllmi, K.B. (1996): The phosphorus cycle, phosphogenesis and marine phosphate-rich deposits. – Earth Science Reviews, 40, 55-124 Glenn, C.R., Föllmi, K.B., Riggs, S.R., Baturin, G.N., Grimm, K.A., Trappe, J., Abed, A.M., Galli-Oliver, C., Garrison, R.E., Ilyin, A.V., Jahl, C., Roehrlich, V., Sadaqah, R.M.Y., Schidlowski, M., Sheldon, R.E., Siegmund, H. (1994): Phosphorus and phosphorites: sedimentology and environments of formation. – Eclogae geologicae Helvetiae, 87, 747-788
Phosphatization of Carbonate Rocks
Lucas, J., Prévost-Lucas, L. (2002): Phosphorite and limestone, two independent end-member products of the range of bio-productivity in shallow-marine environments. – In: Glenn, C.R., Prévost-Lucas, L., Lucas, J. (eds.): Marine authigenesis: from global to microbial. – SEPM, Special Publications, 66, 117-127 Nothold, A.J.G., Jarvis, I. (eds., 1990): Phosphorite research and development. – Geological Society of America, Special Publications, 52, 326 pp. Soudry, D. (2000): Microbial phosphate sediments. – In: Riding, R., Awramik, S.E. (eds.): Microbial sediments. – 127-136, Berlin (Springer) Trappe, J. (1998): Phanerozoic phosphorite depositional systems. A dynamic model for a sedimentary resource system. – Lecture Notes in Earth Sciences, 76, 316 pp. Trappe, J. (2001): A nomenclature system for granular phosphate rocks according to depositional texture. – Sedimentary Geology, 145, 135-150 Further reading: K066
Plate 110 Phosphatization of Carbonate Rocks Microfacies studies of phosphate rocks and phosphorites reveal striking textural similarities within carbonate rocks. Phosphate fabrics are the result of biologic and authigenic processes as well as replacement of carbonate particles and sediment. The term phosphate rock defines a rock composition with less than 50% phosphate components, and phosphorite those of more than 50% phosphate components. Most ancient phosphate deposits consist of granular phosphoritic rocks (Soudry and Nathan 1980). Paralleling the texture-based carbonate classification of Dunham (1962), orthochemical/microbial phosphorites, characterized by in situ phosphatic precipitates, and granular fabrics with allochemical phosphate grains can be distinguished (Trappe 2001). The grains are derived from: (a) the reworking of orthochemical/microbial phosphorites resulting in structureless phosphate grains, variously called pellets, peloids or ovoids. The spectrum of these grains also includes phosphatic ooids, oncoids, pisoids, and fecal pellets. (b) Epigenetic replacement of carbonate grains, mostly bioclasts and intraclasts; (c) phosphatic bioclasts, e.g. vertebrates and inarticulate brachiopods, some arthropods, and conodonts (see Fig. 4.9). Strong phosphatization of grains may be characteristic of lag deposits (–> 2, 3, 5); see the Standard Microfacies Type 14. Except for the difference in mineral composition, thin sections of phosphorites look very much like thin sections of limestones. The plate displays texture types and grain types of granular phosphorites and phosphatized carbonates. Rock and grain names follow the suggestions of Trappe (2001). 1 Phosclast grainstone. Silicified bone bed. Phosphatic skeletal grains (‘phosbioclasts’; predominantly fish bones and shark teeth, S), reworked and redeposited phosphate concretions (RPC) and black rounded lithoclasts (‘phosclasts’, PC) derived from the reworking of phosphatized hard grounds. Cretaceous/Tertiary (Maastrichtian-Ypres): Imi-n-Tanoute, Morocco. 2 Phosooid packstone with layers of fish bones (FB) and phosphatized lithoclasts and phosooids within a carbonate matrix. Note the open internal meshwork (arrow) and the solid exterior covering of the fish bones. Cretaceous/Tertiary (Maastrichtian-Ypres): Chichaoua, Morocco. 3 Phosclast rudstone. Large phosphatized lithoclasts (L) and phosphatic skeletal grains. Early Tertiary (Thersirea Formation, Lutetian): Chichaoua, Morocco. 4 Phosphatized ooids (OO) and small rounded skeletal grains. Same locality as –> 3. 5 Phosphatized lithoclasts. Grains exhibit complete or only marginal phosphatization. Early Carboniferous (Delle Phosphatic Member, Oseagean-Meramecian): Eureka Mining District, Mammoth Mountain, U.S.A. 6 Hardground with microbial phosphate crusts (arrows). The rock is a phoslitho/bioclastic packstone. Note the high compositional variability of grains comprising lithoclasts, ooids, quartz grains and skeletal grains including bryozoans (B), echinoids (E) and serpulids (S). Early Tertiary (Amniter Formation, Thanet-Ypres): Iris-n-Sebt, Morocco. –> 1–6: Courtesy of J. Trappe (Bonn)
Plate 110: Phosphatization of Carbonate Rocks
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13.2 Geochemical Proxies This section summarizes trace element analysis and the use of stable isotopes in evaluating paleoenvironments and the diagenetic development of carbonate rocks. The text focuses on the potential of combined geochemical and microfacies studies. See Brand and Veizer (1980) and Morse and Mackenzie (1990) for basic information on the geochemistry of sedimentary carbonates.
13.2.1 Trace Elements Trace elements in carbonate rocks occurring at concentrations between 10-2 and 10-3 wt.%, or less, are either bound to the carbonate phase studied in RFA analyses, or to the non-carbonate phase represented by acidresistent residues and measured by AAS and ICP (Bausch 1994). For a time trace elements of limestones were usually investigated in bulk samples, but are now studied in samples differentiating micritic matrix, carbonate cements, fossils and other grains using various microsampling techniques. Trace element studies concentrate on specific trace elements (e.g. strontium or manganese) or investigate a large set of elements followed by statistical analyses. Multivariate analyses are a fundamental prerequisite in testing possible relations between microfacies types and chemical data (e.g. Röhl and Strasser 1995; Kühl et al. 1996) In the 1970s many authors stated that rare element variations in carbonate rocks were controlled by variations in lithology, skeletal composition and diagenesis and that rare element patterns might reflect sedimentary facies patterns (Wedepohl 1970; Veizer and Demovic 1974). A more skeptical view is expressed in modern studies concluding that geochemical data can be used in facies diagnoses only when the lithologic and diagenetic development of the rock unit are well known (Oesterreich 1992). This view results partly from the difficulties relating geochemical data to microfacies types based on minor paleontological or textural differences. A critical combination of microfacies types and trace elements is successful if microfacies types are defined using the major compositional criteria (e.g. Lintnerova and Hladikova 1992; Masaryk et al. 1993).
13.2.2 Strontium and Manganese - Favorite Tools for Facies Studies Many studies deal with common elements such as Sr, Mn or Mg. A relatively new approach is the use of Rare Earth Elements. Mg, Sr and and sometimes Mn are
Trace Elements
linked to the carbonate phase, depending on the similarity of the crystal chemistry of Sr and Ca, and Mg and Mn and the substitution of CaCO3 by MgCO3 in calcite and aragonite. The major proportion of limestone studies using trace elements as facies indicators is concerned with these elements. Examples of the application of strontium and manganese distribution to the facies analyses of carbonates are listed in Box 13.2. Strontium Starting with the investigations by Veizer and Demovic (1973), who recognized an apparently facies-controlled bimodal distribution of Sr in Mesozoic micrites, strontium is used as a proxy for paleoenvironmental and diagenetic conditions of carbonate sediments. A common approach aims at recognizing significant correlations between specific strontium contents and carbonate rock types (expressed by microfacies or lithology) and/or depositional settings (e.g. shallow-marine versus basinal). Sr values decrease from modern sediments to ancient limestones and to dolomites. Modern shallowmarine aragonite is characterized by high Sr values (up to 10 000 ppm). Calcite has lower amounts of strontium. The incorporation of Sr in carbonates depends on the primary mineralogy, water temperature, salinity, vital effects, and average Ca/Sr content of seawater. Most ancient limestones have only a few hundred ppm Sr, mainly due to the loss of Sr during neomorphic processes, such as transformation of aragonite to calcite, repeated dissolution-precipitation processes and recrystallization causing a progressive diagenetic elimination of Sr over time. Diagenesis in an open system influenced by freshwater results in decreasing, low and homogeneous values. Manganese The incorporation of Mn into marine carbonates depends on the primary mineralogy, crystallographic conBox 13.2. Selected references dealing with combined facies and trace element analyses of carbonate rocks. Strontium: Al-Hashimi 1976; Auernheimer 1983; Bausch 1968; Cerny 1978; Cioni et al. 1973; H.W. Flügel and Wedepohl 1967; Gittinger 1969; Imreh et al. 1965; Kranz 1976; Krejci-Graf 1965; Jorgenson 1981, 1986; Llavador et al. 1983; Masaryk et al. 1993; Matheos et al. 1990; Oesterreich 1992; Prasada and Naqvi 1977; Sakae et al. 1981; Santos Garcia 1992; Strohmenger and Dozet 1991; Veizer and Demovic 1973, 1974; Veizer et al. 1971 Manganese: Bencini and Turi 1974; Corbin et al. 2000; Oesterreich 1992; Pingitore 1978; Santos Garcia 1992; Shanmugam and Benedict 1983; Shimmield and Price 1986; Turi et al. 1980
Significance of Trace Elements
trols, the concentration of manganese in seawater, Eh conditions, organic matter and microbial processes. The behavior and the vertical distribution of solid-phase and dissolved manganese and iron in modern marine sediments are well known for pelagic and hemipelagic deepsea sediments (e.g. Aller 1990) and for anoxic coastal muds of temperate shelf environments (references in Berner 1980 and Chester 1990). Geochemical data from tropical shelf sediments, particularly reef-derived carbonates and shallow-marine carbonate muds are less common (Alongi et al. 1992).
13.2.3 Significance of Trace Elements in Facies Studies of Carbonate Rocks A basic assumption in the use of trace elements for facies interpretation is that a relationship exists between the trace element content of a carbonate and the concentration of this element in the water where precipitation takes place. Trace elements are applied to a wide range of topics but many applications need further refinement. Depositional conditions • Differentiating depositional and post-depositional physico-chemical conditions. Paleoenvironmental factors • Recognizing paleosalinity. In neritic and coastal marine environments, the Sr and Mg contents of fossil carbonates are correlated, allowing these elements to be used as indicators for the influence of continental and marine waters (Renard 1986). Sr has been used as a salinity indicator in Triassic lagoonal-reef systems (Kranz 1976). The value of Na as a paleosalinity indicator (Masaryk et al. 1993) is somewhat problematic (Hardie 1987; Buggisch et al. 1994) but sodium is a possible index to the salinity of diagenetic solutions. • Recognizing paleowater depths (Shanmugam and Benedict 1983). • Differentiating low and high energy sediments, e.g. low values in micritic lagoonal carbonates; high values in grainstones (Strohmenger and Dozet 1991). Primary mineralogy • Recognizing the original mineralogical composition of micrites and microspar. Sr and Mg concentrations in micrites mimic the concentrations in the precursor aragonite or Mg-calcite (Wiggins 1986), because these elements are incorporated into calcite at equal rates during diagenesis. During aragonite dissolution and calcite precipitation Sr is excluded from aragonite and builds up in solution. Phanerozoic micrites consisting
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originally of aragonite are usually Sr-rich; those consisting of calcite are poor in Sr. • Recognizing the primary aragonite and calcite mineralogy of fossils (Veizer 1983; Dodd and Stanton 1990). Diagenesis • Tracing marine, meteoric and burial diagenesis (e.g. Al-Hashimi 1982). Commonly used elements are strontium, manganese and magnesium. • Recognizing meteoric diagenesis, typically with enrichment of Fe and Mn and depletion in Sr and Na (e.g. Pingitore 1978; Brand and Veizer 1980; Al-Aasm and Veizer 1982). • Checking possible diagenetic alterations of shells used for stable isotopic studies (Grossman et al. 1996). Sedimentation patterns • Recognizing the shallowing of basins, reflected by decreasing Sr contents (Lintnerova et al. 1988). • Recognizing breaks in sedimentation, discontinuities, hardgrounds (Marshall and Ashton 1980). • Distinguishing allochthonous input and autochthonous sedimentation input. Indication of the input of allochthonous sediment (versus autochthonous origin). Detrital input is indicated by elements related to silicate minerals (e.g. Si, Al, K, Fe). • Understanding bedded limestone sequences. Geochemical analysis of bedded limestones can be improved by sampling bed per bed, also analyzing the intercalated marls and evaluating trace elements against the background of overall non-carbonate contents (Bausch 1994). Two-component diagrams are useful in distinguishing clastic, semiclastic and diagenetic elements and correlating profiles. Reefs and carbonate platforms • Differentiating carbonate sediments formed in different parts of ancient reef complexes (Chester 1965; Davies 1972; Flügel and Flügel-Kahler 1963; Matheos and Morsch 1990; Oesterreich 1992). • Differentiating inner and outer platforms (e.g. Orehek and Ogorelec 1981). Stratigraphy • Assessing stratigraphic boundaries (Jorgensen 1981). • Establishing chemostratigraphical correlations for Mesozoic and Cenozoic pelagic carbonates based on regional (and global) trace-element stratigraphy using Sr and Mg contents (Renard 1986). Long- and short-term paleoceanographic fluctuations • The Sr content of pelagic carbonates varies with time (Renard 1986). Long-term variations are influ-
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enced by diagenesis, variation of the distribution coefficient and Sr/Ca ratio in seawater, e.g. caused by variations in hydrothermal activity. Short-term variations are controlled by sea-level fluctuations. Basics: Trace elements in carbonates Brand, U. and Veizer, J. (1980): Chemical diagenesis of a multicomponent carbonate system – I. Trace elements. – Journal of Sedimentary Petrology, 50, 1219-1236 Carpenter, S.J. and Lohmann, K.C. (1992): Sr/Mg ratios of modern marine calcite: empirical indicators of ocean chemistry and precipitation rate. – Geochimica et Cosmochimica Acta, 56, 1837-1849 Churnet, H.G., Misra, K.C. (1981): Genetic implications of the trace distribution pattern in the upper Knox carbonate rocks, Copper Ridge District, East Tennessee. – Sedimentary Geology, 30, 173-194 Kühl, A., Schreiber, A., Scheffler, E. (1996): Faziesanalyse mit multivariatstatistischen und geostatistischen Methoden des Dunklen Knotenkalkes von Wildenfels. – Freiberger Forschungshefte, C, 462, 212 pp. Morse, J.W., Mackenzie, F.T. (1990): Geochemistry of sedimentary carbonates. – 707 pp., Amsterdam (Elsevier) Popp. B., Anderson, F.T., Sandberg, P.A. (1986): Textural, elemental and isotopic variations among constituents in Middle Devonian limestones, North America. – Journal of Sedimentary Petrology, 56, 715-727 Renard, M. (1986): Pelagic carbonate chemostratigraphy (Sr, Mg, 18O, 13C). – Marine Micropaleontology, 16, 117-164 Veizer, J. (1983): Chemical diagenesis of carbonates: theory and application of trace element techniques. – In: Arthur, M.A., Anderson, T.F., Kaplan, I.R., Veizer, J., Land, L.S. (eds.): Stable isotopes in sedimentary geology. – Soc. Econom. Paleont. Miner., Short Course, 10, 3.1 - 3.100 Further reading: K154, K155, K156
13.2.4 Stable Isotopes Marine and non-marine carbonates have been used extensively in documenting changes in the ratios of isotopic species. Stable isotope analysis is one of the most powerful tools in deciphering the depositional and diagenetic history of carbonate rocks and fossils, and understanding environmental controls and global changes of climate and oceanography. The studies are based on oxygen isotopes (18O, 16O), carbon isotopes (13C, 12C) and strontium isotopes (87Sr, 86Sr; Elderfield 1987). Delta values (‰) describe the difference in isotope ratios between samples and standards. Standards are usually based on CaCO3 (PDB standard; Pedee Formation Belemnite) or H2O (SMOW standard; Standard Mean Ocean Water). Measurements are whole rock data or individual constituent data (crystals, grains) that require microsampling techniques. Basic information on the methods and use of stable isotope analyses can be found in Arthur et al. (1983), Faure (1986) and Hoefs (1997). The following notes
Stable Isotopes
emphasize some applications in the context of integrated facies analyses. Current stable isotope analyses of carbonate rocks aiming for paleoenvironmental or diagenetic interpretations are based on the investigation of micrites, carbonate cements or fossils. Stable isotopes in fossils (predominantly foraminifera, mollusks and corals) are used in reconstructing growth rates of organisms (Wefer and Berger 1991), paleoclimatic fluctuations and paleooceanographic circulation patterns. Isotope signals recorded in planktonic foraminifera offer excellent possibilities for recognizing climate cycles, paleoproductivity and changes in ocean circulation. Carbonate oxygen isotope paleothermometry offers a possibility to evaluate paleo-seawater temperatures. The fractionation of oxygen isotopes is temperature dependent. The ratio of the two most common isotopes (16O and 18O) in carbonate fossils can, in principle, be used to reconstruct the temperatures of ancient oceans. Cenozoic and Mesozoic foraminifera are commonly analyzed in order to determine sea-surface, deep-water and bottom-water temperatures but mollusk and brachiopod shells and carbonate cements are used as well (see Corfield 1995 for discussion). Over the past forty years, the 16O/18O record in micro-organisms preserved in deep-sea sediments has provided basic data in establishing a history of long-term oceanic temperature and global ice volume during the Quaternary. Many investigations are concentrated on the Cenozoic, but increasingly more studies are deciphering the isotopic history of Mesozoic and Paleozoic oceans and carbonates. δ13C shifts during the Phanerozoic and excursions at or near boundaries are believed to represent changes in the global carbon cycle. Phanerozoic isotope carbon curves are used for correlations identifying climate variations, determining the isotopic composition of seawater, and recognizing secular variations of organic carbon and carbonate organic production/burial rates. Cenozoic and ancient (particularly Paleozoic) isotope patterns exhibit considerable differences, possibly due to primary differences in the isotopic composition, e.g. caused by exchange between sea water and oceanic ridges; differences in water temperatures; different biological fractionation of ancient organisms as compared with modern organisms; changes in productivity, and diagenetic alterations. Well-preserved brachiopods (Wadleigh and Veizer 1992; Wenzel 1997; Samtleben et al. 2000) and carbonate cements (Lohmann 1985; Carpenter et al. 1991) are assumed to be the most suitable phases for deci-
Use of Stable Isotopes
phering the original oxygen and carbon isotope signatures of carbonate rocks. The value of stable isotope data in paleoenvironmental and facies interpretation of ancient carbonates and carbonate skeletons depends strongly on the critical assessment of possible diagenetic alterations of the material investigated (Fig. 3.9). Diagenetic alterations of shells may be identified by investigating trace elements, cathodoluminescence and shell microstructures (critical remarks in Barbin and Gaspard 1995). Early meteoric diagenesis and burial diagenesis can greatly alter the original isotope ratios incorporated in skeletons or carbonate rocks (Censi and Montana 1994). Diagenetic alterations of micrites may be reflected by strontium and carbonate contents as well as by SEM and cathodoluminescence data. The use of stable isotopes in facies analyses of ancient carbonates Stable isotope studies have been successfully applied to Phanerozoic carbonates of all ages and to carbonates formed on ramps (Joachimski 1991) and platforms (Weis and Preat 1994; Rasser and Fenninger 2002), reefs (Aharon 1991), slopes and basins (Jenkyns and Clayton 1986; Masaryk et al. 1993). Measurements cover whole rock samples, the micritic matrix (aiming at differentiating allomicrites and automicrites), inter- and intraparticle cements (aiming at differentiating shallow-marine, burial and meteoric conditions), as well as individual grains (e.g. oncoids, ooids) and well-preserved shells. A common tool of these studies is the search for isotopic signals that would indicate changes from marine to meteoric conditions and subaerial exposure of the depositional environment, or depositional breaks. The basic prerequisite for the successful application of stable isotopes to facies analyses of limestones is the combination of the geochemical data with field observations, paleontological features and microfacies criteria (Samtleben 2000; Munnecke et al. 2003). Basics: Stable isotopes Arthur, M.A., Anderson, T.F., Kaplan, I.R., Veizer, J., Land, L.S. (1983): Stable isotopes in sedimentary geology. – Soc. Econom. Paleont. Miner., Short Course, 10, 1-151 Banner, J.J. (1995): Application of the trace element and isotope geochemistry of strontium to studies of carbonate diagenesis. – Sedimentology, 42, 805-824 Barrera, E., Tevesz, M.J.S. (1990): Oxygen and carbon isotopes: utility for environmental interpretation of recent and fossil invertebrate skeletons. – In: Carter, J.G. (ed.): Skeletal biomineralization: patterns, processes and evolutionary trends. Volume I. – 557-566, New York (Van Nostrand) Berger, H., Vincent, E. (1986): Deep-sea carbonates: Reading the carbon-isotope signal. – Geologische Rundschau, 75, 249-269
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Brand, U., Veizer, J. (1981): Chemical diagenesis of a multicomponent carbonate system. - 2: Stable isotopes. – Journal of Sedimentary Petrology, 51, 987-997 Clauer, N., Chaudhuri,S. (eds., 1992): Isotopic signatures and sedimentary records. – Lecture Notes in Earth Sciences, 43, 529 pp. Elderfield, H. (1986): Strontium isotope stratigraphy. – Palaeogeography, Palaeoclimatology, Palaeoecology, 57, 71-90 Faure, G. (1986): Principles of isotope geology. – 589 pp., New York (Wiley) Hoefs, J. (1997): Stable isotope geochemistry. 4th edition. – 201 pp., Berlin (Springer) Hudson, J.D., Anderson, T.F. (1989): Ocean temperatures and isotopic compositions through time. – Transactions of the Royal Society Edinburgh, 80, 183-192 Joachimski,M. (1991): Stabile Isotopen (C, O) und Geochemie der Purbeck-Mikrite in Abhängigkeit von Fazies und Diagenese (Berriasian/Schweizer und Französischer Jura, Südengland). – Erlanger Geologische Abhandlungen, 119, 1-114 Marshall, J.D. (1992): Climatic and oceanographic signals from the carbonate rock record and their preservation. – Geological Magazine, 129, 143-160 Morse, J.W., Mackenzie, F.T. (1990): Geochemistry of sedimentary carbonates. – Developments in Sedimentology, 48, 707 pp., Amsterdam (Elsevier) Munnecke, A., Samtleben, C., Bickert, T. (2003): The Ireviken event in the lower Silurian of Gotland, Sweden - relations to similar Palaeozoic and Proterozoic events. – Palaeogeography, Palaeoclimatology, Palaeoecology, 195, 99-124 Popp, B.N., Anderson, T.F., Sandberg, P.A. (1986): Brachiopods as indicators of original isotopic compositions in some Palaeozoic limestones. – Geological Society of America, Bulletin, 97, 1262-1269 Rasser, M., Fenninger, A. (2002): Paleoenvironmental and diagenetic implications of δ18O and δ13C isotope ratios from the Upper Jurassic Plassen limestone (Northern Calcareous Alps, Austria). – Geobios, 35, 41-49 Schidlowski, M. (2000): Carbon isotopes and microbial sediments. – In: Riding, R., Awramik, S.M. (eds.): Microbial sediments. – 84-95, Berlin (Springer) Schulz, H.D., Zabel, M. (eds., 2000): Marine geochemistry. – 455 pp., Berlin (Springer) Veizer, J., Hoefs, J. (1976): The nature of O18/O16 and C13/ C12 secular trends in sedimentary carbonate rocks. – Geochimica et Cosmochimica Acta, 40, 1387-1395 Veizer, J., Fritz, P., Jones, B. (1986): Geochemistry of brachiopods: Oxygen and carbon isotope records of Paleozoic oceans. – Geochimica et Cosmochimica Acta, 50, 1679-1696 Wadleigh, M.A., Veizer, J. (1992): 18O/16O and 13C/12C in lower Paleozoic articulate brachiopods: Implications for the isotopic composition of seawater. – Geochimica et Cosmochimica Acta, 56, 431-443 Wefer, G., Berger, W.H. (1991): Isotope paleontology: growth and composition of extant calcareous species. – Marine Geology, 100, 207-248 Wenzel, B,C. (1997): Isotopenstratigrapische Untersuchungen an silurischen Abfolgen und deren paläoozeanographische Interpretation. – Erlanger Geologische Abhandlungen, 129, 1-117 (cum lit. !) Further reading: K157, K158 (stable isotopes in sediments), K159 (stable isotopes in organisms and fossils)
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13.3 Organic Matter in Carbonate Rocks The formation of carbonate rocks can be considered as a part of a three-component system (Ricken 1993). Organic matter, carbonate and siliciclastic sediment may settle concurrently, but at different rates, and intermix at the sea floor resulting in a particular sediment composition. A change in the flux of one component can cause diluting or concentrating of the other two components expressed by the ratio of organic carbon to carbonate. Organic matter is a trace constituent of calcareous skeletons (Crick 1989), carbonate cements (Dravies and Yurewicz 1985), carbonate particles (Davies et al. 1978) and speleothems (Shopov et al. 1994). The presence and composition of organic matter in carbonates is used as a tool in • estimating the productivity and the origin of ‘black shales’. • recognizing depositional facies and delineate diagenetic fabrics, • understanding the controlling effect of organic constituents on mineral growth, • providing fingerprints for the source of organic matter in sediments and the origin of source rocks, and • determining the timing of hydrocarbon migration. The molecular structure of organic matter controls mineral growth (Ramseyer et al. 1997). Inorganically precipitated carbonates (speleothems, cements; Dravies and Yurewicz 1985) incorporate during growth adsorbed organic matter between submicroscopic subunits of crystals. Biologically secreted carbonates incorporate biogenic tissue between crystals. The organic matter is derived from soils (speleothem calcite), dissolved organic matter (marine carbonates) and biological decay products (aragonitic coral skeletons). Biogeochemical proxies: Sedimentary compounds within sedimentary rocks supply important biomarkers and paleoenvironmental indicators (de Leeuw et al. 1995). Marine as well as non-marine environments are characterized by the combined presence of specific compounds. Biomarkers are (a) proxies for various groups of bacteria, cyanobacteria, algae, vascular higher plants, and animals (Peters and Moldovan 1993); (b) assist in estimating shoreline proximity of modern carbonates, and allow paralic margins dominated by mangroves, and subtidal mud-banks characterized by sea grass to be distinguished; and (c) are fundamental for evaluating of organic-rich carbonate rocks (Hopf et al. 2001). Paleoenvironmental characteristics, such as photic zone,
Organic Matter
anoxia, hypersalinity, microbial sulfate-reducing activity, algal blooming, paleo-upwelling and paleo-surface seawater temperature are indicated by molecular and/ or isotopic characteristics. Most work on organic petrology and palynofacies concerns siliciclastic rocks. But biomarkers are also promising and sensitive paleoenvironmental and depositional indicators in carbonate rocks as demonstrated by the distinct differences in autochthonous biomarkers in reef, platform and basin carbonates allowing facies interpretations to be refined with respect to salinity, oxygenation and water stratification (Marynowski et al. 2000). Cyclicity and its bearing on sequence stratigraphy is another tool that can be elucidated by organic matter distribution in limestones (Bombardiere and Gorin 2000). The identification of silt-sized remains of organisms may be facilitated by biomarkers as well, e.g. the differentiation of grains derived from calcareous algae and from corals (Böhm et al. 1980). Basics: Organic matter Crick, R.E. (ed., 1989): Origin, evolution and modern aspects of biomineralization in plants and animals. – 536 pp., New York (Plenum Press) De Leeuw, J.W., Frewin, N.L., Van Bergen, P.F., Sinninghe Damste, J.S., Collinson, M.E. (1995): Organic carbon as a paleoenvironmental indicator in the marine realm. – In: Bosence, D.W., Allison, P.A. (eds.): Marine paleoenvironmental analysis from fossils. – Geological Society of London, Special Publication, 83, 43-71 Gautier, D.L. (ed., 1986): Roles of organic matter in sediment diagenesis. – Soc. Econ. Paleont. Miner., Special Publications, 38, 203 pp. Huc, A.Y. (1990): Deposition of organic facies. – American Association of Petroleum Geologists, Studies in Geology, 30, 234 pp. Marynowski, L., Narkiewicz, M., Grelowski, C. (2000): Biomarkers as environmental indicators in a carbonate complex, example from the Middle to Upper Devonian, Holy Cross Mountains, Poland. – Sedimentary Geology, 137, 187-212 Peters, K.E., Moldovan, J.M. (1993): The biomarker guide interpreting molecular fossils in petroleum and ancient sediments. – 363 pp., Englewood Cliffs (Prentice Hall) Ramseyer, K., Miano, T.M., D’Orazio, V., Wildberger, A., Wagner, T., Geister, J. (1997): Nature and origin of organic matter in carbonates from speleothems, marine cements and coral skeletons. – Organic Geochemistry, 26, 361-378 Ricken, W. (1993): Sedimentation as a three-component system. – Lecture Notes in Earth Science, 51, 211 pp., Berlin (Springer) Tyson, R.V. (1995): Sedimentary organic matter. Organic facies and palynofacies. – 615 pp., London (Chapman and Hall) Whelan, J.K., Farrington, J.W. (eds., 1992): Organic matter: productivity, accumulation and preservation in recent and ancient sediments. – 533 pp., New York (Columbia Press) Further reading: K069
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14 Depositional Models, Facies Zones and Standard Microfacies
Preceding chapters of part 2 of this book were focused on defining microfacies types, recognizing paleoenvironmental constraints and using an integrated approach to facies analysis. Chapter 14 introduces carbonate facies models, underlines the importance of environmentcontrolled distribution patterns of organisms, and gives an overview of Standard Microfacies Types and their application to the facies analysis of limestones. The latter topic is expanded in the case studies described in Chap. 15. • A depositional facies model is the essence of the criteria and controls extracted from case studies of modern and ancient sediments. • Facies zones (FZ) are belts differentiated according to the changes of sedimentological and biological criteria across shelf-slope-basin transects. • Standard microfacies types (SMF) are derived from local microfacies types looking at joint paleontological and/or sedimentological criteria. Associations of SMFs are additional criteria for recognizing ancient facies zones.
14.1 Depositional Facies Models A facies model is a generalized summary of a given depositional system (Walker 1992). The information required for this summary is taken from local case studies of modern and ancient examples. Facies models should act as norms for purposes of comparison, as a framework and guideline for future observations, as a predictor of new geological situations, and as an integrated basis for the system that it represents. Sensitive and predictive facies models are important in the exploration of hydrocarbons and limestone resources (see Chap. 17 and 18).
14.1.1 Conceptual, Dynamic and Computer Models Carbonate facies models correspond to different categories and comprise conceptual, dynamic and numerical models.
14.1.1.1 Conceptual Models The first facies models proposed for carbonate rocks were conceptual models trying to show spatial distribution patterns of rock types and biota, usually along a coast-basin traverse. These static models were strongly influenced by the investigation of modern subtropical and tropical carbonates. For a long time the Bahamas, Florida and the Persian Gulf served as paradigms for shallow-marine carbonate platforms. Starting with the ‘X-Y-Z model’ by Irwin (1965) new ideas were introduced into carbonate models that were supposed to reflect the overall environmental controls on carbonate deposition on shelves and include the differences between ancient depositional settings and modern settings. Using the commonly observed successions of facies belts of Holocene and Phanerozoic carbonates, Wilson (1975) established a Standard Facies Model that describes Standard Facies Zones (FZ) of a rimmed tropical carbonate platform along a strongly generalized shore-to-basin transect. This Wilson model (see Sect. 14.1.3) had a strong impact not only on the interpretation of the litho- and biofacies of carbonate rocks but also on the use of microfacies types as indicators of facies zones and depositional conditions. Extending the generalizations made by Flügel (1972) for common microfacies types of Triassic carbonates, Wilson introduced a set of ‘Standard Microfacies Types’ (SMF) which aimed at categorizing common and widely distributed Phanerozoic microfacies types. About the same time, Ahr (1973) recognized the difficulties related to the application of the rimmed platform concept to the facies distribution of many Paleozoic and Mesozoic shelf carbonates and proposed the carbonate ramp model. This model summarizes the facies distribution on a gently sloping surface from the coast to deeper water without a morphological break at the shelf edge. Ramps differ from rimmed shelves in the absence of continuous shelf-marginal reef trends, the location of high-energy deposits near the shoreline and not at the shelf edge, and the lack of shallow-water derived clasts in deep-water parts of the ramp. Ginsburg and James (1974), Kendall and Schlager (1981), Schlager (1981, 1992, 2000), Read (1982,
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_14, © Springer-Verlag Berlin Heidelberg 2010
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Numerical Models
1985), Tucker (1985), Tucker and Wright (1990), James and Kendall (1992), Handford and Loucks (1993), Sanantonio (1993), Read et al. (1995) and Lukasik et al. (2000) provided enlarged and more differentiated models of carbonate depositional environments. These papers discuss the subdivision and internal subdivision of ramps and emphasize the importance of non-rimmed shelves, drowned and pelagic platforms, as well as epeiric platforms and ramps. Some of these models are discussed in Sections 14.1.4 to 14.1.8.
14.1.1.2 Dynamic Models Depositional systems are inherently complex, yet ancient successions are often reduced to very simple models considering just one major control factor, most commonly sea level. These dynamic models stress the changes in depositional systems and environmental controls through time and combine different scales of observations from microfacies to sedimentary sequences (Aigner 1984). Dynamic models have their merits because they try to avoid strict categorizations such as those involved in the Wilson model and because dynamic models describe the development of facies zones during time. Systems tracts deciphered by sequence stratigraphy are the result of gradual changes in marine environments caused by variations in water depths related to sea-level fluctuations, accommodation space and biogenic production. These changes are characterized by shifts in the relative importance of limestone grain types. The shifts reflect temporal continua or relays that can be derived from optimized similarity matrices (see Sect. 14.4).
14.1.1.3
Numerical Models
A third category of facies models are computer models conditioned with appropriate geometric and stratigraphic data from outcrops or wells and process data from modern carbonate depositional environments. These models simulate temporal development, stratal geometries and the architecture of carbonate buildups, and evaluate the rates of processes responsible for the formation of carbonate platforms and slopes. Computer models provide a visual image of the internal architecture of carbonate platforms (e.g. Bosence and Waltham 1990) and platform-slope geometries and the shape of slopes (Aigner et al. 1989; Bosscher and Southam 1992). Modeling allows quantification and evaluation of the major controls on the evolution of platforms and ramps studied in outcrops and derived from seismic
sections. The value of the models is usually tested by comparison with well-studied outcrops (e.g. Harris 1991). Most computer models are based on algorithms that simulate processes acting during the development of carbonate buildups and determining their shape and internal structure. Simulations are performed on a spatial grid in a sequence of time steps from a prescribed set of initial conditions defined by various input parameters. During program application, these parameters are varied within the limits of known geological and sedimentological constraints until the model is consistent with all available data and observations. The programs developed for modeling carbonate and/or carbonate-siliciclastic depositional systems differ in their main goals, input parameters used, and the way modeling results are shown. Carbonate computer models are two-dimensional or three-dimensional. Principal key input parameters of 2-D modeling (see Enos 1991) are sea-level fluctuations, subsidence, clastic sediment input and depth-related carbonate growth potential. Other parameters are sedimentation and erosion rates (Bosence et al. 1994), platform shape and slope angle, and early diagenetic processes and compaction (e.g. Strobel et al. 1990). Model outputs are 2-D cross sections explaining • internal and external platform geometries (Bosence and Waltham 1990; Read et al. 1991), • the growth and geometry of reef structures (Gerhard 1991; Bosscher and Schlager 1992; Hüssner 1997), • the evolution and geometry of shelf-basin transitions (Aigner et al. 1990; Kendall et al. 1991; Bosence et al. 1994), • the development of carbonate slopes (Bosscher and Southam 1992), • the estimation of paleobathymetric conditions (Aurell et al. 1995), • the sea-level control on cyclic sedimentation (Goldhammer et al. 1991), • the evaluation of depositional models based on sequence stratigraphic concepts (Burgess 2001). Problems involved in computer models are the difficulty in separating global sea-level signals from local or regional processes, obtaining a high time resolution, and quantifying the controls on the biogenic carbonate production potential. In addition, many models provide only rough lithologies or strongly generalized lithofacies types. Thin-section data on microfacies scale are important parameters in modeling component cycles (see papers in Franseen et al. 1991).
Common Facies Belts
Current 3-D modeling incorporates the basic processes of carbonate accumulation and diagenesis in relation to sea-floor fluctuations, basement subsidence, wave-base erosion and transport of sediment. Common input parameters are initial bathymetry and sea-floor geometry, sea-level changes, changes in paleoenvironmental conditions (e.g. salinity, temperature), and paleoclimate. 3-D modeling of carbonate basins has become a very successful technique in reservoir modeling, particularly where the geological and physical reservoir parameters show extreme heterogeneities (e.g. Vahrenkamp and Grötsch 1994). Modeling of the relationships between reservoir lithofacies and petrophysical criteria reflected by logs requires the differentiation of well-strained lithofacies units. The consideration of microfacies types is an important tool in defining these units.
14.1.2 Basic Elements of Carbonate Facies Models Basic elements used in carbonate facies models describing transects from shallow-marine to deep-marine settings are • changes in the relief of the sea bottom expressed by differently dipping shelf slopes and the shallow-water or deeper-water position of a distinct break in the morphology of the shelf, • vertical boundaries affecting the sea bottom, and represented by high and low tide levels, the fair-weather wave base and the storm wave base, and • lateral differences in the composition of sediments and benthic biota allowing facies zones to be distinguished. These basic elements have been discussed in Chap. 2 of this book.
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with supratidal microbial laminites, sometimes with inland freshwater marsh deposits, coal, or siliciclastics. • Lagoonal facies. Developed behind barrier complexes. Mainly bedded limestones with micritic or peloidal lime mudstones, or cherty burrowed skeletal packstones to mudstones, sometimes with local patch reefs of colonial metazoans. Minor, thin interbeds of fenestral or cryptalgal carbonates reflect periods of shallowing of the lagoon to tide level. • Shoal-water sediments consisting of banks, reefs and ooid/pellet shoals. Occurring as shallow-ramp skeletal banks or lime-sand shoals, or as shelf-edge skeletal reefs and lime sands that may pass gradually downslope into deep ramp facies. On deeply sloping platform edges, the facies passes downslope into foreslope and slope deposits adjacent to deep shelves or basins. • Deep shelf and ramp facies. Nodular, sometimes cherty bedded skeletal packstone and wackestone with abundant whole fossils and diverse open-marine biota. The limestones may have upward-fining and stormgenerated beds. Water depths range from some tens to more than 100 m. The depositional base is largely below the fair-weather wave base, but may be influenced by storms. • Slope and basin facies. Foreslope and slope deposits adjacent to steeply sloping platforms have abundant breccias and turbidites interbedded with periplatform lime muds and terrigenous muds. Adjacent to most ramps, slope and basin deposits are thin-bedded periplatform lime muds or terrigenous muds with generally only few sediment-gravity flow deposits. Paleozoic basinal deposits are commonly shales, with carbonate content increasing toward the platform. Mesozoic and Cenozoic basinal deposits may be pelagic limestones or shales. The slope and basin floor can be anoxic and lacking benthic organisms. These deposits are laminated and non-burrowed. Oxic slope and basin deposits are burrowed and fossiliferous.
14.1.2.1 Common Facies Belts 14.1.2.2 Common Depositional Patterns Most models used in interpreting ancient shallowmarine to deep-marine carbonate rocks consider five major facies belts along a shoreline-to-basin transect that include platforms or ramps and continue via a slope into basins (Read 1995): • Tidal flat facies. Generally arranged in cyclic, shallowing-upward units. Sequences in arid climate exhibit burrowed to non-burrowed lagoonal limestones overlain by intertidal microbial and algal laminites, supratidal evaporites, or eolian-fluvial clastics. Humid sequences are burrowed subtidal-intertidal limestones
Sedimentation on shelves often results in meter-scale shallowing-upward sequences, because the accumulation rates commonly outspace the combined rates of subsidence and sea-level rise (James 1984; Osleger and Read 1991). Consequently, the tops of these sequences show signs of intertidal to supratidal exposure, or they are eroded. Such small-scale cycles can be stacked to form larger sequences displaying transgressive and regressive trends. Stacking of sequences is a characteristic criterium of many ancient platforms. Larger-scaled sequences are commonly attributed to third-order sea-
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level fluctuations, small-scale sequences to sea-level fluctuations in the Milankovitch band (e.g. Read and Goldhammer 1988; Goldhammer and Harris 1989; Strasser 1994). However, autocyclic processes such as migration of bars and islands, delta-switching and local progradation may also create shallowing-upward sequences in shallow-marine environments (Strasser 1991; Goldhammer et al. 1993). These processes are local and can not be followed over large lateral distances.
14.1.2.3 Different Depositional Settings Require Different Facies Models The basic differences of carbonate depositional settings were discussed in Chap. 2 of this book. Fig. 2.6 gives a brief survey of some of these settings. The discussion in the present chapter is focused on facies zones and models forming the base for the interpretation of ancient carbonate environments. The discussion starts with facies zones occurring on rimmed shelves (platforms) and adjacent slopes and basins (Sect. 14.1.3). This Wilson model is contrasted with the ramp model (Sect. 14.1.4) describing sedimentation on a gently dipping slope, and with the non-rimmed platform model (Sect. 14.1.5). The specific conditions of isolated platforms and atolls are examined in Sect. 14.1.6. The sedimentation on extremely extended epeiric platforms and epeiric ramps is the topic of Sects. 14.1.7 and 14.1.8. –> Note that these facies models are limited with regard to their predictive capacity owing to their static view of time and of the relative sea-level changes that control facies shifts. These difficulties can be partly overcome by models based on sequence stratigraphy demonstrating how depositional sequences respond to lowstand, transgressive and highstand conditions of relative sea level under humid and arid conditions (Handford and Loucks 1993). Sequence stratigraphy models can not replace the static models which have their own merits, but they complete our understanding of facies development through time (see Sect. 16.1.2).
14.1.3 Facies Zones of Rimmed Carbonate Platforms: The Wilson Model The most frequently used facies models are those hypothesizing on platforms and ramps. Carbonate platforms are dynamic systems that change through time
Standard Facies Zones
and space. Platforms may expand as their margins grow outward; grow upward while their margins remain stationary; or retreat at their margins backward (e.g. Jansa 1981; Blendinger 1986). Platforms are born, grow and die. Growth is caused by aggradation and/or progradation. Demise is connected with the decrease in and cessation of carbonate production due to (1) drowning caused by rapid eustatic sea-level rise or tectonic subsidence, (2) subaerial exposure caused by sea-level fall or tectonic uplift, (3) high siliciclastic input, or (4) paleooceanographic effects causing changes in water circulation, temperature and salinity. Major variables influencing the evolution of platforms are tectonic setting and subsidence, sea-level fluctuations, carbonate productivity and sediment transport, the nature of the sedimentation at platform margins, the evolution of reef-building organisms through time, and variations in diagenetic processes. Many platforms are ‘rimmed’ at their platform-ocean margins (e.g. South Florida Shelf, the Belize Shelf, or the Queensland Shelf with the Great Barrier Reef) or they are ‘non-rimmed’ (e.g. Western Florida Shelf). Both rimmed and non-rimmed platforms may be attached to the continent or occur detached from the continent as isolated platforms and atolls. Rimmed carbonate shelves and platforms are characterized by • a nearly flat platform top, • a landward low-energy lagoon, • an outer wave-agitated depositional, bypass or erosional margin, marked by • a continuous to semicontinuous reef-dominated or sand-shoal dominated rim along the margin that restricts circulation and wave action, • a pronounced increase in slope, ranging from a few degrees to 60° and more, and • abundant mass-flow deposits (e.g. debris flows, megabreccias, turbidites, slumps) downslope from the shelf edge containing platform-derived material.
14.1.3.1 Standard Facies Zones and the Modified Wilson Model The succession of major facies belts on rimmed tropical carbonate platforms was used by Wilson (1975) to establish a Standard Facies Model depicted as a basinto-shore transect and comprising Standard Facies Zones (Fig. 14.1). The basis of the model is the recognition of consistently recurrent patterns of carbonate facies in the Phanerozoic record and the environmental interpretation of these patterns by using characteristics of Holocene sedimentation patterns.
Standard Facies Zones
The Standard Facies Zones (FZ) defined in Box 14.1 describe idealized facies belts along an abstract transect from open-marine deep basins across a slope, a pronounced platform marginal rim (characterized by reefs or/and a zone with sand shoals), and an inner platform to the coast. The Facies Zones differ in setting, dominant sediments and prevailing biota, and common lithofacies. Carbonates formed within these Facies Zones often exhibit specific Standard Microfacies Types (SMF) assemblages that are used as additional criteria in recognizing the major facies belts (see Sect. 14.3). The definitions of the Facies Zones in Box 14.1 consider Wilson (1970, 1975), recent suggestions by Schlager (2002) and modifications regarding the diagnostic criteria. Facies Zone 9 is subdivided into 9A and 9B to distinguish the significant differences in arid evaporitic and in humid climates. A new Facies Zone 10 is introduced for limestones affected by freshwater dissolution and precipitation processes (e.g. in karstic and pedogenic settings). –> Note that these facies belts are limited to tropical platforms and do not consider platforms in coolwater settings that often correspond better to nonrimmed platforms or ramps. The Wilson model contains more Facies Zones than can normally be found on one platform. Rimmed platforms usually exhibit a reduced number of Facies Zones and often a different lateral arrangement of facies belts (Fig. 14.2).
14.1.3.2 Discussion and Use of Standard Facies Zones The Wilson model has passed the test of time with flying colors despite its static approach and the broad generalizations of the original concept. However, one should be aware that the model is just a snapshot illuminating potential depositional patterns and their lateral relationships. Critical points which must be considered are: • The model is overcomplete. Platforms with a reduced number of Facies Zones are the common case (Fig. 14.2). • The model prefers sharp boundaries between Facies Zones. Lateral facies boundaries, however, are often very gradual (e.g. between the Facies Zones 7 and 8) or can not be distinguished at all (e.g. FZ 3 and 4). • The model underlines the occurrence of litho- and biofacies within specific Facies Zones, but does not consider the development of Facies Zones over time.
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• The model does not show the effects of sea-level fluctuations. It fits predominantly sedimentation patterns formed during transgressive highstand and stillstand phases. • The model is limited to tropical and subtropical platforms and does not include climatic controls, except for the Facies Zones 9A and 9B. Facies distribution of temperate and cool-water platforms can not be interpreted with the model and are better described using models dealing with non-rimmed platforms or carbonate ramps. • The model does not sufficiently consider the changes in the organic platform rims that took place over time in the evolution of reefs. Facies Zone 5, comprising platform-margin reefs and upper slope reefs of various construction, composition and development stages, is too generalized and must be analyzed independently of the other Facies Zones. The specific methods are discussed in Sect. 16.2. • The model says nothing about windward-leeward differentiation resulting in distinct geometrical asymmetries that can be revealed by seismic surveys. • The model neglects the wide facies spectrum of oceanic deep-water settings and gives a much too generalized picture of slope sedimentation. Application of the Standard Facies Zone concept Despite the pitfalls listed above, the facies zonation established by Wilson has been successfully used by many authors involved in basin analyses and reservoir studies, e.g. Hubmann (1992, Devonian), Herbig (1984, Carboniferous), Schott (1984, Late Triassic and Early Liassic), Fezer (1988, Late Jurassic), Morrow and Webster (1991, Mississippian-Pennsylvanian), Mette (1993, Jurassic), Steuber et al. (1993, Cretaceous), Braun (1998, Triassic), Blomeier and Reijmer (1999, Jurassic), and Scheibner and Reijmer (1999, Jurassic). The Wilson model assists in • evaluating outcrops with regard to their paleogeographical position within a likely shelf-to-basin transect; • recognizing potential depositional and environmental relationships between widely spaced stratigraphic sections, particularly in remote areas; • attributing core samples to depositional sites that allow the evaluation of economically interesting facies zones (e.g. within reef or shoal complexes or at platform margins) and of depositional conditions that might have created appropriate poroperm values; • reconstructing ‘lost’ depositional sites by looking for source areas of carbonate clasts found in slope and basinal sediments, and derived from the destruction of carbonate platforms.
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Definition of Standard Facies Zones
Fig. 14.1. Rimmed carbonate platform: The Standard Facies Zones of the modified Wilson model. Box 14.1. Standard Facies Zones (FZ) of the modified Wilson model describing a rimmed carbonate platform. Basin and deep shelf FZ 1A Deep Sea Setting: Below wave base and below the euphotic zone in oceanic deep water. Water depth several hundreds to several thousands of meters. Wide facies belt. Sediments: Entire suite of deep-sea sediments including pelagic clay, siliceous and carbonate ooze, hemipelagic muds, turbidites. Adjacent to platforms mixtures of pelagic and platform-derived material (peri-platform oozes and muds). Bedding highly variable, often thin-bedded. Rock color: dark, reddish or light depending on differences in oxidizing and reducing conditions. Biota: Predominantly plankton, typical oceanic assemblages, sometimes associated with autochthonous benthic fossils. In peri-platform sediments up to 75% shallow-water benthos. Common lithofacies: Pelagic mudstone and wackestone; marls; allochthonous packstone, grainstone, breccia. FZ 1B Cratonic deep-water basin Setting: Below wave base, below the euphotic zone. Water depth about 30 m to several 100s m. Wide facies belt. Sediments: Similar to 1A. Hemipelagic muds very common. Occasionally anhydrite. Sometimes common cherts. Anoxic conditions fairly common (high organic content; lack of bioturbation). Dark thin limestone beds and dark shale beds. Lime mudstones, calcisiltites. Rock color: dark brown and black (due to organic matter) and reddish (due to slow sedimentation). Biota: Predominantly nekton (e.g. ammonites) and plankton (radiolarians, pelagic foraminifera, calpionellids, coquinas of thin-shelled bivalves). Occasionally benthos (abundant sponge spicules). Common lithofacies: Lime mudstone, wackestone, packstones. Marls. Anhydrite. FZ 2 Deep shelf Setting: Below fair-weather wave base but within the reach of extreme storm waves. Within or just below the euphotic zone. Forming plateaus between active platforms and deeper basins. The plateaus are commonly established on top of drowned platforms. Water depth tens of meters to hundreds of meters. Normal salinity, oxygenated waters with good current circulation. Wide facies belt. Sediments: Mostly carbonate (highly fossiliferous limestone) interbedded with marl beds. Skeletal wackestone and whole fossil wackestone; some grainstone and coquinas. Matrix commonly pelmicrite. Some silica. Well bioturbated. Well bedded. Bedding thin to medium, wavy to nodular. Rock color: gray, green, red and brown depending on variable oxidizing and reducing conditions.
Biota: Diverse shelly fauna indicating normal marine conditions. Infauna and epifauna. Minor plankton. Stenohaline biota conspicuous (e.g. brachiopods, echinoderms). Common lithofacies: Wackestone. Occasional grainstones. Marls and shales. Toe-of-slope and slope FZ 3 Toe-of-slope apron (deep shelf margin) Setting: Below wave base and barely at oxygen level. Moderately inclined sea floor (over 1.5°) basinward of steeper slopes. Water depths similar to FZ 2 and perhaps 200 to 300 m. Narrow facies belt. Sediments: Mostly pure fine-grained carbonates, in some places cherty, rare intercalations of terrigenous muds. Pelagic material admixed with fine-grained detritus moved off from adjacent shallow shelves. Grain size highly variable. Typical are well-defined graded beds or breccia layers (turbidites, debris-flow deposits) intercalated in finegrained background sediment. Rock color: dark to light. Biota: Mostly redeposited shallow-water benthos; some deep-water benthos and plankton. Common lithofacies: Lime mudstones; allochthonous packstones and grainstones. Shale partings. FZ 4 Slope Setting: Distinctly inclined sea floor (commonly 5° to nearly vertical) seaward of platform margins. Very narrow facies belt. Sediments: Predominantly reworked platform material and pelagic admixtures. Highly variable grain size. End members are gentle muddy slopes with much slumping, and sandy or rubbly slopes with steep foresets. Rock color: dark to light. Biota: Mostly redeposited shallow-water benthos, encrusting slope benthos and some deep-water benthos and plankton. The facies may be very fossiliferous. Common lithofacies: Mudstone; allochthonous packstone and grainstone; rudstone and floatstone. Breccia. Upper slope reefs and platform-margin reefs FZ 5 Platform-margin reefs Setting: (a) Organically stabilized mud mounds on upper slope; (b) ramps with knoll reefs and sand shoals; (c) wave-resistant barrier reefs rimming the platform. Water depths generally some meters,but some hundred meters for mud mounds. Very narrow facies belt. Sediments: Almost pure carbonates of very variable grain size. Massive limestones and dolomites. Masses or patches of various types of boundstones. Reef cavities filled with internal sediment or carbonate cements; multiple generations of construction, encrustation, boring and destruction. Rock color: light.
Definition of Standard Facies Zones
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Box 14.1. Definitions of Standard Facies Zones (continued). Biota: Almost exclusively benthos. Colonies of framebuilders, encrusters, and bafflers along with large volumes of loose skeletal rubble and sand containing benthic microfossils (e.g. foraminifera, algae). Common lithofacies: Framestone, bafflestone, bindstone, wackestone and floatstone, grainstone and rudstone. Platform-edge and platform sand shoals FZ 6 Platform-margin sand shoals Setting: Elongate shoals, tidal bars and beaches, sometimes with eolianite islands. Above fair-weather wave base and within the euphotic zone, strongly influenced by tidal currents. Very narrow facies belt. Sediments: Clean calcareous, often rounded, coated and well-sorted sands, occasionally with quartz. Sand grains are skeletal grains, or ooids and peloids. Partly with wellpreserved cross-bedding, sometimes bioturbated. Susceptible to subaerial exposure. Rock color: light. Biota: Worn and abraded biota from reefs and associated environments. Low-diversity infauna adjusted to mobile substrate. Common biota are large bivalves and gastropods and special types of foraminifera and dasyclads. Common lithofacies: Grainstone, packstone. Open-marine platform FZ 7 Platform interior - normal marine (open marine) Setting: Flat platform top within euphotic zone, normally above fair-weather wave base. Called lagoon when protected by sand shoals, islands or reefs of the platform margin. Sufficiently connected with the open sea to maintain salinity and temperature close to that of the adjacent ocean. Moderate circulation. Water depths a few meters to tens of meters. Wide facies belt. Sediment: Lime mud, muddy sand and clean sands, depending on the grain size of local sediment production and the efficiency of winnowing by waves and tidal currents. Medium to coarse bedded. Locally patch reefs or organic banks. Terrigenous sand and mud may be common in attached platforms, but are generally absent in detached platforms such as oceanic atolls. Rock color: light and dark. Biota: Shallow-water benthos with algae, foraminifera, and bivalves; gastropods, particularly common. Areas with marine grasses and with patch reefs. Common lithofacies: Lime mudstone, wackestone and floatstone, packstone, grainstone. Restricted-marine platform FZ 8 Platform interior - restricted Setting: As for Facies Zone 7, but less well connected with the open ocean, causing large variations in salinities and temperatures. Within the euphotic zone. Typically strongly differentiated tidal zones with freshwater, saltwater and hypersaline conditions as well as subaerially exposed areas. Shallow, cut-off ponds and lagoons with restricted circulation and hypersaline water. Lagoons behind barrier reefs, within atolls or behind coastal splits. Water depths below one meter and a few meters to a few tens of meters. Wide facies belts. Sediments: Mostly lime mud and muddy sand; some clean sand. Terrigenous influx common. Early diagenetic cementation common. Limestones and dolomites. Rock color: light.
Biota: Shallow-water biota of reduced diversity, but commonly with high number of individuals. Typical are miliolid foraminifera, ostracods, gastropods, algae and cyanobacteria. Marine and freshwater vegetation. Common lithofacies: Lime mudstone and dolomite mudstone, wackestone, grainstone, bindstone. Sedimentary breccia. Arid near-coast evaporitic platforms FZ 9A Arid platform interior - evaporitic Setting: As in Facies Zone 7 and 8, yet with only episodic influx of normal marine waters and arid climate so that gypsum, anhydrite or halite may be deposited beside carbonates. Supratidal. Sabkhas, salt marshes, salt ponds. Wide facies belt. Sediments: Calcareous or dolomitic mud or sands, with nodular, wavy or coarse-crystalline gypsum or anhydrite. Intercalations of red beds and terrigenous eolianites in landattached platforms. Rock color: highly variable; light, yellow, brown, red. Biota: Little indigenous biota except cyanobacteria; ostracods, mollusks, brine shrimps adapted to high salinities. Common lithofacies: Laminated lime and dolomitic mudstones and bindstones alternating layers with layers of gypsum or anhydrite. Humid near-coast brackish regions FZ 9B Humid platform interior - brackish (humid) Setting: Poor connection with the open sea as in FZ 9A but with a humid climate so that water runoff dilutes small bodies of ponded seawater and marsh vegetation spreads in the supratidal flats. Narrow facies belt. Sediments: Calcareous marine muds or sand with occasional freshwater lime mud and peat layers. Rock color: gray, light, brown, dark. Biota: Shoalwater marine organisms washed in with storms plus organisms adapted to brackish-water and freshwater (ostracods; freshwater snails; charophycean algae). Paleokarst, caliche and other terrestrial and terrestrialmarine settings FZ 10 Humid and arid often subaerially exposed, meteorically influenced limestones Setting: Subaerial or subaquatic, formed under meteoric-vadose and marine-vadose conditions. Abundant in karst settings and pedogenic carbonates (continental and near-coast areas), and supratidal and intertidal environments. Sediments: Limestones affected by early diagenetic meteoric dissolution predominantly during phases of subaerial exposure (e.g. paleokarst). Common in caliche crusts. Typically occurring in limestones rich in carbonate cement crusts, but also occurring in micritic caliche or as reworked grains in restricted environments (e.g. coastal ponds or lagoons). Biota: Indigenous biota lacking except cyanobacteria and microbes.
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Carbonate Ramp Model
Fig. 14.2. Common variations in the lateral arrangement of facies zones distinguished in the rimmed platform model (see Fig. 14.1). The arrangement of the facies belts is controlled by the position of the sea level (A highstand, B stillstand, C lowstand). A: The platform margin is characterized by a barrier reef zone (FZ 5) interfingering with an open-marine back-reef zone (FZ 7) of the inner platform. Facies zones FZ 3 and FZ 4 of the Wilson model are not distinctly differentiated due to the mixture of coarse reef debris and calcareous sand from the platform interior. Note the absence of the deep shelf facies zone FZ 2. B: The reef zone (FZ 5) is part of a reef complex consisting of lagoonal back-reef sediments and a slope whose upper part is supplied with talus forming a forereef zone. C: The sketch shows a common case whereby marginal sand shoals interfinger with slope sediments consisting of bioclastic sand transported from the shoals to gentle slopes. Adapted from Schlager (2002).
–> Most studies combine the evaluation of Facies Zones with the determination of Standard Microfacies Types. The methods and potential of this approach are described in Sect. 14.3.
14.1.4 Carbonate Ramp Model A carbonate ramp is a gently dipping sedimentary surface on the sea floor. The facies belts are controlled primarily by energy levels (fair-weather wave base and storm wave base), variations in ramp topography, and material transport by storms, waves and tides. The depositional slope gradient from the shallow-water shoreline or lagoon to the basin floor is of the order of a few meters per kilometer and usually less than 1°. Nearshore, wave-agitated shallow-water carbonates pass gradually offshore into deeper-water, low-energy deposits and then into basinal sediments. Ramps were common during the Phanerozoic, particularly in times when frame- and reef-building organisms were missing or rare. Most ramps have been described from the Cambrian and Early Ordovician, Early Devonian, Early Carboniferous, Early and Middle Triassic, Middle and Late Jurassic, Early Cretaceous and Early Tertiary. Carbonate ramps can develop during the drowning of shelves and during the early stages of platform formation. Often they evolve into rimmed platforms. Contrasted with rimmed shelves, ramps lack the steep slope at the shelf edge, show no continuous reef trends and exhibit no shallow-water derived clasts in deeper-wa-
ter outer ramp deposits. High-energy carbonates are formed near the shoreline in inner-ramp settings or in shoal areas of inner and mid-ramp settings, and not at the shelf edge. Windward ramps are storm- and wavedominated and characterized by grainy sediment at the shoreline. Leeward ramps are characterized by muddy sediments with low grainstone content. Dimension of ramps The widths and lengths of ancient carbonate ramps vary within a wide range. Maximum width ranges between less than 10 km to about 800 km. Most values are below 200 km and many ramps have widths of only <10 to about 20 km. The minimum and the maximum lengths range between 10 km and 1600 km, but some of the Early Paleozoic ramps extending beyond 1000 km are better classified as epeiric ramps. Most values are below 500 km. Ramps with lengths of about 10 km to 200 km are common. The dimensions of Cenozoic ramps are usually within the range of some hundreds of kilometers (West Florida, Cretaceous to present: width 200 km, length > 800 km; Trucial Coast, recent: width 200 km, length > 400 km; Shark Bay, Western Australia, recent: width 100 km, length 200 km; Yucatan Peninsula, Tertiary to present: width 200 km, length more than 600 km). Yet smaller Holocene wave-dominated ramps with a length of only about 50 km are known (Kuwait, northern Arabian-Persian Gulf). Characteristics of ramps The principal criteria of carbonate ramps are • Gentle slopes on which updip nearshore shallow-
Classification of Ramps
water deposits (mainly grainstones) pass gradually downslope without a marked break into progressively deeper water and low-energy deposits, and finally into basinal muddy sediments. • Because oceanic waves and swells sweep directly onto and across the shallow sea floor, the energy level of the shallow-water environments is high. • Complex and differentiated nearshore facies. • Continuous carbonate production takes place in shallow inner and mid ramp settings. • Storm-generated high-energy sediments are commonly deposited in inner and mid-ramp settings. • Autochthonous carbonate production by benthic organisms occurs in outer and deeper ramp settings. • Shallow-water sediments are transported into deeper water ramp parts. • Separate and discrete buildups (e.g. mud mounds or pinnacle reefs) may exist on mid- and outer ramps. Classifications of ramps Several classifications have been proposed for ramp geometry, water depth and the controls on ramps by storm wave base and fair-weather wave base (Ahr 1973; Read 1982, 1985; Aigner 1984; Carozzi 1989; Calvet and Tucker 1988). Read (1982, 1985) divided ramps into two major groups on the basis of the sea bottom profile (Fig. 14.3): • Homoclinal ramps with gentle, relatively uniform slopes without a shelf break exhibiting the same gradient from shoreline to deeper water, and • Distally steepened ramps with an offshore slope break between shallow ramp and the basin in relatively deep water. Distally steepened ramps have characteristics of ramps and of rimmed shelves, but they differ from the latter in that the major break in slope does not occur at the seaward margin of a high-energy rim but instead many kilometers seaward of high-energy shoals in around the mid- or outer ramp. The ramp builds laterally seaward with turbidites and slumps along the steepened slope. The processes of shaping ramps is aggradational in homoclinal and progradational in distally steepened ramps (Gardulski et al. 1991). Homoclinal and distally steepened ramps are often successive stages in the development of carbonate buildups, resulting in the growth of rimmed carbonate platforms. Both homoclinal and distally steepened ramps comprise tidal flat and lagoonal facies and foreshore or offshore complexes of shallow-water oolitic and bioclastic sand shoals, but differ in the sediments of deeperwater and basinal environments. Deeper homoclinal
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ramps are characterized by fossiliferous wackestones/ lime mudstones and marls with open-marine diverse biota. Deeper distally steepened ramps exhibit similar facies but may have slumps, breccias and turbidites along ramp-to-basin margin. Slumps and channeled or sheet-like breccias with clasts derived from the deepwater slopes are also common in the basin sediments consisting of lime mudstones and wackestones with interbedded marls. Ramp types: Read (1985) recognized six major ramp types based on the character of the facies formed under the highest energy conditions in shallow parts of the ramp. The author distinguished ramps with fringing banks, with barrier-bank complexes, with isolated shallow and downslope buildups, with fringing ooid shoals, with ooid-peloid barrier, and coastal beach/dune complexes. This classification addresses the essential role of sand bodies in the formation of ramps and considers the role of sessile organisms contributing to the formation of mounds and small patch reefs within ramp environments. Similar ideas are included in the concept used by Carozzi (1989) to describe the microfacies of ancient ramps. The author distinguished between simple ramps, ramps with bioaccumulated to hydrodynamic buildups, and ramps with bioconstructed buildups. The term ‘bioaccumulated’ refers to the accumulation of loose skeletal fragments (e.g. of crinoids), the term ‘bioconstructed’ to frameworks formed by colonial organisms. Hydrodynamic buildups are formed by the deposition of skeletal grains and ooids, subsequent to storms and current transport. Carozzi’s book summarizes the results of many case studies of ramp and platform carbonates, predominantly from the Paleozoic. Most ramp classifications recognize two critical hydrodynamical interfaces – the fair-weather wave base and the storm wave base. These boundaries can be recognized from sedimentary structures, microfacies criteria and biota. Diagnostic criteria are discussed in Sect.12.1.1.1 and Sect. 12.1.2.1. Combining the fair-weather wave base and the storm wave base, Burchette and Wright (1992) have suggested the subdivision shown in Fig. 14.3 and described in Box 14.2. This model is based on the influence of storms, waves and tides on the sediments in mid- and inner ramp zones. Many Phanerozoic ramps were strongly storm-controlled and therefore exhibit a threefold subdivision. In the case of missing indications of storms, inner and outer ramps can be differentiated by using wave base, biotic composition and dominant limestone textures.
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Carbonate Ramp Model
Fig. 14.3. Carbonate Ramp Model of a homoclinal ramp, modified from Burchette and Wright (1992). The differentiation into an inner, mid- and outer ramp can be applied both to homoclinal and distally-steepened ramps (see inset). Sediment texture, grain size, and biotic criteria depend on the water energy at the sea bottom differing above the fair-weather wave base, between the fair-weather wave base and the storm wave base, and below the storm wave base. The slopedip is strongly exaggerated. Not to scale. Ramp lengths vary between some tens and some hundreds of kilometers.
Box 14.2. Carbonate Ramp Model (Burchette and Wright 1992). Criteria of inner, mid- and outer parts of ramps. Inner ramp The inner ramp comprises the euphotic zone between the upper shoreface (beach or lagoonal shoreline) and the fair-weather wave base.The sea floor is almost constantly affected by wave action. The zone is dominated by sand shoals or organic barriers and shoreface deposits. The shallow inner ramp may consist of (1) a beach barrier-tidal delta complex with lagoons and tidal flats behind (backramp), or (2) fringing sand banks and shoal complexes with intertidal and supratidal flats, but no lagoons behind, or (3) a strandplain of linear beach ridges with depressions. Characteristic sediments are carbonate lime sand bodies formed in agitated, shallow subtidal shoreface areas above the fair-weather wave base. The sands consist predominantly of ooids or various skeletal grains, usually foraminifera, calcareous algae, or mollusks. Peloids may be common in places. Storms contribute to the formation of extended sheet-like sand bodies and sand beaches that may grade into eolian dunes. Offshore storm surges transport shoreface sands to deeper, outer ramp settings. Organic buildups in inner-ramp environments are biostromes and small patch reefs characterized by low-diversity biota (e.g. corals, rudists, oysters). Frequent limestone types are grainstones and packstones. Back-ramp sediments originate in peritidal settings similar to those of inner platforms (comprising mudstones, bindstones and wackestones), and in restricted lagoonal areas (mudstones, wackestones, packstones). Mid-ramp The mid-ramp is the zone between fair-weather wave base and the storm wave base. Water depths reach some tens of meters. The bottom sediment is frequently reworked by storm waves and swells. The sediments reflect varying degrees of storm influence depending on the water depth and bottom relief. Intraclast and breccia beds may be common.
Thick oolitic and bioclastic sand shoals are common. Storm-related features are graded packstone, grainstone beds, hummocky cross-stratification, and tempestite couplets. Skeletal grains exhibit signs of transport. Fair-weather phases are represented by burrowed sediments dominated by lime mud or terrigenous mud forming lime mudstones and marls. Much of the fine-grained sediment might be caused by lateral sediment transport in offshore zones or by transport from the shoreline to midand outer ramp areas. Mid-ramp deposits are often thicker than coeval inner ramp deposits. Organic buildups are represented by pinnacle reefs and mounds. Outer ramp The outer ramp is the zone below normal storm wave base. Water depths vary between tens of meters and several hundreds of meters. The zone is characterized by lowenergy allochthonous and autochthonous carbonates, and hemipelagic sedimentation. Little evidence of direct storm reworking exists, but various storm-related deposits (e.g. graded distal tempestite beds) may occur. Common lithofacies types are bedded, fine-grained limestones (argillaceous lime mudstone and wackestone) associated and interbedded with marl or shale beds. Calcisiltite matrix is abundant. Biota comprise normal marine diverse benthos, sometimes associated with plankton and nekton. Benthic organisms include foraminifera, sponges, bryozoans, brachiopods, mollusks, and echinoderms. Algae may be represented by red algae. Burrows are common. In deeper outer ramp settings restricted bottom conditions may develop. Common organic buildups are mud mounds. The slope break of distally steepened ramps is usually located in a position around the mid- or outer ramp boundary or within the outer ramp. Deposition of slope-derived material may dominate proximal to the break.
Non-Rimmed Platforms
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14.1.5 Non-rimmed Shelves and Platforms
14.1.6 Isolated Platforms and Atolls
A non-rimmed shelf or open platform is characterized by the lack of a barrier at the shelf break. Non-rimmed platforms and ramps are similar in the absence of a rim at the shelf edge. In both systems sediment transport occurs onshore and downslope, and carbonate production takes place in all parts. The major difference, however, is the very gentle slope angle of the ramp, producing different distribution patterns and sizes of facies zones as compared with unrimmed platforms. Non-rimmed shelves occur at the leeward side of large tropical banks and are abundant in all cool-water settings. Fig. 14.4 displays the subdivision of a nonrimmed carbonate cool-water shelf. Box 14.3 summarizes the characteristics. Mid-shelf and outer shelf sediments of cool-water carbonates differ in the composition of the bioclastic sands (see James and Clarke 1997 for case studies).
Isolated platforms are accumulations of shallow-marine carbonates surrounded by deep-water sediments. This non-attached platform type includes very large structures as well as small platforms found on submarine highs and reef atolls growing upon volcanic basements. Most isolated platforms have steep margins and slopes reaching into deep or very deep water. Outer high-energy parts are characterized by marginal reefs and/or sand bodies. Common features of the platform interior are muddy and sandy calcareous sediments. These sediments differ in composition, sedimentary structures, biota and microfacies. High carbonate production around the platform margins and a steady background subsidence may lead to the formation of a rimmed isolated platform with a lagoon in the center. This pattern has been called an ‘empty bucket’ by Schlager (1993).
Fig. 14.4. Hydrodynamic zones and subdivision of a non-rimmed cool-water (temperate) shelf. Base of wave abrasion ranges from 30 m to 70 m, swell wave base may reach 120 m, storm wave base about 250 m. Modified from James (1997). Box 14.3. Criteria used in the subdivision of a non-rimmed carbonate cool-water shelf. After James (1997). Inner shelf Depositional processes: Constant wave agitation. Particle abrasion and bioerosion. Winnowing. Sediment: Zone of sediment movement and active sediment production. Gravels, lithoclastic sands and hard substrates. Subaqueous dunes. Shaved shelf areas. Biota: Coralline red algae, benthic foraminifera, bryozoans, sponges, bivalves, gastropods, serpulids, echinoids. Deposition of epibionts from high-energy kelp forests and low-energy sea grass. Mid-shelf Depositional processes: Frequent storm reworking. Particle abrasion. Sediment transport to outer and inner shelf areas results in sediment-free areas. Bioerosion and burrowing common. Sediment: Zone of active sediment production. Thin
sediment veneer over lithified bedrock. Coarse bioclastic sand. Rippled sands, subaqueous dunes. Biota: Coralline red algae, mollusks, benthic and planktonic foraminifera, bryozoans, brachiopods, sponges, barnacles, echinoids. Outer shelf Depositional processes: Sea bottom reworked by episodic storms. Suspension settling. Bioerosion and burrowing common. Sediment: Zone of carbonate production and accumulation. Fine bioclastic sands. In deeper parts mud (consisting of a mixture of calcitic plankton and skeletal fragments, siliceous sponge spicules, and clay). Burrowed sediments and storm beds. Biota: Bryozoans, sponges, mollusks, brachiopods, benthic and planktonic foraminifera.
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A well-known example of an isolated platform is the Great Bahama Bank (see Pl. 3). Controls on facies distribution are the windward or leeward orientation, open or protected water circulation, or tide- or stormdominated conditions.
Isolated Carbonate Platform
An example of a smaller isolated platform in the Caribbean is shown in Fig. 14.5. The study is based on the thin-section investigation of the grain composition, grain size, and texture of plastic-impregnated sieve samples.
Fig. 14.5. Facies distribution on an isolated carbonate platform. Turneffe Islands, Belize, Central America. Turneffe Islands are surrounded by deep waters. The distribution of the sedimentary facies is controlled by antecedent topography and exposure to waves and currents. The maps show the relationships between environment, facies types and carbonate texture. A: Environments. The platform has an almost continuous surface-breaking reef rim at the windward side. The reef margin is characterized by a cemented coral rubble pavement transected by tidal passes and channels. Windward and leeward forereef areas dip toward the open ocean. Sand aprons are developed in water depths down to 2 m. The bottom of shallow lagoons and deep lagoons (> 5 m) consists of unconsolidated sediment and is either barren or covered with sea grass and udoteacean green algae. Restricted lagoons with depths of 8 m and 3 m are enclosed by mangrove swamps and islands. The bottom of the lagoons is covered by sea grass, Halimeda and sponges. Subtidal ponds are located within mangrove rims. The cays/islands consist of sand and mangroves. B: Facies types expressed by the composition of the sediment. The coral-red algae-Halimeda facies is most abundant in the marginal platform areas as well as in platform interior patch reefs. The mixed peloidal/skeletal facies is characterized by the abundance of non-skeletal grains (predominantly peloids of fecal origin). The mollusk-foraminifera-Halimeda facies is common in platform interior environments. It is characterized by high concentrations of bivalve shells, miliolids and algal fragments. The sediments of the Halimeda facies contain more than 50% algal fragments. It is confined to the restricted lagoons. C: The texture of the surface sediments (using a modified Dunham classification) reflects the hydrodynamic conditions. The distribution of the texture types is similar to the distribution of facies types in ancient isolated reefal platforms. Boundstones characterize the windward reef rim. Grainstone textures are abundant in the reef and forereef environments of the coral-red algae-Halimeda facies. Packstones are most common in this facies in the sand apron, shallow lagoon, and patch reef environments, and occur in the mollusk-foraminifera-Halimeda facies. The mixed peloidal-skeletal facies is characterized by wackestone and packstone textures. After Gischler and Lomando (1999).
Epeiric Platform and Epeiric Ramp
14.1.7 Epeiric Platform Model During the Phanerozoic epeiric seas covered extensive areas of the cratons. These very shallow, low-energy seas extended for hundreds to thousands of kilometers. Epeiric seas first flooded the margins and later the interior of tectonically stable cratons. Modern epeiric seas are rare. Examples of warm-water epeiric seas are the Sunda Sea and Java Sea, cool-water examples are the Baltic Sea, the North Sea, and the Hudson Bay (Edinger et al. 2002). Depositional patterns in epeiric seas are discussed in the epeiric platform model and the epeiric ramp model (Fig. 14.6). The epeiric platform model postulates clear-water sedimentation extending over vast regions from hundreds to thousands of kilometers wide, assuming that tides were damped out by friction with the shallow sea bottom and salinities increased seaward (Irwin 1965). The depositional site should be characterized by a more or less negligible, extremely low slope angle and regionally extended low-energy conditions and distinct salinity gradients. The inner platform is believed to have been characterized by subtidal to intertidal mud flats with widths of tens to hundreds of kilometers, and water depths generally less than 10 m. The model differentiates a low-energy, below wave base zone X of the open sea, a relatively narrow zone Y of higher energy where waves impinge on the sea floor and strong tidal currents develop, and an extended zone Z with restricted water circulation, minimal tidal effects, only periodic storm effects, and hyper- and/or hypo-salinity conditions. The tideless mode of the model in zone Z was questioned by Pratt and James
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(1986) who proposed a ‘tidal flat island model’ assuming the existence of small supratidal low-relief islands and intertidal banks surrounded by open subtidal waters. This model overcomes many contradictions inherent in fitting facies criteria to the epeiric platform model, and adapts the tideless model to many Early Paleozoic limestones and dolostones characterized by supra-, inter- and subtidal cycles.
14.1.8 Epeiric Ramp Model The epeiric ramp model was proposed by Lukasik et al. (2000) to describe temperate Tertiary carbonates from the Murray Basin in Australia (Fig. 14.6B). Epeiric ramps and normal carbonate ramps differ in • the widths that may reach hundreds of kilometers (normal ramp 1–100 km), • very low and negligible slope angles (ramp <1°), • water depths of only a few tens of meters, less than 20 m (ramp 10 to 100 m), • low water energy and subtidal to intertidal mud flats in the inner part of epeiric ramps (ramps: high-energy shoreline with shoal/barrier-lagoon), • tidal flats extending hundreds of kilometers (ramps: within kilometer scale), • dominant processes, such as storms and nearshore tidal currents (on ramps: wind-induced waves and storms). Both the distal and the proximal parts of epeiric ramps are characterized by burrowed silty, fine-grained carbonate, forming mud flats near the coast and in distal sea-grass areas. Grainy sediments, produced by storms, are commonly obliterated by bioturbation. The
Fig. 14.6. Epeiric platform and epeiric ramp models. The models refer to very shallow depositional settings that extend over vast areas, comprising hundreds to thousands of kilometers for epeiric platforms, and hundreds of kilometers for epeiric ramps. In contrast, carbonate ramps usually range between a few kilometers to 100 or 200 km. The inner platform of ramps is characterized by high-energy facies, that of epeiric platforms and ramps by low-energy facies. A: The epeiric platform model of epicontinental seas (Irwin 1965) differentiates three generalized zones, two of which lie in low-energy environments (X and Z), and the third (Y) within higher energy areas. B: The epeiric ramp model (Lukasik et al. 2000) differs from the epeiric platform model in the nature of the slope and the extent of the basin.
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distribution of benthic organisms is controlled by progressively higher trophic resource levels (from lower mesotroph to eutroph) and restricted marine conditions from relatively offshore to nearshore regions. The sediments are typified by pervasive bioturbation and common to abundant fossils. Using variations in fossil content and lithologic composition, four major facies association have been differentiated: • Large foraminiferan-bryozoan facies association, comprising 3 compositional types; • Echinoid-bryozoan facies association, comprising 4 compositional types; • Mollusk facies association, comprising 6 compositional types; • Clay facies. These distal facies types yield predominantly larger foraminifera, bryozoans, mollusks, and echinoderms. Bioturbation is common in all facies types. The matrix is usually a microbioclastic packstone. Common texture types are bioclastic floatstones and rudstones characterized by large fossils within a microbioclastic packstone matrix. Almost all microfacies types would correspond to SMF 8 and SMF 9! The grain association of some facies types corresponds to the biomol, bryomol, foramol and rhodechfor association of cool-water environments. Basics: Carbonate depositional systems and facies zones Anderton, R. (1985): Clastic facies models and facies analysis. – In: Brenchley, P.J., Williams, B.J.P. (eds.): Sedimentology: recent developments and applied aspects. – 3147, Oxford (Blackwell) Edinger, E.N., Copper, P., Risk, M.J., Atmoryo, W. (2002): Oceanography and reefs of Recent and Paleozoic tropical epeiric seas. – Facies, 47, 127-150 Gischler, E., Lomando, A.J. (2000): Isolated carbonate platforms of Belize, Central America: sedimentary facies, late Quaternary history and controlling factors. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems: components and interactions. – Geological Society London, Special Publications, 178, 135-146 Handford, C.R., Loucks, R.G. (1993): Carbonate depositional sequences and system tracks - responses of carbonate platforms to relative sea-level changes. – In: Loucks, R.G., Sarg, H.F. (eds.): Carbonate sequence stratigraphy: recent developments and applications. – American Association of Petroleum Geologists, Memoir, 57, 3-41 Irwin, M.L. (1965): General theory of epeiric clear water sedimentation. – American Association of Petroleum Geologists, Bulletin, 49, 445-459 James, N.P., Kendall, A.C. (1992): Introduction to carbonate and evaporite facies models. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 265-276, Geological Association of Canada Lukasik, J.L., James, N.P., McGowran, B., Bone, Y. (2000): An epeiric ramp: low-energy, cool-water carbonate facies in a Tertiary inland sea, Murray Basin, South Australia. – Sedimentology, 47, 851-881
Basics: Carbonate Depositional Systems
Miall, A.D. (1990): Principles of sedimentary basin analysis. – 2nd edition, 668 pp., Berlin (Springer) Pratt, B.R., James, B.N. (1986): The St. George Group (Lower Ordovician) of western Newfoundland: tidal flat island model for carbonate sedimentation in shallow epiric seas. – Sedimentology, 33, 313-343 Schlager, W. (1992): Sedimentology and sequence stratigraphy of reefs and carbonate platforms. – American Association of Petroleum Geologists, Continuing Education Course Note Series, 34, 71 pp. Schlager, W. (1993): Accomodation and supply - a dual control on stratigraphic sequences. – Sedimentary Geology, 86, 111-136 Schlager, W. (2002): Sedimentology and sequence stratigraphy of carbonate rocks. - 146 pp., Amsterdam (Vrije Universiteit/Earth and Life Sciences) Tucker, M.E. (1985): Shallow-marine carbonate and facies models. – In: Benchley, P.E., Williams, B.P. (eds.): Sedimentology: recent developments and applied aspects. – Geological Society London, Special Publications, 18, 139161, London Walker, R.G. (1992): Facies, facies models and modern stratigraphic concepts. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 1-14, Geological Association of Canada Walker, R.G., James, N.P. (eds., 1992): Facies models. Response to sea level change. – 409 pp., Geological Association of Canada Wilson, J.L. (1975): Carbonate facies in geologic history. – 471 pp., New York (Springer) Further reading: K173, K184, K189, K190, K192, K193 Computer modeling Aigner, T., Brandenburg, A., Van Vliet, A., Doyle, M., Lawrence, D., Westrich, J. (1990): Stratigraphic modelling of epicontinental basins: two applications. – Sedimentary Geology, 69, 167-190 Aigner, T., Doyle, M., Lawrence, D., Epting, M., Van Vliet, A. (1989): Quantitative modeling of carbonate platforms: some examples. – SEPM Special Publications, 44, 27-37 Aurell, M., Bosence, D., Waltham, D. (1995): Carbonate ramp depositional systems from a late Jurassic epeiric platform (Iberian Basin, Spain): a combined computer modelling and outcrop analysis. – Sedimentology, 42, 75-94 Bice, D.M. (1991): Computer simulation of carbonate platform and basin systems. – In: Franseen, E.K., Watney, W.L., Kendall, G.St.C., Ross, W. (eds.): Sedimentary modeling: computer simulations and methods for improved parameter definition. – Bulletin, Kansas Geological Survey, 233, 431-447 Bosence, D.W.J., Pomar, L., Waltham, D.A., Lankester, T.H.G. (1994): Computer modeling a Miocene carbonate platform, Mallorca, Spain. – American Association of Petroleum Geologists, Bulletin, 78, 247-266 Bosscher, H., Schlager, W. (1992): Computer simulation of reef growth. – Sedimentology, 39, 503-512 Bosscher, H., Southam, J. (1992): CARBPLAT - a computer model to simulate the development of carbonate platforms. – Geology, 20, 235-238 Burgess, P.M. (2001): Modeling carbonate sequence development without relative sea-level oscillations. – Geology, 29, 1127-1130 Demicco, R.V., Spencer, R.J., Waters, B.C., Cloyd, K.C. (1991): Two-dimensional computer model of a Cambrian
Biozonation Patterns
shelf deposit. – In: Franseen, E.K., Watney, W.L., Kendall, G.St.C., Ross, W. (eds.): Sedimentary modeling: computer simulations and methods for improved parameter definition. – Bulletin, Kansas Geological Survey, 233, 463-472 Enos, P. (1991): Sedimentary parameters for computer modeling. – In: Franseen, E.K., Watney, W.L., Kendall, G.St.C., Ross, W. (eds.): Sedimentary modeling: Computer simulations and methods for improved parameter definition. – Bulletin, Kansas Geological Survey, 233, 63-100 Franseen, E.K., Watney, W.L., Kendall, G.St.C., Ross, W. (eds., 1991): Sedimentary modeling: computer simulations and methods for improved parameter definition. – Bulletin, Kansas Geological Survey, 233, 524 pp. Harbough, J.W., Watney, W.L., Rankey, E., Slingerland, R., Goldstein, R., Franseen, E.V. (eds., 2000): Numerical experiments in stratigraphy: recent advances in stratigraphic and sedimentologic computer simulations. – SEPM Special Publications, 62 Hüssner, H., Roessler, J. (1997): Mathematical modelling and numerical simulation of reefs. – Courier Forschungsinstitut Senckenberg, 201, 225-236 Read, J.F., Osleger, D.A., Elrick, M. (1991): Two-dimensional modeling of carbonate sequences and component cycles. – In: Franseen, E.K., Watney, W.L., Kendall, G.St.C., Ross, W. (eds.): Sedimentary modeling: Computer simulations and methods for improved parameter definition. – Bulletin, Kansas Geological Survey, 233, 473-488 Slingerland, R., Harbaugh, J.W., Furlong, K.P. (1994): Simulating clastic sedimentary basins. – Prentice Hall Sedimentary Geology Series, 220 pp. Strobel, J., Soewito, F., Kendall, C.G.St.C., Biwas, G., Bezdek, J., Cannon, R. (1990): Interactive simulation (SEDPAK) of clastic and carbonate sediments in shelf to basin settings. – In: Cross, T.A. (ed.): Quantitative dynamic stratigraphy. – 433-444, Englewood (Prentice Hall) Vahrenkamp, V., Grötsch, J. (1994): 3-D architecture model of a Lower Cretaceous carbonate reservoir: A 1 Huwaisah Field, Sultanate of Oman. – In: Al-Husseini. M.I. (ed.): Geo’94. – The Middle East Petroleum Geosciences, 2, 901915 Walkden, G.M., Walkden, G.D. (1990): Cyclic sedimentation in carbonate and mixed carbonate-clastic environments: four simulation programs for a desktop computer. – In: Tucker, M.E., Wilson, J.L., Crevello, P.D., Sarg, J.R., Read, J.F. (eds.): Carbonate platforms: facies, sequences and evolution. – International Association of Sedimentologists, Special Publication, 9, 55-78 Further reading: K197, K198
14.2 Biotic Zonation Patterns Shallow-marine benthic organisms modify substrate and water circulation resulting in the production of sediments and depositional geometries that are recognizably different from those that would originate without organisms. The distribution of benthic communities on platforms and ramps exhibits characteristic zonation patterns that can be used in recognizing facies belts and evaluating paleowater depths.
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14.2.1 Concepts and Methods Biotic distributional patterns across platforms or ramps can be defined by • the lateral distribution of higher systematic units (e.g. groups and subgroups of calcareous algae), • the average distribution and association of genera and their morphologies (e.g. of foraminifera), • the integration of fossil assemblages, limestone composition and microfacies data, and the use of ‘indicator fossils’, and • the distribution of benthic assemblages along shallow marine to deeper marine transects. Lateral distribution of higher systematic units: This approach is limited to groups whose lateral and depth distribution depend on only a few dominant ecological factors, such as light for the algae. Modern calcareous green algae occur in shallow restricted and open platform and inner ramp environments (FZ 8 and 7), calcareous red algae also in platform-edge and upper slope environments (FZ 6) and mid-ramp settings. Calcified cyanobacteria are concentrated in the restricted environments of facies zone FZ 8. This distribution pattern is similar in Tertiary platforms and in some Mesozoic platforms, but must be modified for Paleozoic shelves as demonstrated by Fig. 10.17. Modifications are necessary because of the specific distributional patterns of now extinct subgroups. Average distribution and association of genera and their morphologies: This approach is exemplified by Fig. 14.7. Functionally-controlled morphology and lifestyle of benthic foraminifera can be used to distinguish ramps and rimmed platforms (Geel 2000). Larger foraminifera are valuable paleoenvironmental indicators for Cenozoic carbonates (Hallock and Glenn 1986). The shape and thickness of the tests as well as distribution trends of modern larger foraminifera differ in inner and outer shelf regions. These differences are used in distinguishing sediments with foraminifera in different shelf areas. The distribution patterns reflect the standard facies zones of the Wilson model and can be used in recognizing the major depositional environments. Integration of fossil assemblages, limestone composition and microfacies data, and the use of ‘indicator’ fossils. Recurrent fossil distributional assemblages and limestone textures within specific time slices are the basis of stratigraphically limited facies models. These models represent summaries of case studies. The case studies were evaluated with respect to those sedimen-
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Foraminiferal Distribution Patterns
Fig. 14.7. The Standard Facies Zones (FZ) differentiated by Wilson (1975) for a rimmed carbonate shelf (Fig. 14.1) are reflected by distribution patterns of Cenozoic foraminifera. The model developed by Hallock and Glenn (1986) is based on the relative percentages of three environmentally diagnostic foraminiferal groups (flat, larger rotaliines plus planktonic foraminifera; lenticular larger rotaliines; miliolinids and smaller rotaliines). The three groups are distinguished by test morphologies, habitats of living relatives and R-mode multivariate analysis. All the foraminifera belonging to these categories are counted in thin sections and the proportions are plotted in triangle diagrams. A: Idealized distribution of Tertiary diagnostic foraminiferal groups in sediments of the standard facies zones FZ 1 to FZ 8 across a platform-basin transect with marginal reefs. Water depths are derived from modern data. B: Triangle diagram of four foraminifera assemblages in a core comprising Early Miocene reef-associated wackestones, packstones, grainstones and boundstones. The diagram indicates water depth, wave energy, and open marine depositional environments of the assemblages. Shallow-water, high energy conditions are indicated for faunas plotting to the lower left. Assemblages of protected environments plot to the lower right and deeper open-marine conditions toward the apex. C: Triangle diagram summarizing where foraminiferal assemblages from each of Wilson’s Standard Facies Zones FZ 1– FZ 8 would generally plot. B and C: Modified from Hallock and Glenn (1986).
tary and paleontological criteria that allow water energy and water depth to be interpreted. Figs. 14.8 to 14.11 demonstrate an example for facies zones occurring in Devonian reef, platform and ramp settings. Distribution of benthic assemblages along shallow marine to deeper marine transects: The composition of benthic assemblages varies in different parts of carbonate platforms and ramps. Compositional differences in shallow platforms occur on a species level (e.g. for foraminifera, mollusks, and green algae) as well as on a statistically based ecologically defined group level (e.g. reaction groups, Purdy 1963). Benthic foraminifera have been successfully used as facies indicators in modern and ancient carbonates (Murray 1991). The consideration of textural parameter of foraminiferal tests (grain size, sorting) allows
transported and resedimented tests, and autochthonous assemblages to be differentiated (Gischler 2003). Unlike rimmed platforms, the biotic composition of homoclinal carbonate ramps appears rather homogeneous, and lateral changes are often gradual, specifically in outer ramp environments. In these environments Standard Microfacies Types are of limited value. Ramps can be differentiated if the lateral and vertical distribution of microfossil assemblages and their shifts are used as the main criteria. The shifts can be used to recognize sequence stratigraphic boundaries and maximum flooding surfaces in deeper-water ramps. This is interesting, because in deep and relatively constant outer ramp environments, sequence stratigraphic transitions are very subtle and the boundaries of system tracts tend not to occur as unconformities at the contact of units. For using microfossil assemblages see Fig. 14.12. Text continued on p. 676
Devonian Coral-Stromatoporoid Limestone
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Fig. 14.8. Devonian coral-stromatoporoid limestone. A comparison of Devonian reef and platform carbonates from various parts of the world with regard to the biotic composition and growth types of sessile organisms, biofabric, sedimentary texture and microfacies displays astonishing similarities. These coincidences can be used to designate rough basic facies types (Hladil 1986; Machel and Hunter 1994), which are successful tools in recognizing facies zones in outcrop and subsurface samples. The thin section shows a reef limestone with a widely distributed association consisting of well-preserved tabulate corals (Alveolites – A), rugose corals (Tyrganolites – T) and stromatoporoids (Stachyodes – S, Clathrocoilona – C, Amphipora – AM). Larger growth voids are infilled with tiny micritic clasts, peloids and a few microfossils forming a grainstonepackstone matrix. Remaining void space is occluded by blocky calcite. The criteria of the sample correspond to those of the facies zone IVf of the Machel-Hunter Model (see Fig. 14.9). Both assignments point to a forereef environment situated adjacent to the active zone of reef growth. A ramp model would place the sample in a mid-ramp position (Fig. 14.11). The SMF classification would attribute the sample to the SMF Type 7. This assignment indicates ‘organic reef, platform margin’, but allows no furthergoing differentiation. Late Devonian (Late Frasnian): Moravian Karst, Czech Republic. After Hladil (1986). Practical advice: If you want to extract maximum information from this and similar limestones you should consider the following criteria: Fossils: How are they preserved? Try to differentiate fossils, if possible on a generic or even better, a species level! Determine or estimate the taxonomic diversity! Estimate the frequency of the fossils and their contribution to rock-building! Which fossils are abundant, which are rare? Are sessile organisms in life position, are they overturned, transported or how are they oriented? Are the fossils bored or encrusted? Are there recurrent encrustation patterns? Determine growth forms of the sessile organisms and their guild membership (constructors, bafflers, binders, encrusters, dwellers)! Which organisms are commonly intergrown? Matrix: Describe the matrix with regard to the texture (mud-, wacke-, pack- or grainstone), size, sorting, grading and roundness of matrix grains and type and amount of micrite! Porosity and cements: Porosity type (interskeletal, intraskeletal, shelter pores); cement types and fabric (use polarized light and cathodoluminescence!)
Devonian Reef Model
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Forereef
Backreef
Wave base
Reef core
I-f A Water energy
quiet water below wave base
Turbulence negligible
II-f
III-f
intermittently slightly agitated agitated water water
IV-f moderately agitated water
episodic
moderate
strong
Rounding
original poor fossil shapes to good
moderate to good
good to moderate
Sorting
angular fragments if broken
matrix: good, matrix: good to fossils: poor moderate, moderate to good fossils: moderate to good
B Texture
packstones, grainstones, mudstones
packstones, grainstones, rare floatstones
C Amount of micrite D Porosity types
–> –> –> primary micrite decreasing –> –> –> intraparticle
Pore size Porosity
below 10%
V
strongly agitated water strong
IV-b
strong
very good to moderate good to moderate matrix: moderate fossils: poor
good to moderate
rudstones, bindstones, rudstones, bindstones, framebafflestones rare stones, grainstones rudstones fibrous calcite, sparite, micrite
III-b
moderately slightly agitated agitated water water
II-b
I-b
intermittently intermittently agitated agitated water water
moderate
episodic
episodic
moderate to good
poor to good
original fossil shapes
matrix: matrix: good angular fragmoderate fossils: poor ments if fossils: mod- to good broken erate to good floatstones, packstones, mudstones, grainstones, grainstones wackestones rare rudstones
–> –> –> primary micrite increasing –> –> –>
intraparticle, interparticle, shelter shelter shelter interparticle shelter voids, voids, voids, voids, intraparticle growth interparticle growth framework, framework, interparticle interparticle
shelter voids, growth framework, interparticle
shelter intraparticle, intraparticle, voids, interparticle, fenestral intraparticle, fenestral interparticle
small
large to medium
medium to small
small
up to 40%
up to 10%
up to 20%
medium to small
up to 1% up to 15% mostly < 5%
calcispheres, small E Fossil assemblage foraminifera crinoids
F Indicator fossils
floatstones, rudstones, grainstones, bafflerare stones rudstones
IV / V
large crinoids
medium to small
large to medium
up to 10%
up to 25% up to 10%
tabulate corals, Stachyodes, brachiopods, crinoids, Girvanella, Renalcis
crinoids, tabular stromatoporoids
crinoids / goniatites / brachiopods
medium to small
bulbous + Stachyodes, Amphipora, algae Girvanella, Parathuram- (fenestral massive stromato- Renalcis mina, pores) poroids, ostracods micrite envelopes, syntaxial rim cement
crinoids, Solenopora, Keega
up to 30% algae (fenestral laminites)
Amphipora / calcispheres | Parathurammina / Vermiporella / encrusting Sphaerocodium
Fig. 14.9. Facies model for Middle and Late Devonian reef carbonates after Machel and Hunter (1994). The model differentiates five water energy zones with increasing water energy from I to V, designated as either f (forereef) or b (backreef). Use of the model requires an estimation of frequencies, differentiation of growth forms and some knowledge of fossils. Part A estimates water energy by noting rounding and sorting. Part B refers to the limestone texture. Part C records the amount of micrite. Part D lists the amount, size and the predominant type of porosity. Part E lists assemblages that are common in the different facies zones and have their distributional maxima within the zone(s). Part F lists ‘indicator fossils’ which are usually found only in backreef, forereef, or reef core settings. The criteria used in distinguishing facies zones can be observed in outcrops, cores as well as in large thin sections. See Fig. 14.8 to see how the model is applied.
Fig. 14.10. The location of reef growth and the development of the reef zones depend on the submarine topography. In the case of a more gently dipping slope, maximum reef growth is located deeper than in the example shown in Fig. 14.9. The shallow forereef zone IV-f is missing or only incompletely developed. After Machel and Hunter (1994).
Biotic Ramp Zonation
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Fig. 14.11. Integrated reef/ramp model for Devonian carbonates (modified from Machel and Hunter 1994). The reef model explained in Fig. 14.9 (facies zones indicated by Roman numbers) and the ramp model (facies indicated by small letters) are complementary. They overlap in space and partially substitute for one another. An integration of the models shows the facies distribution from the supratidal zones across a hypothetical carbonate shelf to the basin. The main differences occur in the outer shelf region where reefs develop in the case of a distinct shelf break. In the case of a gently sloping ramp, the litho/ biofacies a and b will develop. Patch reefs (c) are most likely to form within facies b and d. a - Burrowed argillaceous mudstones and wackestones; fossils rare; scattered brachiopods and crinoids. b - Mudstone with carbonate nodules in shaly matrix. c - Patch reefs exhibiting stromatoporoid-coral floatstones to rudstones with bulbous and tabular stromatoporoids; and a wackestone-packstone matrix. d - Amphipora floatstones and grainstones. e - Vuggy, poorly fossiliferous mudstones. f - Laminated and fenestral dolomite. g - Mudstone breccia, clasts correspond to reworked facies f.
Fig. 14.12. Differentiation and sequence stratigraphic interpretation of an outer homoclinal carbonate ramp based on microfossil assemblages. The figure displays the biotic gradients of Early Mississippian ramp carbonates in the Sacramento Mountains of south-central New Mexico (Jeffery and Stanton 1996). Black dots on the horizontal line show the position of measured and sampled stratigraphic sections. The length of the bars indicates the lateral extent of fossil assemblages. The letter S indicates the southernmost occurrence of sponge spicules generally regarded as an indication of relative deep-water environments. The gently dipping thin- to medium bedded carbonates (mostly bioclastic wackestones) of the Alamogordo Member can be laterally traced bed-per-bed for more than 30 km. The thickness of the member is 1.5 m in the North and 10.5 m in the South. The sediment is interpreted as an autochthonous deposition because of the lack of structures indicative of transport, the presence of well preserved non-abraded fossils, patchy distributions of rock types in contrast to down-ramp gradients of depth-sensitive assemblages, common occurrence of fossils in life position, geopetals in undisturbed position, and beds traceable into and through mud mounds formed on the deeper ramp. The biotic components (hexactinellid sponge spicules, crinoid and other echinoderm debris, fenestellid bryozoan hash, and ostracods) are used to define assemblages that vary in relation to their shoreward and basinward ramp position. Assemblage I of the deepest water consists of crinoid/echinoderm debris, fenestellid hash and ostracods. Assemblage II contains the biota of assemblage II plus stacheiins (possible algae), rare salebrids (possible bryozoans) and commonly abundant sponge spicules. Assemblage III is characterized by members of assemblage II plus plurilocular foraminifera. The shallowest assemblage IV consists of assemblage III plus green algae. Within any bed, assemblage I is farthest down-ramp (south) and assemblages III or IV are the farthest up-ramp (north). The positions of the assemblages on the ramp shift laterally in successive beds of the stratigraphic sections, indicating variations in sea level and water depth. The beds 1–6 contain the deepest and deep assemblages I, II and III; in the middle beds 7, 8, 9, 10 and 12 the assemblages shift up-ramp (from right to left of the figure), with the greatest shifts in bed 8. This bed represents the maximum flooding interval within the aggradational sediments of the Alamogordo Member. In beds 11, 13 and 14 the assemblages shift back down-ramp, leaving space for the shallowest assemblage IV in the north. The beds below bed 8 represent a part of the transgressive system tract, the beds above bed 8 document the initial highstand tract. Modified from Jeffery and Stanton (1996).
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14.2.2 Case Study: Foraminiferal Distribution in Late Triassic Reefs and Platforms Distributional patterns of benthic foraminifera evaluated in thin sections of shelf carbonates are important tools in subdividing facies zones of platform, ramp and reef environments, particularly in Mesozoic and Early Tertiary sediments.
Case Study: Triassic Foraminifera
Benthic foraminifera are common constituents of Late Triassic platform and reef carbonates. They have proved to be extremely useful in recognizing facies zones. Open and restricted platforms (Facies Zones FZ 7 and FZ 8), platform margins and foreslope environments (FZ 6, 5 and 4) as well as individual parts of reef complexes can be differentiated and subdivided by distinct foraminiferal assemblages. The concept was
Fig. 14.13. Distributional patterns of benthic foraminifera are successfully used in recognizing and subdividing facies zones and environments of carbonate platform and reef complexes. The example from the Rötelwand near Adnet (Austria) shows the distribution of foraminiferal assemblages in a small patch reef formed on a swell within a muddy shallow intraplatform basin. The facies zones of these Late Triassic limestones are defined by dominating grain types, biota and limestone textures. The central reef facies is characterized by coralline sponges, scleractinian corals and interstitial reef debris. Note the asymmetric facies pattern, comprising bioclastic grainstones of a near-reef oncolithic grapestone facies and an algal-foraminiferal facies at the leeward side, and a mud facies with reef detritus (packstones, wackestones) grading into a marly lime mud facies (wackestones, mudstones) surrounding the reef. The facies zones are distinguished by specific taxa, assemblage composition and relative frequency of taxa. Some of the diagnostic foraminifera are depicted in Pl. 111. Length of the section about 1.5 km. Modified from Schäfer and Senowbari-Daryan (1981).
Basics: Biotic Zonation Patterns
developed for the Northern Calcareous Alps. Studies in other parts of the Tethys (Carpathians, Sicily, Greece, Oman, Papua New Guinea) confirmed the general usefulness of the method in facies studies of outcrop and subsurface samples. Fig. 14.13 and Pl. 111 show the potential of foraminiferal patterns in distinguishing reef environments and adjacent areas. Tethyan Norian and Rhaetian platform and reef carbonates are characterized by specific foraminiferal assemblages. The distribution was controlled by substrate conditions and salinity, nutrient requirements, specific life behavior, the topography of the depositional environment, water depth, and light. The facies-diagnostic foraminifera shown in Pl. 111 are members of various high systematic units (see Sect. 10.2.2.1 for criteria). They include Textulariina characterized by shells with agglutinated walls (Pl. 111/1, 9, 10, 13, 14, 18, 20), Miliolina with porcelaneous walls (Pl. 111/2-8), Involutinina with hyaline aragonitic walls (Pl. 111/15, 16, 17, 19, 22), Rotaliina with hyaline walls Pl. 111/ 23) as well as Fusulinina with microgranular walls (Pl. 111/11, 12, 21, 24). The application of the method requires the following steps: • Quantify the essential microfacies criteria (grain types, fossils, micrite, intergranular cement). • Determine the foraminifera on a genus level, if possible on a species level. • Examine the frequency of the foraminifera by pointcounting, for both the whole sample and the skeletal grains • Check the number of foraminiferal tests representing different genera and species within an area of nine square centimeters. This area was successfully used the studies of Triassic foraminifera. The shell size and density of the shells in thin sections require relatively large fields. • Evaluate the significance of the composition of foraminiferal assemblages using chi-square tests, cluster or ordination analysis (Hohenegger 1974; Dullo 1980)! • Compare your results with the distribution patterns described by Hohenegger and Piller (1975), Piller (1978), Schäfer and Senowbari-Daryan (1981), Gazdzicki (1983) and other relevant papers listed under the Code K111! Basics: Biotic zonation patterns Geel, T. (2000): Recognition of stratigraphic sequences in carbonate platform and slope deposits: empirical models
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based on microfacies analysis of Paleogene deposits in southeastern Spain. – Palaeogeography, Palaeoclimatology, Palaeoecology, 155, 211-238 Hallock, P., Glenn, E.C. (1986): Larger foraminifera: a tool for palaeoenvironmental analysis for Cenozoic depositional carbonate facies. – Palaios, 1, 55-64 Hladil, J. (1986): Trends in the development and cyclic patterns of Middle and Upper Devonian buildups. – Facies, 15, 1-34 Hohenegger, J. (1999): Larger foraminifera - microscopical greenhouses indicating shallow-water tropical and subtropical environments in the present and past. – Occasional Papers of the Kagoshima University Research Center for the Pacific Islands, 32, 19-45 Hohenegger, J., Lobitzer, H. (1971): Die ForaminiferenVerteilung in einem obertriadischen KarbonatplattformBecken-Komplex der östlichen Kalkalpen. – Verhandlungen der Geologischen Bundesanstalt Wien, 1971, 458-485 Hohenegger, J., Piller, W. (1975): Ökologie und systematische Stellung der Foraminiferen im gebankten Dachsteinkalk (Obertrias) des nördlichen Toten Gebirges (Oberösterreich). – Palaeogeography, Palaeoclimatology, Palaeoecology, 18, 241-276 Jeffery, D.L., Stanton, R.J. (1996): Biotic gradients on a homoclinal ramp: the Alamogordo Member of the Lake Valley Formation, Lower Mississippian, New Mexico, U.S.A. – In: Strogen, P., Sommerville, I.D., Jones, G. L. (eds.): Recent advances in Lower Carboniferous Geology. – Geological Society London, Special Publication, 107, 111-126 Kumar, A., Saraswati, P.K. (1997): Response of larger foraminifera to mixed carbonate - siliciclastic environments: an example from the Oligocene - Miocene sequence of Kutch, India. – Palaeogeography, Palaeoclimatology, Palaeoecology, 136, 53-65 Machel, H.G., Hunter, I.G. (1994): Facies models for Middle to Late Devonian shallow-marine carbonates, with comparisons to modern reefs: a guide for facies analysis. – Facies, 30, 155-176 Moody, R.T.J. (1987): The Ypresian carbonates of Tunisia – a model of foraminiferal facies distribution. – In: Hart, M.B. (ed.): Micropalaeontology of carbonate environments. – 82-92, Chichester (Harwood) Murray, J.W. (1991): Ecology and palaeocology of benthic foraminifera. – 397 pp., New York (Longman) Pautal, L. (1987): Foraminiferal assemblages of some early Tertiary environments (bays) of the northern Corbières, France. – In: Hart, M.B. (ed.): Micropalaeontology of carbonate environments. – 74-81, Chichester (Harwood) Piller, W. (1978): Involutinacea (Foraminifera) der Trias und des Lias. – Beiträge zur Paläontologie von Österreich, 5, 1-164 Schäfer, P., Senowbari-Daryan, B. (1981): Facies development and paleoecologic zonation of four Upper Triassic patch reefs, Northern Calcareous Alps near Salzburg, Austria. – In: Toomey, D.F. (ed.): European reef models. – Society of Economic Paleontologists and Mineralogists, Special Publications, 30, 241-259 Venec-Peyre, M.T. (1991): Distribution of living benthic foraminifera on the back-reef and outer slopes of high island (Moorea, French Polynesia). – Coral Reefs, 9, 193-203 Further reading: K111
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Biotic Zonation: Triassic Foraminifera
Plate 111 Biotic Zonation of Shallow-Marine Facies Zones with Benthic Foraminifera: Late Triassic Platform and Reef Limestones Reef: Foraminifera of high-growing coral communities 1 Alpinophragmium perforatum Flügel. The species, known all over the Tethys and from Panthalassan terranes, is characteristic for the central part of Norian and Rhaetian reefs. The fine-agglutinated foraminifer encrusts high-growing corals. Reef limestone, coral framestone. Triassic (Misfah Formation, Hawasina Group, Norian): Oman Mountains. Reef: Foraminifera living in millimeter- to centimeter-sized growth cavities and on low-growing reefbuilders (e.g. sphinctozoid sponges, solenoporacean red algae) 2 Typical reef cavity microfacies with miliolinid foraminifera (1: Ophthalmidium carinatum (Leischner); 2: Galeanella) and peloids. Late Rhaetian patch reef: Feichtenstein, Salzburg, Austria. 3 Reef cavity with Galeanella (1: Galeanella sp.; 2: Galeanella panticae Brönnimann et al.) and micrite clasts. This microfacies occurs in small patch reefs at the margins of intraplatform basins as well as in large platform-marginal reefs. Same locality as –> 2. 4–5 Galeanella panticae Brönnimann et al. This miliolinid foraminifer is abundant in protected parts within central reef parts and occurs also in bioclastic sands deposited in the leeward position of Late Triassic reefs (Facies Zone FZ 7). Norian (Dachstein reef limestone): Gosaukamm, Austria. 6 ‘Sigmoilina’, a common miliolinid foraminifer in the central reef facies. Same locality as –> 4. 7 Ophthalmidium sp. Common in protected areas between reefbuilders and in mud-rich reefs formed at the upper slope (Facies Zone FZ 4). Same locality as –> 4. 8 Ophthalmidium triadicum (Kristan). Oblique section showing the characteristic trumpet-shaped apertura. Common in boundstones but also in muddy platform areas with reef-derived material (Facies Zone FZ 7). Same locality as –> 5. 9 ‘Lituosepta’ sp. Common in reef cavitiy sediments and in sediments between reef patches. Rhaetian: Same locality as –> 2. 10 Kaeveria flügeli (Zaninetti et al.) Common agglutinated foraminifera in protected reef areas. Same locality as –> 4. Back-reef and near-reef carbonate platform 11 Duostominidae. These foraminifera, characterized by trochospiral tests with microgranular walls, are concentrated in platform carbonate sands (Facies Zone FZ 7). They are associated with involutinid foraminifera and dasyclad green algae. Norian (Bedded Dachstein limestone): Dachstein area, Austria. 12 ‘Tetrataxis‘ nanus Kristan-Tollmann. The genus is most common in platform environments, but also occurs in muddy shelf limestones (Facies Zone FZ 8). Rhaetian (Kössen beds): Steinplatte, Tyrol, Austria. 13 Glomospira and Glomospirella are difficult to discriminate in thin sections. The genera are characterized by a streptospirally enrolled second chamber that later becomes planspirally coiled, forming a discoidal test. Glomospirella has been regarded as agglutinated foraminifera but has now been assigned to the involutinids. Common in platform and backreef carbonates (FZ 7), associated with other involutinids and dasyclads. Same locality as –> 4. 14 Biserial agglutinated foraminifera, attributed to Textularia, are more common in platform sediments than in the reef facies. Same locality as –> 12. Platform environments 15 Triasina grainstone. Triasina Majzon is a widely distributed involutinid foraminifer in open-marine far-reef platform carbonates (FZ 7). Note the different preservation of the primarily aragonitic tests. Late Triassic (Rhaetian): Extraclast within an Upper Jurassic breccia. Gschöllkopf, Sonnwendgebirge, Northern Alps, Austria. 16 Aulotortus communis (Kristan-Tollmann), an abundant involutinid foraminifera in grainstones of Late Triassic platforms (FZ 7). Norian (bedded Dachstein limestone): Gosaukamm, Austria. 17 Aulotortus sinuosus Weynschenk. Subaxial-axial section. The species occurs in backreef sands as well as in platform calcarenites and shallow basins. Note the characteristic sickle-shaped chambers. Rhaetian (Kössen beds): Loferbachgraben, Salzburg, Austria. Deeper shelf carbonates 18 Encrusting textularinid foraminifera. Tolypammina gregaria Wendt. Various attached foraminifera are common constituents of peloidal muddy carbonates deposited in deeper areas in front of reefs (FZ 4) or in low-energy platform areas (FZ 8). The foraminifer encrusts bioclasts and acts as binder of sediment grains, sometimes forming oncoids. Rhaetian: Ehrenreit, Bavaria. 19 Planiinvolutina carinata Leischner. Attached, probably miliolinid foraminifera. Intraplatform shelf limestone. Same locality as –> 14. 20 Trochammina alpina Kristan. The trochospiral tests of this agglutinated foraminifer is most common in micritic limestones (deeper shelf carbonates and protected micritic parts of the inner platform (FZ 8), mud-rich reefs of the upper slope, FZ 4). Rhaetian: Gruber reef near Hallein, Austria. 21 ‘Endothyranella’. Muddy shallow-basinal limestones. Rhaetian (Kössen beds): Hochalm, Tyrol, Austria. Far-reef carbonates 22 Agathammina austroalpina Kristan-Tollmann and Tollmann. This miliolinid or finely agglutinated textulariniid foraminifer is abundant in far-reef peloidal muddy settings (FZ 8) and in open-marine inner platform areas (FZ 7). Same locality as –> 17. 23 Austrocolomia marschalli Leischner, a lagenid foraminifer. Note the gradually enlarging chambers and wall with costae. Rhaetian (Kössen beds): Gaissau, Salzburg, Austria. 24 ‘Turrispirillina’ sp. Muddy slope (FZ 4) and basinal facies. Norian (Plattenkalk): Gosaukamm, Austria.
Plate 111: Late Triassic Foraminifera
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14.3 Standard Microfacies Types (SMF) Facies of carbonate rocks is studied on different scales. Stratigraphic relationships of rock bodies, sedimentary structures, and litho- and biofacies are the main targets of outcrop studies. In the subsurface, rock bodies and facies units are seismcally differentiated, using log characteristics and investigating cores and cuttings. Microfacies based on thin-section studies subdivides facies into units of similar compositional aspect that reflect specific depositional environments and controls. This target requires integrating small-scale microfacies data and larger-scale litho- and biofacies criteria. This can be done when the criteria of textures, composition and fossils of the limestones are considered jointly with the help of Standard Microfacies Types. Standard Microfacies Types are virtual categories that summarize microfacies with identical criteria. These criteria are simple, non- or semi-quantitative, and easy to recognize. Most SMF Types are based on only a few dominant characteristics comprising grain types, biota or depositional textures. The SMF concept arose from the recognition of the astonishing compositional and textural similarity of limestones of different age formed in analogous environments (see Fig. 14.28). Originally developed for categorizing common Late Triassic platform and reef carbonates, and based on the combination of texture and paleontological criteria (Flügel 1972), the classification was expanded and more strictly defined by Wilson (1975) for the history of carbonate facies over time. Wilson distinguished 24 SMF Types and used these types as additional criteria in differentiating the major facies belts of an idealized rimmed carbonate shelf (Fig. 14.29). Which criteria are used to differentiate SMF Types? The main criteria used in differentiating SMFs comprise: • Grain types, grain frequency (prevailing grain types) and grain associations. • Matrix types (e.g. micrite, calcisiltite). • Depositional fabrics (e.g. lamination, grading, openspace structures, burrowing, reworking, and redeposition). • Fossils (dominant groups, assemblages, autochthonous/allochthonous occurrence, skeletal concentrations e.g. shell accumulations, specific ecologic groups e.g. planktonic organisms, taphonomic features e.g. preservation; micrite envelopes).
Standard Microfacies Types
• Depositional texture types. Only a few SMF Types correspond to just one texture type; most SMF Types occur in two or more texture types of the Dunham classification and its modifications (Box 14.7). Some SMF Types are very simply defined and rely on the principal and quantitatively dominant kind of grains. For example, limestones with abundant aggregate grains represent a specific SMF Type (Pl. 121/4). Limestones with abundant ooids, or limestones with abundant peloids, or with dominating benthic foraminifera are attributed to three other SMF Types (see Pl. 120, Pl. 121/1-3, and Pl. 122/1). These types are easy to determine but somewhat dangerous in their use as indicators of paleoenvironmental conditions. A thorough inspection of oolitic, peloidal or foraminiferal limestones will reveal structural, textural or compositional differences that must be included in the discrimination of SMF Types, because these differences contain essential information on depositional settings and environmental controls. Oolite grainstones consisting of abundant and well-sorted concentric ooids indicate shallow-marine, high-energy conditions. Grainstones and wackestones with poorly sorted radial ooids indicate low-energy conditions, common in transitionalmarine and also non-marine environments. In the following revision of the SMF Type classification these differences are annotated by the introduction of subtypes (e.g. SMF 15-C and SMF 15-R for oolite limestones) in order to avoid an inflation of SMF types. Subtypes were also proposed for peloid limestones, using differences in the fabrics, and for limestones consisting of foraminifera or calcareous green algae in rockbuilding abundance. Other limestones having in common the same dominant grain types (e.g. oncoids) are attributed to different SMF Types because of unlike carbonate textures and different grain structures, indicating differences in environment and depositional setting (cf. Pl. 119/3 and Pl. 124/2). Some SMF Types are defined by characteristic ecologic groups, e.g. the abundance of planktonic microfossils (Pl. 113), or the occurrence of reef-building fossils (Pl. 116/2, 3). These categorizations are very broad and include many pitfalls. Pelagic limestones with abundant planktonic fossils must be subdivided at least according to the prevailing systematic groups. The use of ‘reef-building fossils’ as diagnostic criteria requires interpretating the reef- and sediment-building potential of benthic organisms. Thin sections with corals are no evidence for the existence of reefs. Corals may be members of the constructor guild and contribute to organic frameworks, but are also bafflers that might or
Standard Microfacies Types
might not take part in reefbuilding, and they are also members of the dweller guild of level-bottom communities. Despite these difficulties, the SMF Types based on ecologic groups have been retained, since they allow a first approximate assignment of samples. Other criteria used in defining SMF Types are specific fabrics, e.g. lamination or fenestral structures (Pl. 123), or combinations of diagenetic and depositional features (laminated evaporitic carbonates, Pl.125/4; pisolitic carbonates, Pl. 126/1, 2). Can the SMF Types be used for other carbonate depositional models? The SMF Types were defined for a model describing the sedimentation on a rimmed carbonate shelf and warm-water platform-reef environments in tropical latitudes. The occurrence and distribution of the SMF Types are strongly dependent on water depth, and that in turn depends on the configuration and topography of the shelf. Some microfacies types of inner and shallow parts of carbonate ramps correspond to SMF Types of platform interior sediments. Back-ramp tidal and lagoonal areas are similar in microfacies associations to tidal flat and lagoonal environments of platforms. Other microfacies types on ramps and platforms, however, differ in their location. SMF 15, characterized by concentric ooids, occurs in the platform model in a platform-edge position, in the ramp model in an inner and mid-ramp position. Outer ramp microfacies types include SMF Types known from the Wilson model, but also types not specifically distinguished in the platform model. Similar situations exist in comparing microfacies types of non-rimmed platforms and epeiric platforms and ramps with SMF Types of rimmed platforms. Major discrepancies exist between the Standard Microfacies Types and microfacies types of cold-water carbonates. The latter differ in composition (lack of ooids, aggregate grains), skeletal grain associations and skeletal mineralogy (see Sect. 12.2), texture and weak cementation. Many temperate and cool-water microfacies types hardly correlate with SMF Types at all. Look at the pictures of cool-water limestones on Pl. 106/1–3 and try to find comparable microfacies types in the SMF classification that were set up for warmwater platform carbonates! The only SMF Type with an approximately similar biotic composition is SMF 9. This SMF is defined by strongly burrowed skeletal wackestones and a wide spectrum of fossil groups also comprising the fossils of these cool-water samples. Because of the too generalized definition of SMF 9, this assignment provides little help in interpreting the environment of the samples. A comparison with microfacies types of ramps is much more advisable.
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14.3.1 Revised Standard Microfacies Types The following pages summarize the diagnostic criteria of the Standard Microfacies Types and indicate their distribution in the Facies Zones of the Wilson model and in the different part of carbonate ramps. Let us start with a short list (Box 14.4).
Box 14.4. List of Standard Microfacies Types. The order of the SMF numbers follows approximately the order of the Standard Facies Zones in the Wilson Model going from the basinal SMF Type 1 to SMF 26 that characterizes subaerial exposed areas. SMF 1: Spiculitic wackestone or packstone, often with calcisiltite matrix. Subtype emphasizes burrowing. SMF 2: Microbioclastic peloidal calcisiltite with fine grainstone and packstone fabrics. SMF 3: Pelagic lime mudstone and wackestones with abundant pelagic microfossils. Subtypes differentiate the groups of planktonic organisms. SMF 4: Microbreccia, bio- and lithoclastic packstone or rudstone. SMF 5: Allochthonous bioclastic grainstone, rudstone, packstone, floatstone, breccia with reef-derived biota. SMF 6: Densely packed reef rudstone. SMF 7: Organic boundstone. Subtypes try to differentiate the kind of contribution by potential reefbuilders to the formation of reefs and other buildups. SMF 8: Wackestones and floatstones with whole fossils and well-preserved endo- and epibiota. SMF 9: Strongly burrowed bioclastic wackestone. SMF 10: Bioclastic packstone and wackestone with abraded and worn skeletal grains. SMF 11: Coated bioclastic grainstone. SMF 12: Limestone with shell concentrations. Subtypes characterize shell-providing fossils. SMF 13: Oncoid rudstone and grainstone. SMF 14: Lag deposit. SMF 15: Oolite, commonly grainstone but also wackestone. Subtypes highlight the structure of ooids. SMF 16: Peloid grainstone and packstone. Subtypes differentiate non-laminated and laminated rocks. SMF 17: Grainstone with aggregate grains (grapestones). SMF 18: Bioclastic grainstone and packstone with abundant and rock-building benthic foraminifera or calcareous green algae. Subtypes describe the systematic assignment of the various groups. SMF 19: Densely laminated bindstone. SMF 20: Laminated stromatolitic bindstone/boundstone. SMF 21: Fenestral packstone and bindstone. Subtypes characterize fenestral voids and the contribution of calcimicrobes. SMF 22: Oncoid floatstone and wackestone. SMF 23: Non-laminated homogenous micrite or microsparite without fossils. SMF 24: Lithoclastic floatstone, rudstone or breccia. SMF 25: Laminated evaporite-carbonate mudstone. SMF 26: Pisoid cementstone, rudstone or packstone. Text continued on p. 712
Standard Microfacies Types
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Plate 112 Standard Microfacies Types: SMF 1 and SMF 2 This plate and the following plates up to Pl. 126 start with a general text explaining the diagnostic criteria and the occurrence of the SMF Type in the different depositional settings. The description of thin sections is focused on characteristics of the SMF Types and indicates their more frequent distribution in generalized facies zones. SMF 1: Spiculite wackestone or packstone, often with a calcisiltite matrix Criteria: Dark, bedded limestones (mudstone, wackestone), commonly argillaceous and organic-rich. Abundant siliceous (or calcified) sponge spicules, often oriented. Parallel orientation of spicules and clay laminae may cause a very fine depositional ‘lamination’ pattern (Wilson 1969). In addition pelagic microfossils (e.g. radiolarians) may occur. The microfossils may be densely packed (–> 1) or sparsely distributed due to intensive burrowing (SMF 1-BURROWED: –> 2). Occurrence: Basinal deep-water environment with slow sedimentation. This microfacies is common in Paleozoic and Mesozoic basinal carbonates (FZ 1), deeper shelf carbonates (FZ 2) as well as in mid-ramp and outer ramp settings. Spiculites are common in deep cold-water settings, but also occur in shallow-marine environments (Gammon and James 2001). Additional pictures: Pl. 44/1, Pl. 139/1. SMF 1-BURROWED: Burrowed bioclastic wackestone with abundant fine pelagic and benthic biodetritus. Criteria: Very small bioclasts, commonly shell debris, scattered within a dense, strongly burrowed matrix. SMF 1-BURROWED differs from SMF 1 in the dominance of sparsely distributed skeletal grains representing a mixture of benthic and planktonic elements. Occurrence: Basin (FZ 1), open sea shelf (FZ 2), and outer ramp. SMF 2: Microbioclastic peloidal calcisiltite. Criteria: Thin- to medium bedded limestones. Very fine-grained packstones or grainstones consisting of peloids and small litho- and bioclasts. The grains are a mixture of off-bank debris and reworked sediment. Common bioclasts are echinoderms and mollusks. Small-scale ripple cross-lamination and laminae exhibiting millimeter-scale gradation are common. Occurrence: Common in deeper basins (FZ 1) and open-marine shelf (FZ 2) as well as in deep-shelf toe-of-slope position (FZ 3). Abundant in outer ramp settings. 1 SMF 1. Radiolaria-bearing spiculite packstone. The bioclastic packstone consists predominantly of accumulations of monaxon and other hexactinellid sponge spicules. Some of the spicules are partly calcified. Larger circular structures are cross-sections of radiolarians (arrows). Note parallel orientation of spicules indicating bottom currents. The matrix is organic-rich micrite. Water energy was sufficient to orientate the spicules, but not to winnow the carbonate mud. Deep basin (FZ 1A). Jurassic of the Northern Calcareous Alps (Roman mosaic stone): Kraiburg, southern Bavaria, Germany. 2 SMF 1-BURROWED. Burrowed wackestone. Skeletal grains are conspicuously small and comprise shell debris, sponge spicules (SP), radiolarians, and small echinoderm fragments (E). Burrowing (B) is abundant, corresponding to the bioturbation grade 4 of Taylor and Goldring (1993). This index describes burrow abundance, amount of burrow overlap and sharpness of primary sedimentary fabric. Grade 4 is characterized by high burrow density with common overlap. Burrowing contributes to the diminution of skeletal grains. Deep open shelf (FZ 2). Early Jurassic (Sinemurian): Korfu Island, Greece. 3 SMF 2. Fine-grained peloid grainstone (FPG) and peloid-bioclast packstone (PBP). Most bioclasts are echinoderms with syntaxial rim cements. The sediment consists of peloids and skeletal debris. White arrow points to a shell fragment, the black arrow to serpulid tubes. The currents that deposited the silt-sized grains also winnowed most of the lime mud. Off-bank facies of an outer ramp. Cretaceous: Switzerland. Fig. 14.14. SMF 2: Cross-bedded fine-grained quartz-peloid packstone. Dark layers consist of densely packed micritic peloids (redeposited intraclasts). Light layers contain angular, silt-sized terrigenous quartz grains. Note the brachiopod shell (arrow). The sample represents a basinal facies. The sediment was deposited in deep-water canyons (see Buchroithner et al. 1980). Note the vertical calcite-filled microcracks displacing an earlier oblique calcite vein. Standard Facies Zone (FZ 2). Late Devonian: Ferragut Vell, Menorca Island, Spain. Scale is 5 mm.
Plate 112: SMF 1 and SMF 2
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Standard Microfacies Types
Plate 113 Standard Microfacies Type: SMF 3 SMF 3: Pelagic lime mudstone and wackestone with planktonic microfossils Criteria: Micrite matrix with common to abundant pelagic microfossils (e.g. foraminifera, radiolarians, calpionellids) or macrofossils (e.g. ‘filaments’ – Triassic to Jurassic, pelagic crinoids, graptolites). As shown by SEM photographs, the micritic matrix of many Mesozoic limestones consists of accumulations of disintegrated calcareous ultraplankton (e.g. coccolithophorids, Pl. 7/3; nannoconids, Pl. 77/23). The type of the dominant pelagic fossils should be indicated by letters e.g. SMF 3-CALP (calpionellids), SMF 3 -F OR (pelagic foraminifera), SMF 3-RAD (radiolarians), SMF 3-TENT (tentaculitids). Occurrence: Basin (FZ 1-B) and open deep shelf (FZ 3). Additional pictures: Pl. 6/6, Pl. 66/5, Pl. 73/3 (SMF 3-FOR), Pl. 76/1 (SMF 3-RAD), Pl. 76/7 (SMF 3-CALP), Pl. 91/3 (SMF 3-TENT). 1 SMF 3-FOR. Pelagic foraminifera wackestone with single-keeled globotruncanids. Tests are generally well-preserved due to the primary Low-Mg calcite skeletal mineralogy. Pelagic deep-water basinal facies (FZ 1). Late Cretaceous (Scaglia rossa): Central Apennines, Italy. 2 SMF 3-FIL. Thin-shelled pelagic bivalve (‘filaments’) wackestone. The thin, curved calcitic shells were originally regarded as planktonic algae and therefore called ‘filaments’ (see Sect. 10.2.4.2). The fossils are pelagic bivalves representing in parts shells of planktonic larval stages. Note the parallel arrangement of the shells pointing to some current control. Dense accumulations (arrows) are caused by burrowing. The biofabric is characterized by dense and loose packing; little variations in the sizes of the bioclasts, and good sorting. The paleobathymetric position of limestones with filaments is controversial. Many authors regard the facies as an indication of bathyal deep-water settings, but others go for a littoral environment. Deep shelf facies (FZ 3). Early Jurassic: Korfu Island, Greece. 3 SMF 3-CALP. Calpionellid wackestone. Arrows point to the few sections which can be used in taxonomic differentiation. The fine-grained matrix consists predominantly of nannofossils and micropeloids better to be seen in SEM. Black patches are bituminous matter. Deep basinal facies (FZ 1). Late Jurassic (Tithonian): Seebühel, Switzerland.
Fig. 14.15. SMF 3-TENT. Densely packed tentaculitid wackestone. All skeletal grains are shells of pelagic mollusks. Tentaculitids contributed largely to the formation of Ordovician to Late Devonian open-marine and basinal micritic carbonates (see Sect. 10.2.4.5). Deep basin (FZ 3A). Early Devonian: Frankenwald, Germany.
Plate 113: SMF 3
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Plate 114 Standard Microfacies Types: SMF 4 and SMF 5 SMF 4. Microbreccia, bioclastic-lithoclastic packstone or rudstone Criteria: Fine-grained breccias, debrites and turbidites consisting of grains of various origin. Grains are commonly worn and rounded and may consist of both locally derived bioclasts, imported shallow-water material and previously cemented lithoclasts. Grains may be either polymict in origin or of uniform composition. Grading textures are common. Grains can be mixed with quartz, cherts and extraclasts. This rock type includes both fine talus and also debris laid down by turbidites (e.g. in ‘allodapic limestones’, Meischner 1965). Occurrence: Foreslope talus, resedimented limestones. Basinal settings (FZ 1) and toe-of-slope settings (FZ 3). SMF 5. Allochthonous bioclastic grainstone, rudstone, packstone and floatstone or breccia Criteria: Densely packed whole fossils and fossil fragments with a high percentage of reef-derived organisms. The bioclasts occur in a chaotic order or they form depositional structures with characteristic patterns (e.g. turbidite sequences). Occurrence: Reef-flank facies, slope (FZ 4), adjacent to reefs. Deposited in forereef position and reef slopes, or in backreef settings and in lagoons. Inner ramps. Additional picture: Pl. 58/1. 1 SMF 4. Polymict clast-supported microbreccia, consisting of carbonate litho- and bioclasts and quartz grains (most white particles). Carbonate grains are subangular to subrounded (fields G, I and L of the Krumbein chart; see Fig. 4.30). Most quartz grains correspond to the fields P and Q (low sphericity and low roundness). Carbonate grains are lithoclasts (L) and bioclasts including bivalves (B), bryozoans (BR), echinoderms (E), foraminifera (F), finely agglutinated orbitolinid foraminifera (OF), and rudist bivalves (R). Debris flow. Toe-of-slope (FZ 3). Late Cretaceous: Northern Calcareous Alps. 2 SMF 4, SMF 2 and SMF 1. Allochthonous and autochthonous deep-water sedimentation. The coarse-grained lower part of the basal unit yields thin brachiopod shells (B) deposited in ‘convex-up’ positions. The center part (packstone, PS) consists of micrite clasts and thin, parallel arranged shells. The overlying layer is a fine-grained grainstone (GS) deposited within a relief (arrows) of the underlying packstone. Packstone and grainstone display indistinct lamination caused by parallel arranged thin shells. Note the differences in grain sizes and the sharp boundary between the fine-grained grainstone and the overlying brachiopod-bearing lime mudstone (M). Shell orientation patterns indicate the existence of weak currents. The sample exhibits three SMF Types: The basal layer with brachiopod shells represents SMF 4. The packstone and grainstone layers correspond to SMF 2. The uppermost part reflects the basinal background sedimentation (lime mudstone with scattered fossils, SMF 1). Sedimentation took place in a deep shelf environment (FZ 3), see Fig. 14.16. Uppermost Silurian (Lochkovian): Klonk near Suchomasty, southwest of Prague, Czech Republic. 3 SMF 5. Phylloid algal floatstone. Characteristic criteria of SMF 5 are the abundance of reef-derived fossils, the chaotic fabric and the packstone texture of the matrix. Most bioclasts of the sample shown here are fragmented plates of udoteacean green algae (see Sect. 10.2.1.5). Some fragments are coated and encrusted by foraminifera. The plates are jumbled together and form a self-supporting framework that was infilled by the matrix. Shells (S) of gastropods and bivalves and fusulinid foraminifera (F) are of minor importance. The matrix is an organic-rich packstone consisting of silt-sized peloids and some quartz grains. The sample comes from flank beds of a shelf-margin mound that was subaerially exposed several times. The mound material was redeposited near the mound core. The algal plates are strongly leached and partly filled with gray micritic sediment (black arrow). The original algal structures are completely destroyed. Many phylloid algal buildups of the southwestern U.S.A. are excellent reservoir rocks. Porosity developed through early vadose leaching of skeletal grains and dissolution of parts of the matrix, combined with fracturing and brecciation. Late Pennsylvanian (Holder Formation): Yucca Mound, northern Sacramento Mountains, New Mexico, U.S.A. –> 2: Hladil 1992 Fig. 14.16. Microfacies assists in differentiating allochthonous and background sedimentation in deep-marine environments. The sample on Pl. 114/2 comes from the international stratotype of the Silurian-Devonian boundary. Questions concerning a possible turbiditic origin of the boundary bed have been clarified by a microfacies study (Hladil 1992). Common microfacies types are hemipelagic laminites, currentderived laminites and lime shale and muddy limestones with graptolites. Coarse calciturbidites are rare. The microfacies of bed 14 exhibits criteria of currentcontrolled deposition (packstones and grainstones; SMF 2), deposition of argillaceous and calcareous material (mudstone, SMF 1), and only minor allochthonous sedimentation (brachiopod accumulation at the base of the bed, SMF 4). Most of the other limestone beds of the Klonk section exhibit similar microfacies composition. Modified from Hladil (1992). A: Laminite; B: Current-generated limestone; C: Coarse turbidite; D: Lime shale and muddy limestones with graptolites.
Plate 114: SMF 4 and SMF 5
687
688
Standard Microfacies Types
Plate 115 Standard Microfacies Types: SMF 5 and SMF 6 SMF 5. For a general explanation of SMF 5, see Pl. 114. SMF 6. Densely packed reef rudstone Criteria: Coarse gravels of biogenic material and lithified sediment derived from ‘reef’ tops or flanks and deposited in high-energy or low-energy slope settings by rock fall and various mass-flow processes. The size of the clasts ranges from some millimeters to tens of centimeters. Often with low matrix content. Sorting is poor (debris flows) or good (turbidites). Occurrence: Slope (FZ 4). SMF 6 is common within and adjacent to reef complexes, at platform margins, or in foreslope sediments. Additional pictures: Pl. 45/3, Pl. 82/2. 1 SMF 5. Turbidite. The sample is a graded turbidite consisting of reef-derived material, and deposited in a basin adjacent to small coral reefs. Reworked grains are bioclasts including bivalve shells (BS), corals (C), echinoderms (E), gastropods (G) and algae (A). The 3.5 cm thick bed exhibits the characteristic calciturbidite sequence: An erosive base (black arrows), graded bedding consisting of a thicker floatstone layer (FS) and a thinner grainstone layer (GS), and a planar (even) top (white arrows). The bed is covered by marls (M). Shallow slope (FZ 5). Late Triassic (Zlambach Formation, Rhaetian): Northern Calcareous Alps, Austria. 2 SMF 6. Debris flow. Lithoclastic rudstone. The sample is from the lower part of a debris flow eroded from small reefs located on an upper slope (FZ 4). Note the apparently good rounding and high sphericity of both large and small micritic grains (fields C, D, M and O of the Krumbein chart; Fig. 4.30). Many grains are clasts of packstone sediment (PS) yielding reef microfossils, cemented lithoclasts (L) and isolated fossils including the microproblematicum Plexoramea (P), echinoderms (E), shells (S) and some porostromate algae (A). Lithoclasts are biogenically encrusted (arrows). Middle Triassic (Ladinian): Mahlknecht cliff, Seiser Alm, Dolomites, Italy. –> 1: Matzner 1986
Fig. 14.17. Highstand or lowstand slope deposition? Evidence from microfacies analysis: The figure shows two models describing the possible influence of sea level on the formation of allochthonous foreslope deposits in the Middle Triassic Schlern/Sciliar-Rosengarten/Catinaccio platform of the Dolomites (modified from Yose 1991). Two tongues of carbonate foreslope talus are separated by an onlapping wedge of volcaniclastic conglomerates. The carbonate megabreccias are cited as a type of allochthonous lowstand wedge (model 1; Sarg 1987). However, megabreccias may also record platform progradation and shedding of platform material during relative sea-level highstands (model 2; Schlager 1994). Slope deposits recording third-order sea-level lowstands and related subaerial erosion of platforms should be characterized by a mixture of limestone clasts derived from the upper slope, platform margin, backreef and platform interior showing evidence of subaerial diagenesis. Foreslope talus accumulating adjacent to productive platforms during third-order sea-level highstands should principally be composed of margin-derived clasts characterized by marine diagenesis. Microfacies analysis of the clasts of the lower talus tongue (black) strongly supports model 2 (Brandner et al. 1991). Clast composition and biota as well as early diagenetic marine cements within the clasts indicate reworking of platform margin/upper slope buildups (–> 2) perhaps triggered by earthquakes of volcano-tectonic origin. Inter-clast pores were occluded by burial cements.
Plate 115: SMF 5 and SMF 6
689
690
Standard Microfacies Types
Plate 116 Standard Microfacies Types: SMF 6 and SMF 7 SMF 6: For general explanation of SMF 6 see Pl. 115. SMF 7. Organic boundstones In the Standard Microfacies classification, the multifold kinds of reef limestones have been amalgamated within a single SMF Type (boundstone) characterized by in-situ organic growth of potential reefbuilders. The type was subdivided by growth forms and assumed growth patterns (framestone, bafflestone, bindstone; see Sect. 8.2). SMF 7 and also the subdivisions are of limited value for microfacies studies. ‘Reef limestones’ look very different! Compare Pl. 116/2 and Pl. 116/3, Fig. 14.18 and the reef limestones depicted in Figs. 8.2, 8.3 and Pl. 41/1 and Pl. 42. Reef limestones require specific investigation methods that are discussed in Sect. 15.2. Additional pictures: Pl. 41/1 (SMF 7-BAFFLESTONE), Pl. 42 (SMF 7-FRAMESTONE). 1 SMF 6. Mass-flow deposit. Reef rudstone. The sample shows grains eroded from ‘sponge reefs’ and redeposited on a slope. Grains are subrounded, micritic lithoclasts, reef dwellers and peloids, sometimes representing ‘tuberoids’ (interpreted as result of carbonate deposition during the decay of sponges and microbial crusts). Arrows point at longitudinal and cross-sections of Terebella lapilloides. This name designates an annelid worm tube characterized by thick agglutinated micritic walls. Terebella is a facies-diagnostic fossil in Late Jurassic sponge reefs. Originally living on the lower macerated surfaces of siliceous sponges, the tubes became transported from the reef to the near-reef detrital facies. The fossil is diagnostic of near reef deposits. Black structures with white ‘eyes’ (left margin) are reworked sessile foraminifera. Late Jurassic (Kimmeridgian): Ludwag quarry northeast Bamberg, Northern Franconian Alb, Germany. 2 SMF 7-FRAMESTONE. Coralline sponge-microbe crust framestone characterized by in-situ growth of sessile organisms. The upright growth form of the sponges (SP) and the close intergrowth with microbial crusts (MC) resulted in the formation of an organic framework. The framework exhibits a large number of cryptic habitats between and below reef-building organisms and in small reef cavities. Cavity-dwelling microfossils (foraminifera, microproblematica, microbes, sponges and others) are common criteria of this SMF Type, known since the Cambrian. Late Triassic (Dachstein limestone, Norian): Gosaukamm, Austria. 3 SMF 7-BAFFLESTONE. Coral bafflestone. Cross-section of a fasciculate coral colony (Retiophyllia) showing trapping of sediment between the delicate coral branches. Solution cavities are filled with internal sediment (arrows). Note the excellent preservation of the corals. Late Triassic (Zlambach Formation, Rhaetian): Gosaukamm, Austria. –> 2, 3: Wurm 1982
Fig. 14.18. SMF 7. Micro-framework reef limestone. Many Permian reefs were formed by a micro-framework consisting of very small, encrusting and low-growing organisms, abundant fine-grained trapped bioclastic sediment (often more than 50%) and marine carbonate cement (Weidlich 2002). Large reefbuilders are of minor or no importance in these reefs formed in platform-marginal or upper slope positions. The sample shows ‘Tubiphytes’ (T), trapped fine-bioclastic and peloidal sediment (S), and recrystallized calcite cement (C). Middle Permian (Wordian-Early Capitanian): Hawasina Complex, eastern Oman Mountains.
Plate 116: SMF 6 and SMF 7
691
692
Standard Microfacies Types
Plate 117 Standard Microfacies Types: SMF 8 and SMF 9 SMF 8. Wackestones and floatstones with whole fossils and well preserved infauna and epifauna Criteria: Predominantly sessile organisms rooted in micrite and some mobile organisms. The micrite contains scattered, often very small skeletal debris contributing to a finely bioclastic matrix. Burrowing is common. Many fossils are well preserved (e.g. double-valved mollusks), but some isolated fragments may occur. Characteristic organisms are mollusks, sponges, corals and some calcareous algae. Occurrence: Shelf lagoon with circulation. Low-energy environments below wave base (FZ 2 and FZ 7). Outer and mid-ramp settings. The sediment is formed in quiet water below the fair-weather wave base. Additional pictures: Pl. 44/3, Pl. 86/1, Pl. 88/1, Pl. 89/1. SMF 9. Strongly burrowed bioclastic wackestone Criteria: Micrite with common to abundant fragmented fossils that are jumbled through burrowing. Bioclasts often micritized. Common fossils in Paleozoic limestones are crinoids, brachiopods, bryozoans, and scattered rugose corals. Mesozoic and Tertiary examples are characterized by bivalves, gastropods, and echinoderms. Occurrence: Shallow lagoon with open circulation at or just below the fair-weather wave base (FZ 7) and deep shelf (FZ 2); outer and mid ramp settings. Additonal pictures: Pl. 19/1, Pl. 44/2. 1 SMF 8. Whole fossil wackestone/floatstone. The sample shows solenoporacean red algae (SA, Parachaetetes), corals (C) and a few dasyclad algae (DA, Jodotella), as well as foraminifera (F) and bored shells (S). Most fossils are well preserved and show no signs of reworking. This SMF is common both in the deep open shelf environment (FZ 2) and in shallow open shelf lagoons (FZ 7). The presence of calcareous algae supports the attribution to FZ 7. Early Tertiary (Paleocene): Kammbühl near Ternitz, Austria. 2 SMF 9. Burrowed bioclastic wackestone/floatstone. Fossils are fragments of gastropods (G), corals (C), and foraminifera (F). White arrows point at burrows filled with fecal pellets and foraminifera. The interpretation as fecal pellets is supported by the dark color and the marked bimodal grain size. The matrix is a fine-grained pelmicrite, in parts pelsparite. The stratigraphic sequence reflects a tropical lagoon surrounded by coral reefs. Thin-bedded, platy bituminous micrites without benthic fossils but rich in fishes represent the lagoonal center. Bedded limestones with larger foraminifera and mollusks (this sample) represent the shallow part of the lagoon. The composition of the gastropod fauna reflects changes from open-marine to more restricted conditions. Early Tertiary (Eocene): Pesciara, Monte Bolca, Verona, northern Italy. Fig. 14.19. SMF 8: Bioclastic wackestones and floatstones. Many samples of the SMF Type 8 are characterized by bioclastic wackestones and floatstones with various shells (A, B, C) and/or by nearly monospecific biota (D). An other common criterion is the fine-bioclastic micrite matrix. A: Bivalve-gastropod wackestone. Normal marine, low-energy, shallow subtidal environment. Facies zone FZ 7. Late Cretaceous (Turonian): Sinai, Egypt. B: Bivalve floatstone. The shells occur together with rounded micritic intraclasts (center). Note large burrow at the top right. Protected shallow subtidal environment. Mid-ramp setting (FZ 7). Locality as in A. C: Bioclastic floatstone with hyolithid shells (see Fig. 10.47) and trilobite fragments (bottom right). The matrix is a quartz-bearing burrowed pelmicrite. Deep shelf environment (FZ 2). Cambrian (Wadi Nasb Limestone Formation): Wadi Al Ghadaf well, eastern Jordan. D: Wackestone with the tabulate coral Thamnopora. The matrix is a fine-bioclastic pelmicrite. FZ 7, openmarine lagoon behind a stromatoporoid-coral reef. Late Devonian (Frasnian): Carnic Alps, Italy. Scale for all figures 2 mm. A and B: Bauer et al. (2002); C: Shinaq and Bandel (1992); D: Galli (1985).
Plate 117: SMF 8 and SMF 9
693
Standard Microfacies Types
694
Plate 118 Standard Microfacies Types: SMF 10 and SMF 11 SMF 10. Bioclastic packstones and grainstones with coated and abraded skeletal grains Criteria: Worn and coated bioclasts deposited within a fine-grained matrix. Forming packstones and wackestones, sometimes also fine-grained grainstones. This type indicates textural inversion. The dominant particles have been transported from high-energy to low-energy environments (e.g. from shoals to swales in the proximity). Occurrence: Shelf lagoon with open circulation (FZ 2) and open sea shelf (FZ 7). Common in inner and mid-ramp settings. Additional picture: Pl. 74/2. SMF 11. Coated bioclastic grainstone with sparry cement Criteria: Most grains are coated bioclasts, exhibit micrite envelopes or are completely micritized. Additional grain types may be rounded intraclasts, peloids and some ooids. Occurrence: FZ 6 (winnowed platform edge sands) and in reefs (FZ 5). Rare in inner ramp and mid-ramp settings. The sediment formed in areas with normal marine salinity, constant wave action at or above the fair-weather wave base or between the wave base and the storm wave base. Coating due to micritization by microborers occurs preferentially in very shallow environments (see Sect. 4.2.3 and Sect. 9.3.4). 1 SMF 10. Worn foraminiferal (fusulinid) packstone. The foraminiferal shells (Triticites) are worn and micritized indicating transport before deposition. Note lithoclasts (black arrows). The matrix consists of silt- to fine-sand sized peloids, micrite and sparite. The sample consists of 81% skeletal grains and 19% matrix according to point-counter analysis. Only very few fusulinids show axial sections (white arrow) necessary for species determination. The mass occurrence of foraminifera is a diagnostic criteria of SMF 18-FOR. The attribution of the sample to SMF 10 regards the textural inversion of the biofabric. Fusulinids living under high-energy conditions were transported downslope to local protected low-energy areas. Textural inversion and transport are indicated by infilling of outer whorls with sediment and by the loss of outer whorls in many tests. Open platform (FZ 7). Early Permian (Leavenworth limestone): Osage County, Oklahoma, U.S.A. The Leavenworth limestone is a uniform sheet, 500 km long and at least many tens of kilometers broad. The mid-shelf limestone unit represents the transgressive part of a megacyclothem, and comprises skeletal wackestones, sometimes associated with mudstones, oncoid wackestones, and the fusulinid packstones shown in this figure. 2 SMF 11. Coated bioclastic grainstone. Most grains are strongly micritized coated skeletal grains associated with some small-sized benthic foraminifera. The sample comes from an isolated Bahamian-type carbonate platform affected by incipient drowning. Middle Jurassic: Monte Kumeta, western Sicily, Italy. 3 SMF 11. Coated bioclastic grainstone. Most skeletal grains have micrite envelopes or biogenic encrustations and correspond to cortoids and small oncoids (black grains). Lime mud with peloids occurs as infilling in the grain-supported fabric (arrows), but most intergranular pores are occluded by calcite cement. Note the conspicuous wavy boundary between areas with micrite and small peloids and areas characterized by calcite cement. The boundary reflects differences in the lithification within the sandy sediment. Bioclastic grains are mollusk shells, dasyclad green algae (DA) and foraminifera (F). In the Late Triassic the association of involutinid foraminifera and dasyclads characterizes open-marine parts of platform interior settings (see Sect. 14.2.2). In addition to the small coated grains, larger ‘fenestral oncoids’ (ON) occur, formed by microbes (Bacinella). The sediment originated in the high-energy area of backreef sands deposited on an outer platform. Late Triassic (Dachstein limestone, Norian): Gosaukamm, Austria. –> 1: Toomey 1983; 2: Di Stefano 2002; 3: Wurm 1982
Fig. 14.20. SMF 11. Coated grainstone. Cortoids are associated with ooids and strongly micritized bioclasts. The pore space between the grains is filled with stubby fibrous rim cement (arrow) and crystal silt, indicating a post-compactional vadose overprint of the platform sediment. The sample comes from platform-marginal sand shoals adjacent to a grapestone facies. Middle Jurassic (Esfandiar limestone): Shotori Mountains, Tabas area, Iran. Crossed nicols. After Fürsich et al. (2003).
Plate 118: SMF 10 and SMF 11
695
696
Standard Microfacies Types
Plate 119 Standard Microfacies Types: SMF 12 and SMF 13 SMF 12. Limestones with shell concentrations Criteria: Bioclastic rudstones or densely packed floatstones characterized by accumulations of commonly one-type shells or echinoderm fragments. Shells may be bivalves, brachiopods, gastropods, or crinoids. The prevailing fossil group should be noted by additional letters: SMF 12-S (coquina composed of shells without specific assignment), SMF 12-BS (coquina composed of bivalve shells), SMF 12-BRACH (brachiopods), SMF 12-GASTRO (gastropods). Most echinoderm accumulations occur from crinoids (SMF 12-CRIN). Occurrence: Rock-building concentrations of shells originate in various environments, from the coast to the deep sea. Bivalve shell beds are formed in platform interior settings including restricted platforms and tidal flats (FZ 8) and open platforms (FZ 7), reefs and slopes (FZ 5 and FZ 4), and toe-of-slope and deep-marine settings (FZ 3, FZ 2 and FZ 1). Shell concentrations are very common in mid-ramp settings. Accumulations of bivalve shells are caused by various processes including current concentrations, storm wave and tempestite concentrations, transgressive lag and condensation concentrations, or prolonged continuous accumulation of shells, because sedimentation fails to keep up with hardpart accumulation (see Fig. 3.6 and Fig. 3.7). Crinoid concentrations: Limestones consisting of abundant and densely packed echinoderm fragments (e.g. ‘encrinites’) represent a specific facies type formed in various settings including slopes, protected platforms, reefs and mounds. This type has been included by Wilson in SMF 12, but should be specifically differentiated because of its paleoenvironmental significance. Be aware that crinoid packstones also may be the result of chemical compaction (indicated by clay seams, stylolites, and corroded bioclasts). Crinoid accumulations occur both in micritic limestones or they form packstones in which the crinoid columnals are cemented by syntaxial calcite. The accumulations may be autochthonous (open sea shelf - FZ 2; foreslope - FZ 4; mounds - FZ 5) or allochthonous (deep shelf margin - FZ 3; foreslope - FZ 4). Allochthonous accumulations often are debris flows or turbidites. Crinoid concentrations are common in mid-ramp settings. Additional pictures: Pl. 94/7 (SMF 12TRILO), Pl. 95/2 (SMF 12-CRIN). The original definition of SMF 12 was restricted to bioclastic rudstones or grainstones characterized by abundant shell hash, and deposited in an environment of constant wave or current action that removed carbonate mud by winnowing. Actually, mud may be removed by winnowing (bioclastic rudstones), but can be still preserved within and between the shells (bioclastic floatstone, –> 1). Winnowing, therefore, should not be regarded as a critical definition criterion. SMF 13. Oncoid rudstones and grainstones Criteria: Millimeter- to centimeter-sized oncoids, predominantly cyanoids and porostromate oncoids, forming a grain-supported fabric, sometimes in association with ooids and fine-grained bioclasts. Bimodal grain-size distributions are common. Note that this SMF does not categorize oncoid floatstones (see SMF 22!). Occurrence: Ancient oncoids were formed in different settings (see Sect. 4.2.4.1), both in high-energy and low-energy environments. Oncoid rudstones and grainstones were common in open-marine winnowed edge sand areas (FZ 6), open shelf lagoons (FZ 7), the environs of platform patch reefs, and in backreef areas behind larger reef complexes. On ramps grain-supported oncoids originated predominantly in shallow inner ramp settings. Additional picture: Pl. 12/1. 1 SMF 12-BS. Bivalve floatstone. Concentration of bivalve shells resulting from high population density on a slope (FZ 4). The absence of bioerosion, abrasion and encrustation contradicts transport. The recrystallized bivalves are double-valved (center) or occur as single valves. Matrix is a lime mudstone. Late Triassic: Steinplatte, Tyrol, Northern Alps, Austria. 2 SMF 12-CRIN. Crinoid floatstone. Crinoid stem fragments and arm plates (arrows) are embedded within a bioclastic wackestone matrix containing fine-grained echinoderm debris. Note the mixture of complete and broken elements. This crinoid limestones originated at the flanks of a mud mound (FZ 5; see Pl. 144). Mississippian (Waulsortian facies): Muleshoe Mound, Sacramento Mountains, New Mexico, U.S.A. 3 SMF 13. Oncoid-ooid rudstone. Large oncoids formed by skeletal and non-skeletal cyanobacteria and encrusting foraminifera (white arrows) in association with smaller bioclasts and radially-structured ooids. The ellipsoidal shape of the oncoids is controlled by the shape of the cores (bivalve shells). Note the irregular knobby surface (black arrows), indicating the absence of reworking of the oncoids. The oncoids shown at the bottom right are compacted and dissolved at their boundary. The oncoids formed in moderately high-energy, shallow waters near a platform edge (FZ 6). The radiallystructured ooids were swept into the environment of oncoid formation from protected parts of the platform. Late Jurassic: Switzerland.
Plate 119: SMF 12 and SMF 13
697
698
Standard Microfacies Types
Please note: SMF 14 has not been forgotten. This type is different from all other SMF Types. It characterizes lag deposits originating during periods of reduced sedimentation or non-deposition. You can find a discussion and a picture of it on Pl. 9/7. Plate 120 Standard Microfacies Types: SMF 15 SMF 15. Oolite (commonly ooid grainstones but also oolitic wackestones) Carbonate sand accumulations with a high percentage of ooids occur on or near the seaward edge of platforms, banks and shelves and they form within the platform interiors and in inner and mid-ramps (Halley et al. 1983). Ooid-rich limestones were originally regarded as one SMF Type characterized by grainstones consisting of abundant, well-sorted and often overpacked ooids, usually occurring in cross-bedded units formed in highenergy environments of oolitic shoals, tidal bars and beaches (FZ 6, winnowed edge sands). This definition did not consider significant differences in the microfabric of the ooid cortices (Fig. 4.20) reflecting differences in environmental conditions and depositional settings (see Sect. 4.2.5; Fig. 4.26 and Fig. 14.22). Concentric ooids with laminae composed of tangentially arranged crystals commonly originate in high-energy marine settings on oolitic shoals, tidal bars and beaches (FZ 6 and FZ 7; mid- and outer ramp settings). Radial ooids characterized by laminae consisting of radially arranged crystals are formed under moderate to low-energy conditions, often in sea-marginal or lacustrine environments with elevated or lowered salinities. Many of these ooids originate in restricted environments (e.g. lagoons, lagoonal ponds; FZ 8). Micritic ooids can be generated by random growth of crystals in marine environments characterized by reduced sedimentation (e.g. outer ramps), growth within biofilms (FZ 8, FZ 9B), or are the result of pervasive micritization. These categories should be distinguished within SMF 15 by using additional letters: SMF 15-C (concentric), SMF 15-R (radial) and SMF 15-M (micritized). More than other carbonate grains, ooids are easily reworked and transported. Oolitic limestones, therefore must be checked to see whether the ooids are autochthonous or parautochthonous, or whether they have been transported by gravity-mass flows and turbidity currents (forming calciturbidite beds in slope and basin sediments; FZ 1 and FZ 4) or whether they were accumulated by wind action (forming eolian dune carbonates; see Fig. 15.1). Additional pictures: Pl. 13/4 (SMF 15-R), Pl. 13/6 (SMF 15-C), Pl. 45/1 (SMF 15-C). 1 SMF 15-C. Ooid grainstone with concentric ooids. The ooids exhibit distinct tangential structures (arrows), are multiplecoated and well-sorted. Some ooids are micritized and a few are dissolved (bottom left). Interparticle pores are filled with drusy cement. Ooid shoals near an outer platform margin (FZ 6). Late Triassic (Upper Rhaetian): Northern Alps, Austria. 2 SMF 15-R. Ooid grainstone with radial ooids. In contrast to SMF 15-C, the laminae of the ooid cortices are radially structured (arrow) or exhibit tangential/radial structures. Other features different from SMF-C are poor sorting, the abundance of non-spherical, highly irregularly shaped grains, abundant reworked and broken ooids, the occurrence of composite grains acting as cores for renewed ooid coatings, and the rareness of fossils associated with the ooids. Limestones with abundant radial ooids occur as ooid wackestones or packstones with ooids within micrite matrix, or ooid grainstones composed of transported, broken and sometimes regenerated radially-structured ooids (this figure). The sample comes from restricted near-coast marginal-marine parts of a carbonate platform (FZ 8). Early Cretaceous: Apennines, Italy. 3 SMF 15-M. Quartz-bearing ooid wackestone with micrite ooids. Characteristic features of these ooids are the wide range in sizes and shapes, poor sorting and ferruginous coatings (black). The ooids are allochthonous grains. They were reworked and redeposited in deep outer ramp environments with discontinuous sedimentation. Sparse fossils are gastropods. Late Jurassic (Oxfordian): Torrecilla-Barranco near Zaragoza, Spain.
Fig. 14.21. Sites of ooid and pisoid formation. Ooids accumulate on land in eolian dunes, and subaqueously in marine shoals, banks and sheetlike bodies on attached and isolated platforms. Allochthonous accumulations occur in turbidites and debris flows deposited on slopes and in basins. On homoclinal ramps ooid accumulation is common near the shoreline in inner ramps and in mid-ramp shoals. Note the different setting of SMF 15 in platforms and ramps! Pisoids and pisolites (SMF 25) occur in terrestrial settings, e.g. caves and caliche, on the shores of some lakes, and in tidal ponds.
Plate 120: SMF 15
699
700
Standard Microfacies Types
Plate 121 Standard Microfacies Types: SMF 16 and SMF 17 SMF 16. Non-laminated peloidal grainstone and packstone and laminated peloidal bindstone Criteria: Accumulations of very small, grain-supported, subrounded or subangular peloids. Two subtypes are common: SMF 16-NON-LAMINATED, characterized by tiny, equal-sized peloids associated with benthic foraminifera, ostracods or calcispheres. SMF 16-LAMINATED is characterized by peloids of different size, forming irregularly distributed fine-grained packstone (pelmicrite), grainstone (pelsparite) and sometimes also packstone fabrics. An other criterion are wavy micritic laminations. The peloids of SMF 16 are mud peloids, fine-grained fecal pellets (definitions in Fig. 4.11), and microbial peloids (SMF 16-LAMINATED). Occurrence: SMF 16-NONLAMINATED is common in shallow platform interiors comprising protected shallow-marine environments with moderate water circulation (FZ 8) and in inner ramp settings, and may also occur in evaporitic arid platform interiors (FZ 9A). SMF 16-LAMINATED contributes to the formation of mud mounds and the stabilization of platform slopes. Additional pictures: Pl. 10/2 (SMF 16-NON-LAMINATED), Pl. 66/1 (SMF 16-NON-LAMINATED). SMF 17. Aggregate-grain (grapestone) grainstone Criteria: Aggregate-grain grainstone. The grainstones and grainstone-rudstones consist predominantly of arenitic and ruditic lumps and grapestones (see Sect. 4.2.7 for definitions), associated with peloids and some coated and micritized skeletal grains. Fossils are usually rare, except for a few foraminifera and algae. Occurrence: Characteristic for platforms, very rare in ramp settings. FZ 7 and FZ 8 (open and restricted shallow platforms). Very abundant in isolated platforms of the Bahama type but also occurring in attached platform settings. Additional pictures: Pl. 15/4, Pl. 15/5, Pl. 15/8, Pl. 43/1. 1 SMF 16-NON-LAMINATED. Fine-grained peloid grainstone. Some of the peloids correspond to tiny rounded micrite clasts. Fossils are very rare except for a few smaller foraminifera. The sample comes from a grainstone bed at the top of a sedimentary cycle. It represents the shallowest subtidal deposition on a relief formed by algal mounds. Pennsylvanian (Holder Formation, Virgilian): Alamogordo, Sacramento Mountains, New Mexico, U.S.A. 2 SMF 16-NON-LAMINATED. Fine-grained peloidal grainstone. Note the similarity with Fig. 1. Arrows point to ‘calcispheres’ (see Sect. 10.2.1.9). The peloids are angular and rounded micrite clasts. The paleoenvironment was a restricted inner platform lagoon (FZ 8) adjacent to an inner ramp. Early Carboniferous: Belgium. 3 SMF 16-LAMINATED. Bindstone consisting of fine-grained laminated peloidal grainstone grading into laminated peloidal packstone. The sample exhibits distinct alternations of thicker pelsparite layers (PS) and thin pelmicrite (PM) layers separated by irregular, probably microbial, micrite laminae. This texture is characteristic for mud mounds, slopes, and protected platform interior environments (FZ 8). Late Jurassic (Oxfordian): High Atlas Mountains, Morocco. 4 SMF 17. Aggregate-grain grainstone. Most grains are micritized aggregate grains (lumps) exhibiting the characteristic lobate outline. Angular lumps indicate some transport and reworking. The composite grains consist predominantly of peloids and some bioclasts bound together by microcrystalline calcite. Most grains have rims of thin early-marine cement. Remaining interparticle pores of the grain-supported fabric were occluded by coarse burial calcite. The sample comes from a rimmed carbonate platform (FZ 6). This grapestone facies originated within a distance of only a few kilometers behind the platform edge. Late Triassic (Rhaetian): Steinplatte, Tyrol, Austria.
A
B
Fig. 14.22. Relationship between grapestone facies (SMF 17) and limestone resources. Aggregate grains originate preferentially in platform settings with moderate but fluctuating water energy, low sedimentation rates and low terrigenous input. The latter condition and the high amount of interparticle calcite cement are the reason for the formation of chemically high-quality limestones. The figure shows an example. A: The Late Triassic Dachstein limestones of the Northern Calcareous Alps were formed on an attached platform (FZ 9 and FZ 8) rimmed by marginal reef complexes (FZ 5, FZ 4, and parts of FZ 7). Grapestones originated in the back of the reef zone. B: A survey of the Triassic Dachstein limestones revealed that the reef facies and the grapestone facies have the highest calcium carbonate content, are relatively low in MgCO3 and contain very little insoluble residue. These criteria characterize both facies types, particularly the grapestones facies, as targets for the successful exploration of industrially used pure limestone resources. A and B: Modified from Flügel (1977 and 1981).
Plate 121: SMF 16 and SMF 17
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Standard Microfacies Types
Plate 122 Standard Microfacies Type: SMF 18 SMF 18. Bioclastic grainstones and packstones with abundant benthic foraminifera or calcareous green algae Criteria: The main feature of this SMF Type is the high abundance of benthic foraminifera or calcareous green algae. Other grains are peloids, cortoids and composite grains. Common textures are grainstones and packstones. Species diversity is low to moderate. The dominant biota should be indicated by additional letters (SMF 18-FOR, foraminifera; SMF 18-DASY, dasyclad algae; SMF 18-UDO, udoteacean green algae; SMF 18GYMNO, gymnocodiacean algae). Occurrence: Sediments of SMF 18 occur as bars and channels, and in sand shoals heaped up by tidal currents in shallow lagoons and bays (restricted platform, FZ 8) and in shelf lagoons with open circulation (FZ 7). Foraminiferal sands form also in backreef areas and on reef flats (Fig. 14.23). SMF 18 sediments are common in inner ramp settings. Additional pictures: Pl. 57/7 (SMF 18-GYMNO), Pl. 57/7 (SMF 18-GYMNO ), Pl. 61/1 (SMF 18-DASY), Pl. 75/6 (SMF 18-FOR ), Pl. 105/2 (SMF 18-DASY), Pl. 145/4 (SMF 18-DASY). 1 SMF 18-FOR. Foraminiferal grainstone with abundant miliolids. Miliolid foraminifera are very common in lagoonal environments (sometimes with elevated salinity) of Mesozoic and Cenozoic restricted inner platforms and inner ramps. Most associated particles are aggregate grains (AG). The sample comes from a restricted platform lagoon (FZ 8). Late Cretaceous (Albian to Cenomanian): Trnovo, Dinaric Platform, Slovenia. 2 SMF 18-DASY. Dasyclad grainstone characterized by concentrations of Spinaporella andalusica Flügel and FlügelKahler. Spinaporella is characterized by long spines. Most of the smaller particles (e.g. bottom center) are broken algal elements. Note the geopetal infilling of the algal thalli (black arrows). The sample comes from a backreef lagoon behind a reef formed predominantly of serpulids and synsedimentary carbonate cement. Open-marine platform environment (FZ 7). Late Triassic (Norian): Sierra Nevada near Granada, southern Spain. 3 SMF 18-GYMNO. Gymnocodiacean algal packstone characterized by the concentration of debris of gymnocodiacean algae (see Sect. 10.2.1.5). The micritic matrix was formed by disintegration of the algae, a mode that corresponds to the ‘Halimeda model’ (see Sect. 4.1.2; Fig. 4.3). The limestone originated in a low-energy, lagoonal inner ramp environment. Permian (Yeso Formation): Cloudcroft east of Alamogordo, New Mexico, U.S.A.
Fig. 14.23. SMF 18-FOR. Foraminiferal grainstone. The massive calcarenite consists of densely packed rotaliid and lepidocyclinid foraminifera with Nephrolepidina praemarginata. The skeletal sand is part of a reef flat situated between the back reef area and the front of a fringing reef. The strongly varying orientation of the tests, poor sorting and the admixture of some corallinacean material (grains here white, because negative print!) indicate high-energy deposition by storms. The sample would be attributed to Facies Zone FZ 7 according to the abundance of benthic foraminifera, but to Facies Zone FZ 5 (‘reef’) if one would only consider the place from where this sample was taken. Negative print. Tertiary (Castro Limestone, Late Oligocene): Salento Peninsula, southern Italy. After Bosellini and Russo (1992).
Plate 122: SMF 18
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Standard Microfacies Types
Plate 123 Standard Microfacies Types: SMF 19, SMF 20 and SMF 21 SMF 19. Densely laminated bindstone Criteria: Bindstone with a laminated fabric consisting of millimeter thick laminae that differ in internal composition. Finer laminae correspond to micrite layers. Somewhat coarser laminae consist of densely spaced, very small peloids or sparry calcite. The biofabric is characterized by couplets with irregular, sometimes wavy boundaries. Fossils are rare except ostracods and some foraminifera, gastropods and a few algae. Occurrence: Near-coast platform interior (FZ 8 and FZ 9). Tidal flats of attached and isolated platforms and inner ramps. Additional picture: Pl. 132/1. SMF 20. Laminated stromatolitic bindstones/mudstones Criteria: Planar or variously dome-shaped laminated bindstones composed of fine- or coarse-grained laminae sometimes exhibiting microbial or algal structures (see Sect. 9.1.4; Pl. 50). Occurrence: Very common in the intertidal zone, but also in supratidal and shallow subtidal environments. Open platforms (subtidal, FZ 7), tidal zones of restricted lagoons (FZ 8) and arid coasts (FZ 9A, supratidal). Additional picture: Pl. 132/5. SMF 21-FEN. Fenestral packstones and bindstones SMF Type 21 was originally defined by a ‘spongiostrome mudstone fabric’ with algal tufts. Relicts of porostromate algae or calcimicrobes may be preserved (SMF 21-PORO; see Pl. 124/1) or not. More useful criteria in distinguishing this microfacies are fenestral fabrics (SMF 21-FEN) typically developed in FZ 8 and FZ 9A. Pl. 123/3 and Pl. 124/1 display two common variants of SMF 21 occurring in association or separated. Criteria: Bindstone characterized by variously sized fenestral cavities within a framework formed by sedimentary and biogenic grains. Smaller fenestrae correspond to birdseyes, larger structures may have stromatactoid shapes (compare Sect. 5.1.5 and Pl. 20). Occurrence: Supratidal and intertidal environments (restricted lagoons, FZ 8) and evaporitic lagoons (FZ 9A). Additional pictures: Pl. 20/2, Pl. 41/2, Pl. 46/2, Pl. 50/5, Pl. 132/6. SMF 19, SMF 20 and SMF 21 are common and in facies analyses often used as indicators for intertidal and supratidal environments. The main criteria are lamination textures caused by microbial/algal mats and fenestral fabrics. Differences in the morphology of the mats (–> 1 and 3) in combination with other sedimentary, organic and mineralogical features are used in estimating the ‘exposure index’ within the peritidal zone. This index reflects the percentage of time during which the area where the laminated sediment was formed was exposed (Ginsburg et al. 1977). This approach also can be applied to ancient peritidal carbonates and allows microfacies samples to be allocated to predominantly intertidal, mostly supratidal but with some intertidal, or only supratidal subenvironments (Smosna and Warshauer 1981). 1 SMF 19. Laminated bindstone consisting of alternating couplets of micritic and peloidal layers (FZ 8). This microfacies may grade into bindstones with fenestral fabrics. Late Permian (Main Dolomite, Zechstein): Subsurface, Miloszewo well, Northern Poland. 2 SMF 20. Laminated stromatolite bindstone. The sample shows finely wrinkled laminae arranged in dense and closelyspaced growth patterns exhibiting vertically stacked protuberances (arrows). This fabric meets the definition of stromatolites, favored by Riding (1999): ‘A stromatolite is a laminated benthic deposit’. The irregular extended structure marked by SP probably represents soft non-rigid sponges. The irregularly-shaped white calcite-filled structures indicate burrowing and disintegration of the microbial mats by grazing organisms. The sample comes from cyclic bedded limestones (member B of the ‘Lofer cycle’). Intertidal environment (FZ 8). Late Triassic (Dachstein limestone, Norian): Zelenica, Begunjscica Mountains, northern Slovenia. 3 SMF 21-FEN. Fenestral packstone characterized by small spar-filled cavities (birdseyes, white arrows) and larger irregular voids some of which are filled with fine-grained sediment at the bottom (stromatactoid fabric, black arrow). The sedimentary framework formed by peloids and cyanobacteria tufts (CB) encloses transported dasyclads (D). Porostromate cyanobacteria tufts are characteristic constituents of micritic, non-laminated intertidal carbonates. The sample comes from the intertidal member of a Lofer cycle (see Sect. 16.1.2). Early Jurassic: Korfu Island, Greece. –> 1: Peryt 1986
Plate 123: SMF 19, SMF 20 and SMF 21
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Plate 124 Standard Microfacies Types: SMF 21 and SMF 22 SMF 21-PORO. Fenestral bindstones, mudstones and packstones with porostromate microstructures Criteria: Like the SMF 21-FEN, the fabric exhibits small fenestrae, but unlike this type the sedimentary framework consists of densely growing tufts of porostromate cyanobacteria and algae (see Sect. 10.2.1.1). Ostracods and a few foraminifera may be present. Occurrence: Platform interior (FZ 8). Common in intertidal pools and very shallow subtidal environments. SMF 22 Oncoid floatstones or wackestones Criteria: Millimeter- to centimeter-sized agglutinated oncoids, consisting of sedimentary grains (e.g. tiny skeletal grains, terrigenous quartz) trapped and bound together by non-skeletal microbes and algae (compare Pl. 11/1, 2). Occurrence: Low-energy environments of shallow lagoons and tidal zones (FZ 8). Typically found on the edges of ponds with brackish water. Additional picture: Pl. 12/7. 1 SMF 21-PORO. Fenestral bindstone with porostromate microfossils. A framework consisting predominantly of cyanobacteria (arrows) is well-preserved. Some of the large voids (V) may represent desiccation structures. The sample comes from a very shallow backreef lagoon behind platform-edge reefs. Late Triassic (Dachstein limestone, Norian): Gosaukamm, Austria. 2 SMF 22. Oncoid floatstone. The oncoids were formed by trapping and binding of sedimentary grains (quartz silt, tiny bioclasts). Note the differences in organic matter (black) and in the frequency of grains within the irregularly developed laminae. The cores of the oncoids are gastropods. The oncoids correspond to the ‘spongiostromate’ oncoid type. Compare the distinctly different structure of oncoids formed in subtidal normal-marine environments (SMF 13; Pl. 119/3). The sample comes from transitional-marine brackwater deposits. Cretaceous: Switzerland. –> 1: Wurm 1982 Fig. 14.24. Many fenestrae originate from the decomposition of algae and cyanobacteria. A: Layers with carbonate grains and layers with filaments of Scytonema are common features of tidal algal mats in the Bahamas. B: Oxidation of the mat destroys the non-calcified cyanobacteria, leaving the fenestrae (black) within the grain-supported framework. Depending on algal growth patterns, ‘stromatactoid’ voids with irregular roofs (this figure) or isolated spar-filled voids (Fig. 14.25A) will develop. Modified from Monty (1976).
Fig. 14.25. SMF 21-FEN. A: Fenestral packstone characterized by birdseyes and a pelmicritic framework. Note the thin cement rim within the birdseye structures. B: Red fenestral limestone. The red rock color is caused by the geopetal infilling of many birdeyes with red vadose silt indicating intermittent subaerial exposure of tidal platform interior carbonates. Protected part of FZ 7. Early Jurassic: Djbel Bou Dahar, High Atlas, Morocco. After Scheibner and Reijmer (1999).
Plate 124: SMF 21 and SMF 22
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Standard Microfacies Types
Plate 125 Standard Microfacies Types: SMF 23, SMF 24 and SMF 25 The plate displays two SMF Types distinguished by Wilson (1975) and the new SMF Type 25 characterizing carbonate-evaporite transitional environments. SMF 23. Non-laminated homogeneous micrite and microsparite without fossils Criteria: Unfossiliferous lime mudstone or fine-grained dolomicrite, sometimes with authigenic evaporite minerals. Occurrence: Tidal flats (FZ 8) and arid evaporitic coasts (FZ 9A). Deposited in saline or evaporative environments, e.g. in tidal ponds. Remarks: Note that micrites without fossils or other grains may also occur in anoxic deep basins (FZ 1B), and that the ‘lack’ of fossils can only be a sampling effect. SMF 24. Lithoclastic floatstones, rudstones and breccias Criteria: Coarse, up to several centimeter-sized angular carbonate lithoclasts. Elongated micrite clasts are abundant, often edgewise arranged, and sometimes cross-bedded, forming an intraformational breccia. Occurrence: Platform interior (FZ 8). Lag deposits in tidal channels and on tidal flats. SMF 25. Laminated evaporite-carbonate mudstone Criteria: Alternation of fine-crystalline carbonate (limestone, dolomite) which might be microbially induced, and diagenetically deformed layers with evaporite crystals (gypsum). Occurrence: Upper intertidal to supratidal sabkha facies (FZ 9A) in arid and semiarid coastal plains and evaporitic lacustrine basins. 1 SMF 23. Non-laminated unfossiliferous mudstone. Irregularly distributed voids result from the replacement of evaporite minerals by calcite. FZ 9A. Middle Triassic: Northern Calcareous Alps, Austria. 2 SMF 23. Non-laminated unfossiliferous dolomite mudstone. Note the calcitized anhydrite rosettes within the fine-grained idiotopic dolomite matrix. The micrite matrix corresponds to type 1 distinguished in Fig. 4.1, and is probably of physicochemical origin, as indicated by the very fine crystal size, the evaporite minerals, and the complete lack of fossils. Supratidal environment. FZ 9A. Late Tertiary (Nullipore Formation, Miocene): Subsurface, Gulf of Suez, Egypt. 3 SMF 24. Coarse lithoclastic rudstone-floatstone. Imbricated flat pebble conglomerate consisting of partially dolomitized, laminated ‘algal’ clasts (AC; algal mat chips) and exhibiting desiccation cracks (DC). Note the stylolitic joints of the pebbles, indicating compaction and consolidation of the sediment prior to erosion and redeposition. The matrix is formed by skeletal grains and neomorphic sparite. The conglomerate has been interpreted as storm sediment deposited within tidal flat areas (FZ 8). Late Triassic (Calcare di Zu, Rhaetian): Denti della Vecchia, Lombardian Basin, northern Italy. 4 SMF 25. Laminated evaporitic lime mudstone composed of alternating layers of organic-rich calcimudstone (M), and evaporitic layers (E) with tiny gypsum crystals. Some layers show early diagenetic entherolitic folds characterized by antiform buckles. Note the deformation of folded and non-folded layers by the growth of large gypsum crystals (arrows). The sample comes from the deep evaporitic part of a lacustrine basin. Early Tertiary (Eocene): Mormoiron Basin, France. –> 3: Lakew 1990
Fig. 14.26. SMF 25 is a characteristic of sabkha facies originally described from the Abu Dhabi coast on the Arabian Gulf. Sabkha is an Arabian word for a salt flat found near sand dunes. The sabkha cycles start with subtidal shelly lime mud and muddy skeletal sands commonly deposited in a lagoonal environment, overlain by intertidal laminated sediments characterized by algal/microbial mats containing gypsum. The supratidal part of the cycle consists of anhydrite layers whose nodules display a ‘chickenwire’ texture as the result of displacive growth. The ancient counterpart of this cycle is represented by subtidal fine-grained dolomites or lime mudstones with rare fossils (mudstones, sparse wackestones), followed by intertidal layered dolomicrites or micritic limestones exhibiting parallel and/or irregular, wavy laminations, and sometimes vanished evaporites. The supratidal sabkha facies can be recorded by anhydrite beds. See Warren and Kendall(1985) for a review of sabkha criteria. Modified from Achauer (1982).
Plate 125: SMF 23, SMF 24, SMF 25
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710
Plate 126 Standard Microfacies Types: SMF 26 Figs. 1 and 2 of this plate show a SMF Type introduced here. Main criteria are abundant calcareous pisoids/ vadoids (see Sect. 4.2.6), formed by early diagenetic meteoric carbonate precipitation, predominantly but not always, during phases of subaerial exposure (–> 2). This microfacies is abundant in humid and tropical settings, e.g. paleokarst limestones, caliche carbonates, and travertines. It is associated with SMF 9A in evaporitic settings (FZ 9B). Pictures 3 and 4 on this plate show non-carbonate pisolitic structures. Additional pictures: Pl. 14/ 1, Pl. 14/3, Pl. 34/6, Pl. 128/6. SMF 26. Pisoid cementstones, rudstones and packstones Criteria: Accumulations of pisoids both autochthonous or allochthonous characterize this important faciesdiagnostic SMF Type. Diagnostic criteria are variously shaped, millimeter- to centimeter-sized, densely packed pisoids, commonly cemented by meteoric cements. The cores of the pisoids are usually broken pisoids or cement crusts. Reworking and redeposition is common. Many pisoids are inversely graded, exhibit polygonal fitting, and are associated with tepee structures. Indigenous biota are usually lacking. Occurrence: Terrestrial subaerial, or terrestrial and transitional terrestrial-marine subaquatic settings, formed under meteoric-vadose and marinevadose conditions (FZ 10). 1 SMF 26. Pisoid rudstone. These pisoids are the grains described in the classical paper by Dunham (1969; see Sect. 4.2.6 and Pl. 14/1) and interpreted as caliche soil pisoids formed by subaerial vadose diagenesis during episodic lowstands of sea level. At present, the range of depositional interpretations of these pisoids includes caliche formation in continental or coastal-spray-zone supratidal settings, vadose-marine inorganic precipitation in inter- and subtidal environments or formation in marine seepage or groundwater springs (Handford et al. 1984). The pisolitic cementstone consists of large pisoids/vadoids linked by outer laminae and meniscus cement. Most pisoids are well-laminated, their cortices consist of accretionary, essentially isopachous cement laminae. The non-skeletal nuclei are preexisting pisoids or peloids. Arrows point to vadose cements. Successive laminae differ in thinner darker micritic and clearer microspar lamellae. Most black nuclei are broken fragments of preexisting pisoids. White nuclei correspond to reworked pieces of laminated interpisoid cement. Late Permian (Carlsbad Group, Guadalupian): Entrance to the Carlsbad Cavern, New Mexico, U.S.A. 2 SMF 26. Vadose pisolitic crusts indicating cyclic exposure of lagoonal carbonates. The thin section displays inversely graded pisoids (also called vadoids because of their formation within vadose diagenetic environments) made of pisoid fragments, which in turn are cores of younger vadoids (V). Some vadoids exhibit radial cracks (RC), indicating sporadic desiccation. The pisoids originated during subaerial exposure phases of parts of intertidal and shallow subtidal lagoonal sediments in the back of platform-margin sponge-coral reefs. The pisolitic crusts may form beds reaching a thickness of tens of centimeters to a few meters. Middle Triassic (Wetterstein limestone, Ladinian): Scharnkopf near Reichenhall, Northern Calcareous Alps, Germany. Non-carbonate pisoids 3 Silicified pisolite breccia. Cretaceous: Subsurface. Ras Al Khaimah, United Arabian Emirates. 4 Pedogenic pisoids. Pisoids occur both in carbonate and non-carbonate rocks. The sample is a lateritic pisolite representing a pedogenic sediment. The pisoids have a diameter up to 5 mm and consist of a kaolinitic or pisolitic core and a cortex composed of iron oxide. Note the shrinkage cracks within the cores (arrows). The matrix consists of hematite (black) and kaolinite (white) representing reworked material of a weathered Precambrian basement (crystalline rocks). Sedimentary structures indicate mass-wasting and fluviatile transport mechanisms. Late Cretaceous (Nubian Group): Wadi Kalabsha southwest of Assuan, Egypt. –> 2: Henrich 1984; 4: Fischer and Germann 1987
A
B
Fig. 14.27. Genetically different pisoids differ in their morphology. A: Phreatic-vadose pisoids often exhibit irregular shapes, polygonal fitting, abundant laminae and downward bulging. B: Caliche pisoids are characterized by centrifugal and accretional growth, and a few discontiuous laminae. A: Modified from Esteban and Pray 1983. B: After Calvet and Julia 1983.
Plate 126: SMF 26
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14.3.2 Discussion of Standard Microfacies Types How many samples fit into the SMF Type categories? A common experience is that not all thin sections can be precisely assigned to a specific SMF Type. A review of fifty papers describing and illustrating the microfacies of Paleozoic and Mesozoic carbonate rocks shows that about 70% of the local microfacies distinguished by the authors correspond to established SMF Types. Difficulties occur with local microfacies types represented by wackestones and grainstones containing various types of grains and of fossils. Many of these samples are sediments of the platform interior (FZ 6– FZ 8) and of the deep shelf (FZ 2). In these environments mixing of grains by burrowing has a strong impact on the composition of the sediment. In addition, storms frequently contribute to sediment mixing in interior platform environments. ‘Mixed grain compositions’ are also common in ramp carbonates (Sect. 14.3.4.2). These samples should be studied using the ‘compositional maturity’ concept proposed by Smosna (1987), see Sect. 6.2.1.4. SMF Types assist in facies interpretations but must be refined More than two third of the SMF Types distinguished by Wilson (1975) describe sediments deposited above the fair-weather wave base and between the wave base and the storm wave base. Microfacies types of deepwater sediments are clearly underrepresented. SMF types comprise autochthonous and parautochthonous sediments deposited more or less within the place of their origin, and allochthonous sediments, transported from shallow to deeper depositional environments. The autochthonous background sedimentation in basinal and deep shelf settings is represented by SMF 1 and SMF 3. SMF 2 indicates mixing of autochthonous sediment and fine-grained eroded and redeposited material. The muddy and often burrowed sediments of SMF 1 and SMF 3 may present indications for bottom currents. The microfacies of these bioclastic wackestones is common in limestones characterized by nodular fabrics (see Sect. 5.1.6). SMF 4, SMF 5 and SMF 6 are allochthonous sediments, whereby the composition of SMF 5 and SMF 6 provide some hints to the existence of reefs. These reefs might have developed at platform margins, in an upper slope position, or on non-rimmed platforms. Microfacies and fossils in clasts of breccias or foreslope rudstones allow source areas to be recognized (see Sect. 16.2.5). Note that reef material as well as lithified sand
Standard Microfacies Types: Discussion
shoal material may be transported by storms onshore as well. The interpretation of SMF 4 sediments requires information on the geometry and thickness of the allochthonous limestone beds and their position within the sequence. One of the difficulties of the Standard Microfacies concept is the highly generalized definition of SMF Types of fossiliferous limestones. These limestones, represented by SMF 8, SMF 9, SMF 10 and SMF 11, form the bulk of bedded shelf carbonates. SMF 8 to SMF 11 are common both in the Facies Zone FZ 7 of the open shelf lagoon in euphotic depths, and in the Facies Zone FZ 2 of the deep shelf within depths of tens to hundreds of meters. Environmental interpretations derived from these SMF Types must be checked by palecological and biofacies criteria (see Chap. 15 for examples). The assignment of samples to SMF 8 and SMF 9 can be difficult, because both microfacies types may show strong variations in the frequency of skeletal grains and other grain types. SMF 12 summarizes accumulations of various shells formed in different ways. A sound interpretation of samples attributed to this SMF depends on information on the stratigraphic and geometric characteristics of the shell bed, thickness and lateral extension, position within the sedimentary sequence, and the kind and taphonomy of fossils (Kidwell et al. 1991; Fürsich and Oschmann 1993; Fürsich 1995). SMF 13 and SMF 22 are characterized by oncoids. These types are well-defined. Most of the environmental information is included in the composition of the oncoid cortices. Take a look at the organisms contributing to the growth of the oncoids, the type of the lamination, and consider breaks in growth indicated by encrusting and boring organisms. SMF 14: A very specific SMF Type is SMF 14 (Pl. 9/7) introduced by Wilson for lag deposits. Lag deposits are residual accumulations remaining on a surface after the finer material has been removed. Lags represent slow accumulations of coarse material in the winnowing zone. The particles are chemically and physically resistent through long periods of non-deposition and persistent winnowing of sediment on the sea floor. The criteria of carbonate lag deposits are coated, rounded and worn particles, in places mixed with peloids or ooids that are blackened and iron-stained. Phosphate grains e.g. fish bones and teeth and phosphatized organic remains are common (Pl. 110/2, 3, 5) as well as glauconite and quartz grains. Coarse lithoclasts may be present, sometimes forming lag conglomerates. Fossil shells are often leached; steinkern preservation is common. Lag deposits are characteristically thin beds.
SMF Types: Discussion
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Fig. 14.28. SMF Types assist in recognizing facies belts independently of the age of the rocks. Note the amazing similarity in grain composition and depositional texture of the two samples from Jurassic and Cretaceous limestones. Both samples represent SMF 17, characteristic for sediments formed in platform interior environments when water depth has remained relatively constant (late highstand phase). The similarity includes not only the grainstone texture and the dominance of aggregate grains but also the occurrence of benthic foraminifera and algae, indicating very shallow warm-water sedimentation. A: Thick-bedded limestone within a subtidal shallowing-upward sequence. Foraminifera are textulariids and Trocholina (T). Green algae are represented by fragments of Clypeina (arrow). Late Jurassic: Sulzfluh, Graubünden, Switzerland. B: Thin-bedded lagoonal limestone at the base of a mixed siliciclastic-carbonate sequence. Miliolid foraminifera are Cyclogyra (right) and Quinqueloculina (center top). Late Cretaceous (Cenomanian): Monastery St.Anthony, Eastern Desert, Egypt. Scale valid for both pictures. After Kuss (1986).
Occurrence: Removal of fine-grained sediment by strong winnowing occurs in high-energy near-coast environments (FZ 7), but also through bottom currents, e.g. in deep shelf and basinal environments (FZ 2 and FZ 1). SMF 15 characterized by ooids, SMF 16 defined by accumulations of peloids, and SMF 17 typifying carbonates rich in aggregate grains are frequent constituents of shallow platform carbonates. SMF 15 and SMF 16 are also common in inner- and mid-ramps in contrast to SMF 17 that appears to be very rare in ramp carbonates, but is an excellent indicator for platform interior environments (Fig. 14.28). These three SMF Types are easy to recognize if the diagnostic grains dominate quantitatively. Many samples, however, contain mixtures of aggregate grains and peloids, or aggregate grains, peloids and skeletal grains. The assignment of these samples to one of the SMF Types may be somewhat tricky. SMF 18 characterizing accumulations of foraminiferal shells or calcareous green algae is well-defined but unites accumulations having different origins. The miliolid foraminiferal sand of Pl. 122/1 is an autoch-
thonous sediment as indicated by the good preservation of the foraminiferal tests and their association with aggregate grains. Pl. 122/2 demonstrates the formation of a coarse algal sand through the disintegration of dasyclad thalli. Disintegration was stronger for the gymnocodiacean algae (Pl. 122/3) that fell apart into smaller particles than the dasyclad algae. Both samples show how algae contribute to in-place sediment production, but the dasyclad living in high-energy areas, where most of the mud has been winnowed, produced skeletal sand recorded by a grainstone fabric, whereas the gymnocodiacean algae living in a low-energy environment produced lime mud. Contrasting these examples, the Cretaceous foraminiferal grainstone depicted in Fig. 14.23 was formed by a short storm event that may have occurred in different parts of the platforms and ramps. A closer look at the taxonomic composition of the foraminiferal fauna may tell you in which ecological zones the tests have been reworked (see Fig. 14.7). The joint characteristic of SMF 19, SMF 20 and sometimes also of SMF 21 is a fine biogenic lamination whose origin is related to trapping and binding activities of microbes and algae, and to the precipitation
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Problems with SMF Types
Plate 127 Problems with Standard Microfacies Types: Textural Changes in Bedded Limestones Fluctuations in water energy and sediment transport can result in the formation of sedimentary layers differing in depositional texture, constituent composition as well as the size and packing of grains. These differences might give rise to SMF Type assignments that do not necessarily reflect the factual sedimentary environments. To recognize these environments you need a combination of field data, paleontological criteria, and microfacies. Small-scaled, short-time changes in depositional texture and the association of grain types are common in platform interior areas and in inner and mid-ramp environments influenced by storm-induced currents. These changes in texture and fabric differ from changes caused by turbidity currents and documented by distinct vertical succession patterns and wide areal extensions of sedimentary structures. 1 Changes in depositional energy. The conspicuous characteristic of the sample is the alternation of cortoid grainstone (G) and peloid packstone (P) fabrics. The lower grainstone layer (G1) and upper layer top (G3) consist of skeletal grains (mostly corals - C, some dasyclad algae, DA, and a few mollusk shells, MS). Larger black grains are porostromate oncoids (O). All skeletal grains are coated by micrite envelopes. Some grains are encrusted by foraminifera. Intergranular pores are occluded by thin rims of early marine cements surrounding the grains, and late diagenetic cements (white) occluding the remaining pore space. The intercalated packstone layer (2) consists of small peloids, most of which are reworked and transported micrite clasts. The coated grainstone shows criteria of SMF 11, the packstone of SMF 16-NON-LAMINATED. The assignment to SMF 11 is supported by field data and paleontological criteria. The sample comes from bedded carbonates corresponding to calcareous sands formed at the margin of a platform in a backreef position. An assignment of the packstone, however, to SMF 16 would suggest tidal flat sedimentation and not reflect the genuine depositional environment. The packstone layer originated from a very short high-energy event causing a redeposition of peloids from nearby protected areas. Note the transitional character of the lower boundary of the packstone layer and the infilling of peloids in the underlying grainstone. White arrows point to geopetals (deposition of grains on top of shells). Strong shearing of the limestones is indicated by different orientation and filling of microfractures. Late Triassic (Dachstein limestone, Norian): Loser, Styria, Austria. 2 Storm sedimentation. Transition of ooid grainstone (OG, 1) to bivalve shell- ooid rudstone (BOR, 2) at the top. Ooids are conspicuously small and more or less the same size, indicating a pre-depositional sorting of the grains. The microstructure of the shells shows reworking of ostreid bivalves. The blurred boundary between the two textures suggests that deposition of both layers occurred during a single event. Ooids have settled down, coming to rest in the cup-shaped valves (note the geopetals indicated by arrows), and filling shelter pores beneath the shells. The pores were occluded by calcite cement. The upper part of the sample shows similarities with SMF 12-BS (bivalve shell hash). The dominance of ooids would suggest an assignment to SMF 15 pointing to deposition in ooid shoals. However, the intimate relations between ooids and shells as well as the chaotic orientation of the shell fragments characterize the sample as a storm deposit which can not be categorized by a specific SMF Type. The sample comes from a storm-dominated inner ramp. Late Jurassic (Oxfordian): Osterwald, Lower Saxony, Germany. 3 Changes in water energy causing differences in grain packing. The succession exhibits a lower packstone/bindstone fabric (P/B, 1), separated by a distinct boundary (arrows) from an upper grainstone fabric (G, 2), which again grades into a packstone/bindstone fabric (P/B, 3). The densely packed peloids, coated grains, oncoids and algae of the lower part are bound together by microbial crusts. The grainstone of the upper part contains reworked fragments of porostromate algae (PA) and gastropod shells (GA). Note oblique microfractures (MF) occluded with spar and gray silt. The sharp boundary between the two textures suggests a more rapid stabilization and cementation of the bindstone than the grainstone. The textural changes reflect short-term changes in water energy conditions. The packstone/bindstone texture corresponds approximately to SMF 16, indicating deposition within a shelf lagoon (Facies Zone FZ 7). Paleontological data support this interpretation. The sample comes from an isolated Bahamian-type platform (see Pl. 134). Late Jurassic (Tithonian): Sulzfluh, Graubünden, Switzerland. 4 Break in shallow-marine sedimentation. The lower part of the succession (1) consists of variously sized, angular and rounded lithoclasts embedded within a fine organic network formed by microbes and synsedimentary cements. The upper part (3) exhibits a similar fabric but the lithoclasts are larger and of different composition. Some clasts are reworked aggregate grains, others are extraclasts. The gray central part (2) consists of micritic cement, and differs from the lower and upper layers by the absence of lithoclasts. The erosion at the top of the cement layer (arrows) and micro-encrusters (ME) characterize a hardground evidencing sporadic breaks in deposition. Attribution of the sample to SMF Types might lead to SMF 4 (microbreccia deposited on slopes) or SMF 5 (forereef slope). Neither designation fits the real paleoenvironment. As shown by field data, the sediment originated in a near-coast environment where semi-indurated muds with algal and aggregate grains were reworked and stabilized by microbial mats. Lithoclasts are storm-transported chips. Late Cretaceous: Central Apennines, Italy.
Plate 127: Problems with Standard Microfacies Types
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of carbonate induced by the life cycle and the decay of these organisms. The environmental interpretation of these SMF Types should consider the classification of benthic microbial carbonates described in Sect. 9.1.5 and summarized in Fig. 9.1. SMF 23 is a poorly constrained type. The lack of fossils in micritic limestones or dolomites may have many reasons, one of which could be an elevated salinity. SMF 23 should only be considered if pseudomorphs after gypsum, anhydrite or halite crystals, or the stratigraphic position of the micritic beds support the interpretation. SMF 24 represents a common characteristic of tidal deposition but can also originate from storms in shoreface settings. Coastal and inland sabkhas are characterized by SMF 25, subaerially exposed areas in various climate zones by SMF 26. Both types reflect non-marine conditions. Depositional variability and Standard Microfacies The depositional texture and composition of bedded limestones can change within a scale of some millimeters and a few centimeters. These changes are reflected by different microfacies that are of local importance only. However, some of these small-scale microfacies types indicate substantial changes that are reflected by Standard Microfacies Types. Examples are shown in Pl. 127. SMF Types do not cover the whole spectrum of microfacies Some samples characterized by conspicuous microfacies criteria can not be attributed to an SMF Type of the rimmed platform model, owing to the lack of appropriate SMF Types. Examples are burrowed lime mudstones (e.g. Pl. 25/4) common in deep shelf and deep-shelf margin environments, and intraclast packstones and wackestones occurring in restricted platform interior, in toe-of-slope settings and common in distally steepened ramps. Nodular limestones (Pl. 21 and Pl. 139) represent another example. Some of these microfacies types are included in ‘Stratigraphic Microfacies Types’. 14.3.3 Stratigraphic Microfacies Types Geologists working in different parts of the world know the amazing similarities of the facies of carbonate rocks of similar age. Limestones of specific long-range time intervals are characterized by recurrent microfacies types exhibiting a wide geographic distribution and a limitation to particular major facies belts. The significant global coincidences of some of these ‘stratigraphic
Microfacies Types of Carbonate Ramps
microfacies types’ are sometimes caused by climatic controls determining biogeographical distribution patterns of shelf organisms. One example is the congruent microfacies types of Late Triassic platform and reef carbonates in different parts of the Tethys. Other reasons for time-dependent microfacies types are the combined effects of specific oceanographic conditions and evolutionary stages, exemplified by the worldwide distribution of nodular deep-shelf carbonates in the Jurassic, and the Late Cretaceous pelagic chalk facies triggered by evolutionary bursts of the calcareous plankton. Lists of the common stratigraphic microfacies types for the Devonian, Late Paleozoic, Late Triassic, Jurassic and Cretaceous can be found in Wilson (1975). 14.3.4 Common Microfacies Types of Carbonate Ramps The SMF Types, describing sediments of tropical rimmed shelves, can only be sometimes found in carbonate shelves formed in differing latitudinal and climatic settings and developing different geometries. The following sections compare the common microfacies criteria of carbonate ramps with those of rimmed platforms and summarize the common and recurrent microfacies types of ramp carbonates. Many microfacies criteria of platform carbonates can also be seen in thin sections of limestones formed in ramp settings. However, there exist differences concerning the distribution and type of the prevailing matrix, grains and fabrics. The occurrence of these criteria on inner, mid- and outer ramps is controlled by offshore and onshore transport by storms, the growth of mud mounds and reefs in different parts of the ramps, and the bathymetric dependence of benthic organisms that contribute to carbonate production. 14.3.4.1 Microfacies Criteria of Carbonate Ramps The fine-grained matrix is micrite and calcisiltite. Micrite is common in deeper outer ramps and in protected areas of inner ramps. Both allomicrites and automicrites contribute to carbonate production. Allomicrites are characterized by poorly sorted matrix exhibiting very fine peloids within a matrix, with small bioclastic debris probably caused by bioerosion, variations in packing and mixed microfossil faunas. Automicrite matrix is a common constituent of mud mounds. Clues to inplace carbonate production in deeper ramps (Jeffery
Microfacies of Carbonate Ramps
717
and Stanton 1996) are down-ramp sequential distribution of depth-sensitive organisms (see Fig. 14.12), well preserved and unabraded fossils, beds and bed sets that are traceable along ramp slopes for tens of kilometers, and that are also traceable through mounds growing on the ramps, geopetals in original position, and lack of sedimentary structures indicative of transport. Calcisiltite is abundant in outer ramps, characterizing a particular microfacies type (RMF 1). The origin of these calcisiltites is enigmatic. Abrasion in storm-influenced ramp parts and transport to deeper ramp parts may be one of the sources in Paleozoic ramps. Pelagic fallout contributed to the formation of calcisilite and allomicrite of Mesozoic and Cenozoic ramps. Skeletal mineralogy might have had some influence on the development of ramps that are more abundant in calcitedominated times with maxima in the Cambrian–Ordovician, Mississippian and Jurassic than in aragonite-dominated times.
isms by bottom currents is indicated by fossils that are uncommon in the depositional environments where they are found (e.g. larger foraminifera in outer ramp settings).
The dominant grain types of carbonate ramps are skeletal grains and ooids, followed by peloids and intraclasts. Other grain types are quantitatively of minor importance (cortoids, oncoids), or are very rare or totally absent (pisoids, aggregate grains). Abundant aggregate grains are characteristic of carbonate platforms but missing in ramp settings.
Peloids occur as fecal pellets in inner ramps and as mud peloids in outer ramps. Cortoids are present in current-washed sand shoals of inner ramps, but appear to have not the same importance as in rimmed platform carbonates. Oncoids occur in inner and outer ramps. Oncoids from shallow and deep environments differ in nuclei, organisms contributing to the cortices, lamination type, and the composition of the encruster association. Oncoids built by cyanobacteria and algae are present in inner ramps; those built by red algae (rhodoids) in inner ramp, but more frequently in outer ramp settings. Micrite oncoids are common in outer ramps and basins.
Common skeletal grains in outer ramps are calcimicrobes, smaller foraminifera, sponge spicules, bryozoans, brachiopod and mollusk as well as arthropod shells, serpulids and echinoderms. Shallow inner ramp sediments exhibit similar bioclasts in comparison to the normal marine and restricted platform interior facies of platforms: Calcimicrobes, benthic smaller and larger foraminifera, calcareous green algae, mollusks, serpulids, ostracods and echinoderms. These shells are also major constituents of storm and current induced skeletal shoals in mid-ramp and innerramp settings. Organisms taking part in the formation of mud mounds, reef mounds or patch reefs are microbes, some green and red algae, sponges including archaeocyathids, stromatoporoids, coralline and siliceous sponges, tabulate and scleractinian corals, and some pelecypods (rudists). Burchette and Wright (1992) suggest that the main sediment producers both in organic buildups and in level-bottom communities have migrated from outer and mid-ramp settings to inner ramp settings since the Late Jurassic. Taphonomic and preservation criteria including shape, rounding, breakage, and size and sorting of fossils are good indications of transport and allochthonous deposition. Out-of-habitat transport of benthic organ-
Ooid accumulations originate in three ramp settings: (1) in inner ramps near shorelines and at coasts (here also contributing to eolian sand dunes), (2) in mid-ramp and inner ramp settings forming shoals and thick grainstone sheets; and (3) as grains having been transported from mid-ramps to outer ramps and deposited in thin tempestite beds intercalated with lime mudstones and wackestones, marls or shales. Most of the spherical ooids correspond to concentric or micrite ooids. Parautochthonous spherical and non-spherical radial ooids may occur in inner ramps. Relations between water energy, ooid size and sorting and the thickness of ooid cortices are indicators of transport mechanisms (see Fig. 4.25 and Fig. 4.26).
Intraclasts may occur in all ramp settings, but are abundant in distally steepened ramps (RMF 9) within debris flows accumulating near to the outer ramp slopes. These intraclasts are micritic, poorly sorted, angular to subrounded grains embedded within a micrite matrix. Sedimentary mass flow breccias derived from ramp slopes are monomict, with medium to high matrix contents and clast-supported fabrics. The decimeter- to meter-thick breccia beds have sharp boundaries. Important depositional fabrics in differentiating inner, mid- and outer ramps are bioturbation and lamination. Burrows occur all over the ramps, but are particularly abundant below the wave base and below the storm wave base. Mid-ramp and outer ramp burrows exhibit differences in density, spatial distribution and ichnotaxa association. Burrows that have not caused major breakage of shells, and burrows that have produced extensive breakage are proxies for the intensity of burrow-
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ing, which is greater in the second type that is common in outer ramp settings. Depositional lamination is frequent in deep ramp settings with hemipelagic sedimentation. Fenestral fabrics are common in tidal areas in back-ramp regions (RMF 23). Mid-ramp sediments are characterized by biofabrics indicating storm deposits (cross-bedding, imbrication patterns, reverse grading bedding; see Sect. 12.1.2.1). Good characteristics are disordered fabrics typified by chaotic arrangement of non-sorted, variously sized, commonly fragmented and broken bioclasts whose interspaces are infilled with calcisiltite or bioclastic micrite (RMF 28). Many ramp carbonates, particularly packstones and grainstones, show low compositional maturity (see Sect. 6.2.1.4) expressed by a mixture of different skeletal grains, peloids, intraclasts and sometimes also ooids. This microfacies is common in packstones and grainstones of mid-ramp- and inner ramp settings. Characteristic microfacies types are RMF 8 and RMF 14.
14.3.4.2 Microfacies Types of Carbonate Ramps Detailed studies of carbonate ramps are biased with regard to the Early Carboniferous ramps of the ‘Waulsortian facies’, and Late Mesozoic and Tertiary ramps. Many of the papers listed in Box 14.5 contain data on microfacies criteria. But relatively few investigations define microfacies types, describe their distribution in different parts of ramps, and use the microfacies types to subdivide ramp environments or evaluate the sequence stratigraphic development (e.g. Bachmann 1994; Bakush and Carozzi 1986; Rüffer 1995; Bachmann and Kuss 1998). A survey of some of the papers listed in Box 14.5 show the existence of several microfacies types that appear to dominate in specific parts of the ramp model (Box 14.6). The distribution and frequency of these microfacies types are different in low-energy, muddominated ramps that are mainly controlled by waves and tides, and in high-energy, grain-dominated ramps strongly controlled by storms. Low-energy ramps have a low, high-energy ramps a good reservoir potential. Some of these Ramp Microfacies Types (RMF) correspond in their criteria to SMF Types of carbonate platforms, but other RMF Types do not. The RMF Types should not be regarded as obligatory categories comparable with the SMF Types. The latter are better defined and are based on more case studies than the RMF Types, which only reflect a summary of the current state of the art.
Microfacies Types of Carbonate Ramps
Box 14.5. Selected case studies on Phanerozoic carbonate ramps. More references, specifically on the importance of carbonate ramps as reservoir rocks, can be found in Burchette and Wright (1992). Tertiary: Betzler et al. 1999, Buxton and Pedley 1989; Cunningham 2002; Cunningham and Collins 2002; Danese 1999; Franseen et al. 1997; Gietl 1998; Kulbrok 1996; Obrador et al. 1992; Pedley 1992; Pedley et al. 1993; Pomar 2001; Surlyk 1997 Cretaceous: Bachmann 1994; Bachmann and Kuss 1998; Caranannte et al. 2000; El Gadi and Brookfield 1999; Gomez-Perez 1998; Harris et al. 1984; Kulbrok 1996; Pascal et al. 1993; Scott and Warzeski 1993; Surlyk 1997 Jurassic: Aurell et al. 1998; Azaredo 1998; Badenas and Aurell 2001; Baumgartner and Reyle 1995; Braun 1998; Budd and Loucks 1981; Herrmann 1996; Schmid and Jonischkeit 1995 Triassic: Aigner 1984; Calvet and Tucker 1988; Calvet et al. 1990; Hips 1998; Michalik et al. 1992; Rüffer 1995; Török 1997; Tucker et al. 1993 Permian: Gérard and Buhrig 1990 Pennsylvanian: Flügel and Zhou 1986; Harris 1990 Mississippian: Ahr 1989; Bachtel and Dorobek 1998; Bakush and Carozzi 1986; Brandley and Krause 1997; Burchette et al. 1990; Caplan et al. 1996; Chatellier 1988; 1992; Chen and Webster 1994; Diaby and Carozzi 1984; Elrick et al. 1991; Faulkner 1988; Gawthorpe 1989; Gawthorpe and Gutteridge 1990; Handford 1988; Jeffery and Stanton 1996; Lasemi et al. 1998; Rankey 2003; Reid and Dorobek 1993; Sebbar et al. 2000; Simpson 1987; Sommerville and Strogen 1992; Wright 1986; Wright and Faulkner 1990 Devonian: Döring and Kazmierczak 2001; Kaufmann 1998 Silurian: Coburn 1986; Dixon and Graf 1992; Frykman 1989; Hurst and Surlyk 1983; Lavoie and Bourque 1993; Sami and Desrochers 1993; Whitaker 1988 Ordovician: Bova and Read 1987; Grimwood et al. 1999; Lavoie 1995; Lee and Kim 1992; Lee et al. 2001; Read 1980, 1982 Cambrian: Ahr 1973; Barnaby and Read 1990; Markello and Read 1981
14.3.5 Tracing Facies Zones with Microfacies Types Since common microfacies types reflect the environmental conditions controlling depositional patterns and the distribution of organisms, the assemblage of characteristic Standard Microfacies Types can be used as a proxy for specific facies belts. Fig. 14.29 shows the distribution of SMF Types in the rimmed carbonate platform model. Fig. 14.30 displays the approximate distribution of the more common microfacies types over a carbonate ramp.
Microfacies Types of Carbonate Ramps
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Box 14.6. Common microfacies types of Paleozoic and Mesozoic ramp carbonates. The list starts with microfacies types of the deeper outer ramp zones, and continues with microfacies types of mid-ramps and shallow inner ramp zones. Microfacies types of the bioclastic and oolitic carbonate sand shoals and banks formed in inner and mid-ramp settings are listed at the end of the box. Numbers in brackets indicate corresponding SMF types. Outer ramp Thin- to medium-bedded, fine-grained, often burrowed limestones and marls. Laminated marls alternate with lime mudstones. Skeletal grains commonly well preserved, not worn. Mudstones, wackestones, packstones. Some tempestite beds (grainstones). Mud mounds (see Sect. 16.1.2). RMF 1: Calcisiltite and mudstone with peloids, very fine skeletal debris, sponge spicules, sometimes fine-laminated (SMF 1). RMF 2: Argillaceous burrowed mudstone and wackestone; rare agglutinated foraminifera, ostracods, echinoderms. RMF 3: Burrowed bioclastic wackestone and packstone with diverse, common to abundant fossils (bivalves, brachiopods, echinoderms) and peloids. Skeletal grains not worn; whole fossil preservation common (SMF 8). RMF 4: Peloidal wackestone and packstone (~ SMF 2). RMF 5: Pelagic mudstone with planktonic microfossils and open-marine nektonic fossils (e.g. ammonites) (SMF 3). RMF 6: Graded, laminated and finely cross-bedded bioclastic and peloidal grainstone (tempestite). Mid-ramp Medium-bedded, fine-grained bioclastic limestones and marls, often burrowed. Skeletal grains often worn. Echinoderms common. Mudstone, wackestone, packstone, some grainstones. Distally steepened ramps exhibit slumps and various signs of reworking and transport of fine-grained ramp material. Distal steepening occurs in outer ramp settings or near the mid-ramp-outer ramp boundary. Calcareous sand shoals and sand banks (see right side). Various kinds of reefs occur in mid-ramp and inner ramp settings (e.g. coral patch reefs, coral-red algal reefs). RMF 2 , RMF 3, RMF 5 and RMF 7: Bioclastic packstone with abundant echinoderms and common bivalves and foraminifera. Skeletal grains worn (SMF 10). RMF 8: Burrowed packstone and grainstone with various skeletal grains, intraclasts, oncoids and peloids. RMF 9: Wackestone, packstone, floatstone with micritic intraclasts and ramp-derived bioclasts; sometimes microbreccias (distally steepened ramps) (~SMF 5). RMF 10: Limestone conglomerates (distally steepened). RMF 11: Marls with intraclasts and limestone pebbles (distally steepened). RMF 12: Boundstones comprising coral and coral-crust framestone, red algal framestone (SMF 7). Inner ramp Inner ramp sediments are bedded, microfacially differentiated limestones and dolomites forming relatively thin sequences. Marls are of minor importance. The inner ramp comprises open-marine environments with good water circulation, protected environments with restricted water circulation, sand shoal and bank environments characterized by oolitic and bioclastic grainstones and packstones (see below), lagoonal environments behind shoals or islands and peritidal environments. The latter are separated from other inner ramp parts as back-ramp environments (Fig. 14.3). Common texture types of open and protected inner ramps are bioclastic packstones and wackestones. Coral and coral-
ed algal boundstones (RMF 12) and other reef limestones (bivalve reefs) are common in open inner ramps. Other reefs, e.g. serpulid biostromes, occur in protected innerramp settings. Open inner ramp environments RMF 13: Bioclastic wackestone and packstone with abundant larger foraminifera (e.g. orbitolinids) (SMF 18-FOR). RMF 14: Bioclastic packstone and wackestone with skeletal grains, various amounts of intraclasts and some ooids (near-shoal). RMF 15: Bioclastic floatstone with various reef-derived material (near-reef coral-, algal- or bivalve-debris) (~SMF 6). Protected and low-energy inner ramp environments A common microfacies type in these environments is RMF 7 characterized by common to abundant echinoderm fragments. Associate skeletal grains occurring in different quantity are bivalve shells, gastropods, bryozoans, and benthic foraminifera. Other microfacies types are packstones and grainstones with larger foraminifera (RMF 13). RMF 16: Mudstone, wackestone or packstone with abundant miliolid foraminifera (SMF 18-FOR). RMF 17: Bioclastic wackestone with dasyclad green algae (SMF 18-DASY). RMF 18: Bioclastic wackestone with ostracods. Lagoonal environments RMF 17 (low-diverse flora) and RMF 19: Non-burrowed lime mudstone. RMF 20: Bioclastic wackestone and packstone with calcareous algae and benthic foraminifera. RMF 21: Oncoid packstone and floatstone (SMF 22). Peritidal zones RMF 19 and RMF 21 and RMF 22: Fine-laminated dolomitic/lime mudstone (SMF 19). RMF 23: Fenestral bindstone (SMF 21). RMF 24: Intraclast mudstone and packstone. RMF 25: Laminated evaporite-carbonate bindstone (SMF 25). Carbonate sand shoals and banks Storm-dominated ramps are characterized by accumulations of ooids, skeletal grains and peloids. These grains occurring separately or associated, form grainstone and packstone textures. Bedding and cross bedding are common. These sediments originate in mid-ramp and innerramp settings. Bioclastic banks may be formed by storminduced but also by coast-parallel bottom currents. RMF 26: Medium- and coarse-grained bioclastic grainstone and packstone with various benthic skeletal grains. RMF 27: Bioclastic grainstone and packstone composed of few dominant skeletal grains (e.g. predominantly echinoderms, or predominantly foraminifera). RMF 28: Bioclastic floatstone and rudstone exhibiting a strongly disorganized fabric. RMF 29: Ooid grainstone with densely packed concentric ooids (common in shoals and banks) (SMF 15-C). RMF 30: Shelly ooid grainstone and packstone (common in banks).
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Key for Determining SMF Types
SMF Types are valuable tools in facies analysis but do not demand too much from these virtual units. Be careful in solely using distribution patterns of SMF and RMF assemblages to recognize and trace facies zones of shelf-basin transects and • check your interpretations based on microfacies types against outcrop criteria (e.g. lithofacies, sedimentary structures), • examine the assignment to specific facies zones in the context of paleontological and palecological data, and • do not fit your data to the model instead of doing the other way round!
14.3.6 Determining Standard Microfacies Types: A Practical Guide Having avoided the pitfalls shown in Pl. 127 and discussed in Sect. 14.3.2, you are ready to determine the SMF Type of your samples. Take the following steps!
• Discriminate samples showing bindstone/lamination fabrics, breccia fabrics, shell and echinoderm concentrations, and signs of autochthonous reef growths from other samples. -- Bindstone and lamination fabrics: These fabrics characterize the SMF Types SMF 16-LAMINATED (irregular pelmicrite and pelsparite layers consisting of peloids of different size), SMF 19 (consisting of couplets of finer micritic laminae and somewhat coarser pelsparitic laminae), SMF 20 (planar or variously dome-shaped fine- and coarse-grained laminae), SMF 21 (laminated fenestral spar-filled cavities within a framework formed by sedimentary and biogenic grains; note that the fenestral fabrics also occur in limestones without lamination) and SMF 25 (alternation of fine-crystalline calcimudstone or dolomicrite layers and evaporitic layers with gypsum or anhydrite). -- Breccia fabrics are developed in SMF 4 (microbreccia consisting of skeletal grains and lithoclasts, often graded), SMF 5 (fine-grained breccia with a
Mudstone Calcisiltite Wackestone
Floatstone
Packstone
Grainstone
Rudstone
3 Abundant planktonic microfossils 23 Micrite or microsparite without fossils
5 Densely packed whole fossils and fragments of fossils, often reef-derived 8 Whole fossils, fine bioclastic micrite matrix 22 Millimeterto centimetersized agglutinated oncoids 24 Millimeterto centimetersized lithoclasts
1 Sponge spicules, often calcisiltite matrix 4 Microbreccia, small bio- and lithoclasts 5 Densely packed whole fossils and fragments, often reef- or platform-derived 10 Abraded and worn skeletal grains 16-NON-LAMINATED Very small equally-sized peloids 18 Abundant rock-building benthic foraminifera or calcareous algae 21 Spar-filled voids within a micritic or pelmicritic framework 26 Pisoids
5 Densely packed whole fossils and fragments, often reef- or platform-derived 11 Abundant coated skeletal grains 13 Millimeter- to centimeter-sized oncoids with tube-like structures 15-C Ooids with concentric structures 15-R Ooids with radial or radialconcentric structures 16-NON-LAMINATED Equally-sized peloids 17 Abundant aggregate grains 18 Abundant rockbuilding benthic foraminifera or calcareous algae
4 Microbreccia, small bio- and lithoclasts 5 Denselypacked whole fossils, often reef- or platform-derived 6 Millimeter- to centimetersized reefderived bioclasts and fossils 13 Millimeterto centimetersized oncoids with tube-like structures 24 Millimeterto centimetersized lithoclasts 26 Pisoids
2 Microbioclastic peloidal fine-grained packstone/ grainstone fabric
1 Sponge spicules, often calcisiltite matrix 3 Abundant planktonic microfossils 8 Whole fossils, fine bioclastic micrite matrix 9 Abundant fragments of fossils, bioturbation 10 Abraded and worn skeletal grains 15-M Scattered micritic ooids 22 Millimeter- to centimeter-sized agglutinated oncoids
Box 14.7. Key to the determination of Standard Microfacies Types. The key does not include reef limestones, carbonates with bindstone/lamination fabrics, and limestones exhibiting concentrations of shells or echinoderms (see text for explanation). Start with attributing your sample to texture types and continue with the search for SMF Types characterized by adequate criteria.
SMF Types of Rimmed Platform
Facies Zone (FZ) Spiculite wackestone/packstone Microbioclastic peloidal calcisiltite Pelagic mudstone/wackestone Microbreccia, bio-lithoclastic packstone Allochthonous bioclastic grainstone/ rudstone/packstone/floatstone, breccia SMF 6: Densely packed reef rudstone SMF 7: Organic boundstone, platform-margin ‘reef’ SMF 8: Whole fossil wackestone/floatstone SMF 9: Burrowed bioclastic wackestone SMF 10: Bioclastic packstone/wackestone with worn skeletal grains SMF 11: Coated bioclastic grainstone SMF 12-S: Limestone with shell concentrations SMF 12-CRIN: Limestones with crinoid concentrations SMF 13: Oncoid rudstone/grainstone SMF 14: Lag deposits (found in different facies zones) SMF 15-C: Ooid grainstone with concentric ooids SMF 15-R: Ooid grainstone with radial ooids SMF 15-M: Ooid grainstone with micritic ooids SMF 16-NON-LAMINATED: Peloidal grainstone/packstone SMF 16-LAMINATED: Peloidal bindstone SMF 17: Aggregate-grain grainstone SMF 18: Grainstone/packstone with abundant foraminifera or algae SMF 19: Densely laminated bindstone SMF 20: Laminated stromatolitic bindstone/mudstone SMF 21: Fenestral packstone/bindstone SMF 22: Oncoid floatstone/packstone SMF 23: Homogeneous, non-fossiliferous micrite SMF 24: Lithoclastic floatstone/rudstone/breccia SMF 25: Laminated evaporite-carbonate mudstone SMF 26: Pisoid cementstone/rudstone/packstone
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10
9
8
7
6
5
SMF 1: SMF 2: SMF 3: SMF 4: SMF 5:
4
3
2 1 ========= ============== ==== ==== ========= ==== ==== ====
==== ==== ==== ====
==== ==== ====
========= ======================================= ========= ==== ============== ========= B ========= B ========= ===== ===== ==== ========= ========= ========= A, B
========= A ============= A,B ========= ==== A ========= ==== A ===== ====
Fig. 14.29. Distribution of SMF Types in the Facies Zones (FZ) of the rimmed carbonate platform model. Nearly all facies belts are characterized by assemblages consisting of several SMF Types. The distribution of SMF 14 (lag deposit) is not specially indicated, because non-deposition or strongly reduced sedimentation occurs in many deep-marine as well as shallowmarine facies zones. A: evaporitic, B: brackish.
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SMF Types of a Carbonate Ramp
Inner ramp
RMF 1: RMF 2: RMF 3: RMF 4: RMF 5: RMF 6: RMF 7: RMF 8: RMF 9: RMF 10: RMF 11: RMF 12: RMF 13: RMF 14: RMF 15: RMF 16: RMF 17: RMF 18: RMF 19: RMF 20: RMF 21: RMF 22: RMF 23: RMF 24: RMF 25: RMF 26: RMF 27: RMF 28: RMF 29: RMF 30:
Peloidal calcisiltite with spicules Argillaceous burrowed mudstone/wackestone Burrowed bioclastic wacke-/packstone with diverse fossils Peloidal wackestone/packstone Pelagic mudstone Graded, laminated bioclastic grainstone Bioclastic packstone, abundant echinoderms Burrowed multigrain packstone/wackestone Wackestone/packstone/floatstone with rampderived intra- and bioclasts Limestone conglomerates Marls with intraclasts and limestone pebbles Boundstone Wackestone/packstone with larger foraminifera Packstone/grainstone with various bioclasts Bioclastic floatstone, reef-derived material Mudstone/wackestone with miliolids Bioclastic wackestone with dasyclads Bioclastic wackestone with ostracods Non-burrowed lime mudstone Bioclastic wackestone/packstone with algae and benthic foraminifera Oncoid packstone/floatstone Finely laminated dolomitic or lime mudstone Fenestral bindstone Intraclast mudstone/wackestone Laminated evaporite-carbonate bindstone Bioclastic grainstone/packstone with diverse skeletal grains Bioclastic grainstone/packstone with a few dominant skeletal grains Bioclastic grain-/floatstone, disorganized fabric Ooid grainstone with concentric ooids Shelly ooid grainstone and packstone
Peritidal Lagoon Shoal Restrict. Openmarine
Midramp
Outer ramp
Basin
============ ================== ============ ====== ================= ====== ================== ====== ============ ============ ============ =========== ============ ====== ====== ====== ====== ======
====== ============ ====== ============ ====== ====== ====== ====== ====== ====== ====== ====== ======
Fig. 14.30. Strongly generalized distribution of microfacies types in different parts of a homoclinal carbonate ramp. This is not a ‘model’ and the Ramp Microfacies Types do not have the character of Standard Microfacies Types. The figure just shows some microfacies types frequently described in studies dealing with Paleozoic and Mesozoic carbonate ramps. RMF 9, RMF 10 and RMF 11 occur predominantly in distally steepened ramps. See text for the correlation of some of these RMF Types with SMF Types of platform carbonates.
Determining SMF Types: Advices
high percentage of reef-derived material), SMF 6 (reef talus, fossils and lithoclasts) and SMF 24 (consisting of coarse micrite clasts and laminated clasts, ‘flat pebble conglomerate’). All these SMF types also develop other textures and are, therefore, also included in Box 14.7. -- Concentrations of shells and echinoderms (mostly crinoids): Shell beds and coquinas are represented by SMF 12. The fossils form packstone, rudstone and floatstone fabrics. The type needs further subdivision. -- Reef limestones are assembled in SMF 7 characterized by sessile reefbuilders and framestone, bafflestone or bindstone fabrics. These limestones require specific investigations of the role of organisms in the formation of reefs, growth fabrics, internal sedimentation and cement types (see Sect.16.2). • Distinguish remaining samples according to textural types using the Dunham classification. • Use Box 14.7 to differentiate the samples, considering common and abundant grain types. Note that samples with identical grain composition can show different depositional textures. • Compare your favorite SMF designation with the definition of the SMF Type and with the pictures of the examples. But note that the pictures can not show the whole spectrum of microfacies variations. • Check the occurrence of the SMF Type in the context of the rimmed platform facies model (Fig. 14.29) and think about the reliability of attributing your sample via the SMF classification to a specific facies zone. If you find that your sample can not have originated in the facies zone indicated by the model, have faith in your interpretation and consider the following possibilities: -- The composition and texture may not reflect criteria caused by autochthonous deposition, but rather criteria caused by transport from the original depositional site to other facies belts. Oolite beds intercalated within basinal sediments and representing turbidites or debris flow deposits may display all the characteristic criteria of SMF 15. Concentrations of shells (SMF 12) are often allochthonous deposits formed far from the living sites of the organisms. Storms can transport skeletal grains in offshore and onshore directions, and form multigrain bio- and lithoclastic calcarenites distant from the carbonateproducing area. -- Different mixtures of various skeletal grains, coated grains, composite grains, peloids and intrac-
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lasts may create problems in the accurate classification of SMF Types. The SMF Types SMF 8, SMF 9, SMF 10 and SMF 11 often exhibit a multigrain composition without a pronounced quantitative dominance of a specific grain category. Your problem may be related to this shortcoming in the SMF classification. -- Finally, your assignment may not fit because you have relied on the wrong depositional model. Consider the possibilities of the carbonate ramp model (Fig. 14.3). Perhaps you will find the microfacies characteristics of your sample in one of the RMF Types listed in Sect. 14.3.4.2 and summarized with regard to their more common distribution in Fig. 14.30. Note that several of these RMF Types correspond in their composition and texture to SMF Types, but that the distributional pattern of some of these corresponding types differs from that on carbonate platforms.
14.4 Dynamic Microfacies Types and Environmental Changes The concept of microfacies types includes a strong tendency to categorize. Strict categorization, however, may obscure the question of whether or not the microfacies data set consists of components that vary in a continuous fashion, or of discontinuous grain associations. If the spatial distribution is discontinuous, the stratigraphic or geographical representation of various microfacies types as a mosaic of ‘cells’ is justified and can provide a basis for the interpretation of sedimentary environments. This approach is followed by the SMF concept. However, if the distribution is continuous, the facies mosaic gives only a rough picture of nature. This is specifically important in the facies interpretation of carbonate ramps where environmental changes are progressive and gradual rather than abrupt. Most grain analyses are descriptive with little analysis of relationships in grain composition or grain abundance of samples. One way to detect relationships and patterns in grain-type distributions are ordination methods that explore relationships between the components and detect and analyze relays in carbonate rocks. Relays are spatial or temporal systematic shifts in the relative importance of component grains answering to unidirectional changes of major environmental conditions (Hennebert and Lees 1985). The application of ordination methods to the study of relay phenonema facilitates the recognition of environmental gradients (Hennebert and Lees 1991).
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Fig. 14.31. Microfacies relays. A: Traditional microfacies types based on the abundance of grain types are generally considered as discrete units caused by a discontinuous distribution of the grain types across depositional environments or within a stratigraphic section. Differences in grain abundance are used to distinguish facies which seem to be genetically unrelated. B: The relay method shows the distribution and relative abundance of grain types formed in response to a unidirectional environmental gradient, e.g. water depth or decreasing water energy in shallowing-upward sequences. After Spence and Tucker (1999).
Relays reflecting environmentally controlled compositional changes document the effects of relative third-order sea-level fluctuations and can, therefore, successfully be used in identifying system tracts and hidden sequence boundaries and evaluating microfacies types (Spence and Tucker 1999; Fig. 14.31). Ordination methods use coefficients giving a measure of similarity in sample composition or in grain distributions. Relays are detected using computer-optimized similarity coefficient matrices (e.g. the Jaccard similarity coefficient) based on presence/absence compositional data or applying correspondence analysis techniques (Hennebert and Lees 1985). The latter method can consider absolute frequencies, relative frequencies (percentages), ranked values (e.g. low, medium, high) and binary data (presence/absence) as well.
Dynamic Microfacies Types
Basics: Standard microfacies Bonham-Carter, G.F., Gradstein, F.M., D’Iorio, M.A. (1986): Distribution of Cenozoic foraminifera from the northwestern Atlantic margin analyzed by correspondence analysis. – Computer and Geoscience, 12, 621-635 Carozzi, A.V. (1989): Carbonate rock depositional models. A microfacies approach. – 604 pp., Englewood Cliffs (Prentice Hall) Cheetham, A.H., Hazel, E.J. (1969): Binary (presence-absence) similarity coefficients. – Journal of Paleontology, 43, 1130-1136 Flügel, E. (1972): Mikrofazielle Untersuchungen in der alpinen Trias: Methoden und Probleme. – Mitteilungen der Gesellschaft der Geologie- und Bergbaustudenten in Österreich, 21, 6-64 Hennebert, M., Lees, A. (1985): Optimized similarity matrices applied to the study of carbonate rocks. – Geological Journal, 20, 123-131 Hennebert, M., Lees, A. (1991): Environmental gradients in carbonate sediments and rocks detected by correspondence analysis: examples from the Recent of Norway and the Dinantian of southwest England. – Sedimentology, 38, 623-642 Koopman, B.J., Lees, A., Piessens, P., Sarnthein, M. (1979): Skeletal carbonate sands and wind-derived silty marls off the Sahara coast: Baie de Lévrier, Arguin platform, Mauritania. – Meteor Forschungsergebnisse, C, 30, 15-57 Spence, G.H., Tucker, M. (1999): Modelling carbonate microfacies in the context of high-frequency dynamic relative sea-level and environmental changes. – Journal of Sedimentary Research, 69, 947-961 Wilson, J.L. (1975): Carbonate facies in geologic history. – 471 pp., Berlin (Springer) Further reading: K148, K190
Fig. 14.32. Exercise: Try to determine the SMF Type for this sample using the Guide in Sect. 14.3.6. How old is the limestone (see Chap. 10) and from which environment does it come? Answers can be found in the Appendix.
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15 Basin Analysis: Recognizing Depositional Settings
This chapter is an overview of tools for translating microfacies data into depositional settings. Using case studies of non-marine and marine carbonate rocks, we focus on the diagnostic criteria for recognizing the facies characteristics needed in basin analysis studies. Microfacies is an important element of basin analysis, but must be applied along with all the other field and laboratory observations available for reconstructing depositional conditions. This is exemplified by Fig. 15.1.
15.1 Pedogenic Carbonates Pedogenic carbonates are formed by the accumulation of CaCO3 within carbonate-rich soils. Ancient pedogenic carbonates are studied by means of microscopy, stable isotopes and geochemistry. The term paleosol (or paleosoil) refers to a soil formed on a landscape of the past. Calcite accumulations in terrestrial, near-surface soil horizons, formed in semiarid climate under conditions of sparse rainfall, are called caliche
Fig. 15.1. At first sight this sample would be referred to the Standard Microfacies Type SMF 10, indicating a shallow-marine depositional environment. But the sample is a sandy grainstone formed in eolian dunes. The highly porous sediment consists predominantly of chips of coralline red algae (all black grains), terrigenous quartz grains (Q), bryozoans (B) and bivalve shells (BS), most of which are completely leached and only preserved by the circumcrusts of micrite rims forming cortoids. Additional skeletal grains are echinoid spines (E), gastropods (GA) and miliolid foraminifera (F). The association of coralline red algae, bryozoans and echinoids is typical for shallow-marine carbonates, deposited in temperate seas. But what criteria characterize the sample as an eolian deposit? Diagnostic features are the narrow size range of the grains, good sorting, winnowing out of fine particles, high grain diversity (as compared to marine oolites), weak cementation (see Pl. 29/ 7, Pl. 33/2) and cross-bedding observed in the outcrops. Small-scale cross-bedding is indicated by the white lines. This rock was used by the Punians in ancient Carthage for building stones (E. Flügel and Ch. Flügel 1997). Pleistocene (Tyrrhenian): El Haouria, Cap Bon Peninsula, northeastern Tunisia. E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_15, © Springer-Verlag Berlin Heidelberg 2010
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or calcrete. Caliche has a variety of origins ranging from a purely pedogenic formation to cementation in the phreatic zone. Calichefication modifies the original sedimentary features of carbonate rocks and produces new textures and structures (see Sect. 2.4.1.1 for further information). Caliche originates via a combination of cementation, replacement and displacement, leading to the substitution of the host material. If time permits, a calciterich hard horizon (hardpan layer) will develop, acting as a barrier to further movement of calcium carbonate and favoring the formation of laminar units on the impermeable horizon. The horizon eventually fractures and brecciates, leading to the formation of distinct textures, often associated with pisoids. Paleocaliche are typically micritic deposits that can be laminated or non-laminated. Caliche profiles developing from shallow-marine limestones often exhibit a sequence of massive/brecciated and laminar/pisolitic horizons.
15.1.1 Microfacies Criteria of Paleocaliche A well-illustrated review of micromorphological criteria of caliche was provided by Esteban and Klappa (1983). Note that descriptive terms used by geologists and pedologists differ (Ettensohn et al. 1988)! Caliche exhibits microfabrics of non-biogenic and biogenic origin. The latter, summarized as beta calcrete fabrics (Wright 1990), includes microbial coatings, calcified tubules, rhizoliths, alveolar-septal fabrics, and Microcodium. Microbial and fungal contributions are evident in many calcretes (Verrecchia 2001). Alpha caliche fabrics originate from the replacive growth of carbonate (e.g. dense micrite, microspar, nodules, complex cracks and crystal-filled fractures, rhombic calcite crystals). Some of the criteria listed below reflect the development of soil profiles over time, e.g. an open micrite network can develop to dense laminar secondary lime mudstone, and then to brecciated caliche. Common thin-section criteria are: (1) Texture • Carbonate texture. The general texture of paleocaliche corresponds to lime mudstones, packstones or bindstones. Note that these textures are no depositional but secondary textures! (2) Matrix and constituents • Dense micrite (Pl. 128/5; Fig. 15.2) concentrated either into dark isolated, small subspherical nodules, or
Criteria of Paleocaliche
into ring-like or irregular masses forming a secondary lime mudstone developing from micritic parent limestone. The often used term ‘glaebule’ refers to constituents of soils (e.g. nodules) whose shape, color or composition differ from the surrounding soil matrix. • Clotted micrite consists of rounded to angular micrite peloids, separated by blocky calcite. The clotted texture can lead to total obliteration of the original sedimentary textures and fabrics. A grainstone composed of well-sorted, sand-sized grains may be altered to clotted micrite with peloids (diameter from 0.5 to 1.0 mm). • Micritic peloids. Spherical particles, 0.05–0.5 mm in size. Occur in great abundance, isolated or as densely packed aggregates forming a clotted texture. Interpreted as being of near-surface diagenetic origin (‘grainification’ of a lime mudstone), or as calcified fecal pellets of small soil arthropods. Peloidal and coated fabrics have also been interpreted as calcification products related to root mats (Calvet and Julia 1983). • Micrite coatings (Fig. 15/2; Pl. 128/6). Sand-sized to pebble-sized detrital carbonate grains and bioclasts coated with thin regular to highly irregular laminae of brownish or yellowish micrite, microsparite or fine sparite. Coatings vary laterally in thickness or only occur on one side of the grain. Coated grains may be connected by micritic bridges. Micrite coatings are particularly common in limestone-hosted caliche characterizing initial phases of calichefication. • Pisiform accretionary bodies of micrite or microsparite with a nucleus of carbonate silt. Associated with laminar caliche horizons. • Pisoids. Soil pisoids (Pl. 128/6; Fig. 4.27) are sphaerical or subsphaerical, irregularly laminated concentric micritic grains, formed by chemical or biochemical precipitation of microcrystalline carbonate on a nucleus (Fig. 14.27B; Pl. 126/4). See Sect. 4.2.6 for more information. Commonly associated with laminar caliche crusts. (3) Fabric • Laminar crusts (Pl. 128/1, 6). Laminated or banded millimeter-thick crusts of slightly undulated carbonate laminae alternating in color (yellowish/grayish) and composition (micrite, microsparite, sparite). Lamina thickness varies between 50 µm and about 500 µm. The laminae may form a complex network. The laminae follow the topography of the substrate. See Read (1976) for criteria distinguishing laminar caliche from microbial laminated structures. • Circumgranular cracking (Fig. 15.2; Pl. 128/2, 3). Micrite nodules, up to 10 mm, bounded and locally dissected by crescentic and circular veins of sparite, up to 0.5 mm thick. Formed by alternating shrinkage
Caliche: Root-Related Structures
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Fig. 15.2. Diagnostic features of paleocaliche. One of the most significant features of caliche seen in thin sections is their inhomogeneous texture and the lack of fossils. Characteristic criteria shown in this picture are the breccia-like texture, nodular fabrics (N), clotted fabrics with micritic peloids (PEL), grains coated by variously thick micrite (MC), and spar-filled wrinkled microcracks (white arrows). The latter occur as straight or wrinkled cracks, or surround the nodules (circumgranular cracking). The coatings surround alveolar structures (A; irregular pores infilled with sparite and separated by interconnecting micritic bridges; interpreted as root structures). Black arrows point to rhizogenic structures. The rock is strongly fractured. Early Tertiary (Ilerdian, Paleocene): Serraduy, Tremp Basin, Southern Pyrenees, Spain. After Eichenseer and Luterbacher (1992).
and expansion processes in soils induced by seasonal drying/wetting cycles. (4) Root-related features Rhizoliths, Microcodium, calcified root hairs, pedotubules, and alveolar structures can be combined under the term ‘root-related structures’. Examples of carbonate rocks with evidence of the former activity of plant roots forming ‘rhizogenic calcretes’ (Wright et al. 1995) are known from various Phanerozoic time intervals, particularly from the Late Paleozoic, Jurassic, Cretaceous and from the Tertiary and Quaternary. • Root structures (rhizoliths) occur as root casts or rhizoconcretions. Rhizogenic calcretes are formed by the calcification on, in, or around roots of higher plants. Root casts consist of tubular branched voids (pedotubules; Pl. 128/5) of decayed roots infilled with fine sparite. They are generally less than 1 mm in diameter. Rhizoconcretions are characterized by concentric accumulations of tangential to random fibers of micrite. They are typically 0.5 to 2 mm in diameter, often surrounding root casts or alveolar structures. See Klappa (1980) for a classification of root structures.
• Microcodium (Pl. 128/4). Aggregates of 0.1–1 mm long prismatic petal-like sparry calcite crystals, with a central cavity in each one. The cavity is filled with dark material. The calcite crystals are radially arranged around an axis, or are grouped into irregular masses and crusts. Interpreted as calcified mycorrhizae, a symbiotic association between soil fungi and cortical cells of plant roots, or as bacteria or algae (see Sect. 10.2.6.2). • Alveolar structures (Fig. 15.2). Tubular to irregular pores infilled with sparite and separated by straight or various sinuous interconnecting bridges of micrite. The bridges consist of tangentially or randomly oriented micrite and/or microsparite. The pores may contain arcuate micritic septae (alveolar-septal structure). Individual pores reach 1 mm in width and 20 mm in length. Interpreted as entwined rhizoliths, resulting from the calcification of root hairs. (5) Breccia structures • In-situ brecciation (Fig. 15.2, Pl. 128/6). Micrite fragmented in place, forming a polygonal fracture network infilled with sparite. Infilled micrite polygons are millimeters to several centimeters in size. Brecciation
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is induced by desiccation, displacive crystallization of calcite, root activity and/or extensive periods of dissolution. (6) Carbonate cement • Cement and microsparite. Common cement types in caliche are vadose micritic cement, blocky calcite cement and meniscus cement. Microsparry fabrics are common. 15.1.2 Significance of Paleocaliche and Paleosols Paleosols have been known since the Precambrian. Most case studies deal with examples from the Quaternary and Tertiary, but other papers deal with the Cretaceous, Jurassic, and Late Paleozoic. The recognition of paleosols and paleocaliche is of paramount importance in interpretating past climate (Mack and James 1994; Theriault and Desrochers 1994), the evolution of the atmosphere (Cerling 1991; Ekart et al. 1999), and weathering (Martini and Chesworth 1992). Most geological studies of paleosols focus on caliche. Paleocaliche is hosted by sandstones, alluvial
Microfacies of Paleocaliche
sediments, and lacustrine carbonates and shallow-marine carbonates. Different host sediments cause different associations of calcrete profiles, horizons and associations (Theriault and Desrochers 1994). Since paleosols can form on limestones of marine origin only when the relative sea level is lowered, paleosols are key criteria for bounding surfaces in sequence stratigraphy. The combined use of paleocaliche and paleokarst criteria gives us a detailed picture of the emergence evolution of platforms (Davies 1991; Webb 1994). Basics: Pedogenic carbonates and paleocaliche Bain, R.J., Foos, A.M. (1993): Carbonate microfabrics related to subaerial exposure and paleosol formation. – In: Rhezak, R., Lavoie, D.L. (eds.): Carbonate microfabrics. – 19-27, New York (Springer) Bullock, P., Fedoroff, N., Jongerius, A., Stoops, G., Tursina, G., Babel, U. (1985): Handbook of soil thin section description. – 152 pp., Wolverhampton (Waine Research Publications) Davies, J.R. (1991): Karstification and pedogenesis on a late Dinantian carbonate platform, Anglesey, North Wales. – Proceedings of the Yorkshire Geological Society, 48, 297321 Esteban, M., Klappa, C.F. (1983): Subaerial exposure environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American
Plate 128 Carbonate Paleocaliche and Paleosols: Microfacies Criteria The plate displays a few diagnostic criteria including Microcodium (–> 4), laminar crusts (–> 1, 6), micritic nodular structures (–> 2, 3, 6), spar-filled branching pedotubules (–> 5), circumgranular cracks (–> 3) and pisoids (–> 6). 1 Laminar paleocaliche crust. The foraminiferal grainstone, formed in a platform interior environment, is disconformably overlain by a pedogenic crust consisting of light sparry and dark micritic, slightly undulated laminae. The caliche crust represents a sequence boundary. Early Tertiary: Southern border of the Galala Basin, Egypt. 2 Nodular paleocaliche formed on marginal carbonates of an episodically drying lake. The lacustrine carbonates were pedogenically modified, resulting in the formation of palustrine limestones (see Sect. 2.4.1.2). Caliche criteria are circumgranular cracks (white arrows) around nodules (N) consisting of dense micrite. The black arrow points to sparfilled shrinkage cracks. Evidence of the lacustrine host rock are charophyte algae (CA). Early Tertiary (Claron Formation, Paleocene): Brice Canyon National Park, Utah, U.S.A. 3 Circumgranular cracking around a quartz-bearing nodule caused by alternating shrinkage and expansion processes in soils. The circumgranular crack is filled with calcite. Same locality as –> 2. 4 Microcodium Glück, characterized by elongate, radiating and petal-like calcitic elements grouped into ring-like structures. The structure is indicative of a meteorically influenced environment, e.g. a paleosol, and hence is a criterion for the recognizing continental conditions, cessation of marine sedimentation, and subaerial exposure. Findings of isolated Microcodium within marine strata may be explained by the reworking of paleosols. Microcodium is abundant in many Late Pennsylvanian to Permian, Late Cretaceous, Paleocene, Eocene, and Miocene paleosols. Pleistocene: Crete Island, Greece. 5 Caliche nodule containing irregularly bent and branching voids infilled with fine sparite (pedotubules, arrows) corresponding to decayed root elements of higher plants. Early Cretaceous: Subsurface, Ostermünchen well, southern Bavaria, Germany. 6 Paleocaliche. The sample exhibits several criteria that are diagnostic for alpha fabric caliche: Clotted and inhomogeneous micrite, irregular nodular texture (N), fracturing and incipient brecciations (white arrows), undulating laminar crusts (black arrow), and pisoid grains (P). Note the difference in the size, shape and internal lamination of these grains as well as their irregular grading. Most of the pisoids consist of microspar and exhibit numerous fine pores. SMF 26. Quaternary: Agadir, Morocco.
Plate 128: Paleocaliche and Paleosols
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Association of Petroleum Geologists, Memoir, 33, 1-54 Freytet, P., Plaziat, J.-C. (1982): Continental carbonate sedimentation and pedogenesis. Late Cretaceous and Early Tertiary of southern France. – Contributions to Sedimentology, 12, 1-213 Klappa, C.F. (1980): Rhizoliths in terrestrial carbonates: classification, recognition, genesis and significance. – Sedimentology, 27, 613-629 Klappa, C.F. (1983): A process-response model for the formation of pedogenic calcretes. –In: Wilson, R.C.L. (ed.): Residual deposits. – Geological Society, London, Special Publication, 11, 211-220 Gebhardt, U. (1988): Mikrofaziesanalyse und stratigraphischregionale Interpretation terrestrischer Karbonate der variszischen Molasse (Mitteleuropa, Permokarbon). – Freiberger Forschungshefte, C 427, 30-59 Reinhardt, J., Sigleo, W.R. (eds., 1988): Paleosols and weathering through geologic time: principles and application. – Geological Society of America, Special Paper, 216, 181 pp. Read, J.F. (1976): Calcretes and their distinction from stromatolites. – In: Walter, M.R. (ed.): Stromatolites. – 5571, New York (Elsevier) Retallak, G.J. (2001): Soils of the past. Second edition. – 416 pp., Oxford (Blackwell Sciences) Theriault, P., Desrochers, A. (1993): Carboniferous calcretes in the Canadian Arctic. – Sedimentology, 40, 449-465 Wright, V.P. (ed., 1986): Paleosols. Their recognition and interpretation. – 315 pp., Oxford (Blackwell) Wright, V.P. (1994): Paleosols in shallow marine carbonate sequences. – Earth Science Review, 35, 367-395 Wright, V.P., Tucker, M.E. (eds., 1991): Calcretes. – International Association of Sedimentologists, Reprint Series, 2, 352 pp. Wright, V.P., Platt, N.H., Marriott, S.B., Beck, V.H. (1995): A classification of rhizogenic (root-formed) calcretes, with examples from Upper Jurassic-Lower Cretaceous of Spain and Upper Cretaceous of southern France. – Sedimentary Geology, 100, 143-158 Further reading: K030
Paleokarst and Speleothems
sols, discolored rocks, terra rossa, breccias, and specific cavity fills. The recognition of paleokarst-related disconformities has major implications for the interpreting sea-level history and defining a sequence-stratigraphic framework. One of the most spectacular features of subsurface karst is the development of cave systems and voids originating from meteoric dissolution. These caves are places where the formation of various types of carbonates (speleothems) and the precipitation of aragonite and/or calcite in smaller voids (Pl. 129/4) take place. Many caves exhibit infilling by cave sediments (Pl. 129/ 2, 3) whose composition and fabric provide clues to climate and hydrological conditions. A major feature of surface and subsurface paleokarst are intraformational karst breccias (Fig. 15.4; Pl. 129/1), which may be hosts for mineral deposits, particularly lead-zincsulfide deposits (see Sect. 17.2). Paleokarst develops as depositional, local or interregional paleokarst (see case studies in Choquette and James 1988). Depositional paleokarst is commonly associated with meter-scale carbonate platform cycles. Local paleokarst develops when only parts of platforms (e.g. platform rims) are exposed, due to synsedimentary block faulting or small drops in sea level. This is documented by the case study shown in Pl. 129. Interregional paleokarst is related to major eustatic-tectonic events and characterized by widespread karst terranes. Distribution: Paleokarst features have been reported since the Archean, but became abundant in carbonate
15.2 Paleokarst and Ancient Speleothems Paleokarst is an ancient karst (see Sect. 2.4.1.3), commonly buried by younger sediments or sedimentary rocks. Speleothems are carbonate sediments formed in cave systems. Karst is an aspect of lowstand conditions in the carbonate system. Humid climate favors karstification of exposed carbonate platforms during sea level lowstands. Rainwater charged with CO3 penetrates fissures and voids creating a karst landscape with collapse and solution sinkholes, solution caves, dolines and underground rivers. Karst features indicative of subaerially exposed karst are dolines, rundkarren (rounded solution tunnels), kamnitzas (solution basins with circular to oval outlines, flat bottoms and rounded edges), paleo-
Fig. 15.3. Paleokarst breccias caused by the collapse of cave rocks and speleothem carbonates are characterized by a disorganized fabric, and clast-supported angular limestone clasts in a size range of a few millimeters up to boulders more than several tens of centimeters across. The thin section shows a breccia consisting of angular clasts of pisolitic and fine-laminated speleothem calcite of a paleokarst cave. Early Carboniferous (Viséan): Holy Cross Mountains, Poland.
Paleokarst: Criteria
rocks from the Upper Proterozoic onward (Bosak et al. 1989). Large-scale products of subaerial exposure have remained rather similar since the Mid-Proterozoic. However, the amount of karst formation per time unit should have been lower during the Cambro-Devonian, owing to the absence of terrestrial vegetation cover prior to the Carboniferous.
15.2.1 Diagnostic Criteria of Paleokarst and Paleospeleothem Structures Common features of paleokarst include geomorphicstratigraphic characteristics, field-scale observations and microscopic criteria (Choquette and James 1988; Esteban and Klappa 1983; Wright 1991). The study of paleokarst requires observations at outcrop, macroscopic and microscopic scales as well as geochemical studies. Valuable criteria include breccias, dikes and fissures, cave infill sediments, speleothems, carbonate cements, and silcretes and calcretes. • Breccia: Collapse breccia (Pl. 129/1; Fig. 15.3) formed in ancient caves are clast-supported mega- and micro-breccias consisting of angular, sometimes also subrounded millimeter- to tens of centimeter-sized limestone clasts. Sorting is generally poor. Fitting may be good or poor. The matrix is calcareous or siliciclastic depending on the lithology of the host rock and that of the overlying beds. Karst collapse breccia should not be confused with breccia produced by interstratal dissolution of interbedded sulfates (see Sect. 5.3.3.3). • Dikes and fissures: Filled and cemented fissures (Pl. 129/1) with widths ranging from millimeters to decimeters are common features of paleokarst. Fissures may be cemented by several generations of isopachous and granular coarse calcite cement and filled with internal sediment, consisting of crystal silt, and clasts derived from cave walls as well as from overlying beds. Corrugated marine radial fibrous cement crystals may indicate karst phases (Aubrecht 1997). Fissure filling occurs subaerially by continental sediments or, subsequent to marine transgressions, by marine sediments yielding characteristic cavity-dwelling microfossils (e.g. ostracods). The fossils within the marine infill of fissures and solution cavities are of enormous importance in recognizing the age and timing of emergence and karstification. • Cave infill sediments: Caves can function as giant sediment traps accumulating various sediments and debris. Common structures are (a) lamination (sub-millimeter densely and widely spaced, light and dark parallel laminae; fining-upward; laminations of locally de-
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formed laminae; Pl. 129/2, 3, 5), (b) microbreccias (consisting of angular rock clasts and broken crystals embedded within a microgranular matrix; Pl. 129/1), (c) coarse redeposited calcite crystals and fine crystal silt, and (d) internal vadose sediments as partial fillings on floors of cavities or as matrix sediment in collapse breccias; often grayish-yellow, silty and laminated. • Speleothems are among the more useful features for recognizing karstification (Dreybrodt 1988; Frisia et al. 2000). The surface of cave bedrocks may be coated with white layers consisting of coarse sparry calcite and explained as precipitation from phreatic conduits. Distinctive criteria of paleospeleothems are (a) laminar flowstones consisting of several cement layers, growing directly on host rocks and associated with geopetal vadose internal sediment, (b) dendroid calcite sinter crusts forming stalagmite and stalactite structures (Fig. 15.4), (c) fibrous calcite cement crystals developed as optically length-slow ‘coconut-meat calcite’ (Folk and Assereto 1976) or palisade calcite, often as radiaxial crystals, (d) helictites (excentriques) in fissures (Reinhold 1998), (e) alterations of large and small crystal layers, and (f) coarse blocky calcite, occluding voids. • Carbonate cements associated with local paleokarst were studied by Dunham (1969), Steinen and Matthews (1973) and Schönlaub et al. (1991), those of interregional paleokarst by Meyers (1974, 1978, 1988). Most
Fig. 15.4. Speleothem carbonate. Fingerlike, radiating, cmthick sinter crusts lining tens of cm-sized cavities within a highly porous limestone characterized by ‘sickle cell’ structures (see Fig. 15.6). The speleothem structures are directed upward here, but downward and sideward growth occurs as well. The crusts consist of layers of fibrous calcite. Interframework pores are filled by stalactitic cement followed by drusy calcite mosaic. The sinter crusts formed during subaerial exposure of lacustrine carbonates related to a lake level lowstand phase. Late Tertiary (Miocene): Wallerstein, Nördlingen Ries, Germany. From Pache et al. (2001).
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Paleokarst: Case Study
Plate 129 Paleokarst and Cave Carbonates: An Early Carboniferous Case Study The plate exhibits samples documenting microfacies criteria of a paleokarst breccia (–> 1), cave sediment infillings (–> 2, 3, 5) and relicts of calcite speleothems developed within caves (–> 4). Extensive subsurface karstification of Late Devonian limestones took place during the Early Carboniferous in the central Carnic Alps, a part of the Southern Alps (Schönlaub et al. 1991) as a result of drops in sea level during the Late Tournaisian and synsedimentary block faulting. The paleokarst is associated with an unconformity separating Famennian to Early Carboniferous limestones from overlying siliciclastic sediments (Hochwipfel Formation). The Famennian to Early Carboniferous sections comprise pelagic limestones and shelf limestones, separated from the Hochwipfel Formation by a disorganized collapse breccia (–> 1). The breccia consists of angular, mm to 0.5 m-sized clast-supported and matrix-supported clasts exhibiting different microfacies types. The succession of events documented by field data, microfacies and geochemical signals is shown in Fig. 15.5B. Polymetallic strata-bound sulfide ore deposits are thought to be genetically related to the Late Devonian paleorelief and connected with the process of karstification. 1 Paleokarst collapse breccia formed by subaerial dissolution leading to the development of a cave system. The breccia consists of angular and subrounded limestone and chert clasts (C). The limestone clasts are Famennian fenestral mudstones (FM) and nodular pelagic wackestones (NW). The Early Carboniferous fine-grained matrix is a microbreccia (MB) composed of detrital quartz, sandstone, shale clasts, and cherts. The large limestone clast at the right bears a fissure infilled with radiaxial cements (arrow), pointing to a pre-breccia formation of fissures associated with synsedimentary block faulting processes. Cava di Marmo southeast of the Plöcken Pass, near the Italian/Austrian boundary. 2 Densely laminated internal sediment within karst cavities. The parallel siltstone laminae consist of quartz, calcite and illite/muscovite. Arrows point to rock debris. Each couplet of dark and light laminae originate from one flooding of the cavities (Ford 1985). Early Carboniferous (Middle Dinantian): Malpasso quarry, central Carnic Alps, Italy. 3 Laminated internal sediment composed of calcite, authigenic quartz and mica deposited in karst cavities. Note parallel planar bedding and flame and load structures (arrows). Cava Val di Collina west of the Plöcken Pass, Carnic Alps, Italy. 4 Euhedral calcite crystals lining the cave walls (speleothems). Cathodoluminescence criteria indicate that these nonluminescent cement crystals were formed in a freshwater phreatic environment. Same locality as –> 3. 5 Cave sediment infilling in paleokarst cavities. Laminated filling between fractured scalenohedral calcite crystals formed on the walls of the karst cavities. The spar calcite coating predates the densely laminated infilling. Same locality as –> 3. –> 2–5: Schönlaub et al. (1991) Sedimentary Event
Diagenetic environment
Transgression of the Hochwipfel Fm. marine phreatic
Shale
Formation of collapse breccia – Pl. 129/1
Clasts Cephalopod-trilobite wackestone Fenestral limestone Black chert Brachiopod-crinoid limestone
First generation of isopachous radial cements – Pl. 129/1
V
Bedded sandstone
freshwater phreatic
Erosion of crystal silt marine phreatic
Formation of cave sediment – Pl. 129/2, 3, 5
Matrix Siliciclastic, derived from overlying Hochwipfel Formation
Formation of caves with calcite coating – Pl. 129/4, 5
freshwater phreatic
Formation of fissures Ore concentration Local silcretes and cherts
Fenestral limestone
surface karst/vadose
Extensive karst erosion and limestone solution
A
B
Limestone sedimentation up to Tournaisian-Viséan boundary
marine phreatic
Fig. 15.5. Case study ‘paleokarst’. Timing of sedimentary events and changes in the diagenetic environments in the Early Carboniferous succession of the central Carnic Alps, Austria/Italy. Numbers refer to figures in Pl. 129. Modified from Schönlaub (1991).
Plate 129: Paleokarst
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cements beneath unconformities seem to be of meteoric-phreatic or deeper-burial origin. Shallow phreatic cements are clear, iron-poor, blocky or sparry calcite (Pl. 129/4, 5), commonly non-luminescent with thin zones of bright luminescence. Burial cements are ironrich and commonly inclusion-rich blocky calcite with dull, variably zoned cathodoluminescence. • Calcretes and silcretes: Many paleokarst surfaces are overlain by calcretes. Siliceous breccias and layers of white weathered cherts on top of surface karst have been interpreted as silcretes (Mustard and Donaldson 1990; Schönlaub et al. 1991). Diagnostic criteria of significant dissolution and successive pervasive silica cementation are (a) abundant micropores, (b) needle fiber fabrics, (c) a network of anastomosing walls, and (d) composite void fillings by euhedrally terminated megaquartz, radiating chalcedony and coarse mosaic quartz in the center of voids.
15.2.2 Significance of Paleokarst and Cave Carbonates Karstification may affect various types of carbonate buildups, including carbonate platforms, reefs, carbonate ramps and pelagic swells. Paleokarst structures assist in • recognizing unconformities, subaerial exposure and stratigraphic breaks (Mussman et al. 1988; Vanstone 1998; Haas 1999; Turgeon and Lundberg 2001), • correlating stratigraphic units (Charcosset et al. 2000), • evaluating sea-level fluctuations (Satterley et al. 1994; Calner 2002; compare the case study, Pl. 129), • identifying paleoclimates. Paleokarst, used in conjunction with other subaerial exposure features (specifically paleosols of humid and semiarid phases) to be distinguished can be critically compared with modern karst phenonema, to provide an excellent indicator of paleoclimatic conditions (Jimenez de Cisneros et al. 1993; Wright 1980, 1982, 1988; Burns et al. 1998). However, because of the complex history of subaerial weathering processes, paleo-data regarded as indicators should be analyzed within a detailed microstratigraphic setting. High-resolution of paleoclimatic variations are possible by comparing annual variation patterns in speleothems (Railsback et al. 1994; Shopov et al. 1994; Genty and Quinif 1996). • understanding hydrological conduits (Klimchouk et al. 2000). Paleokarst features describe the long-range effect of descending and ascending waters. • exploring hydrocarbon reservoir rocks and mineral deposits (see Chap. 17).
Travertine
Basics: Paleokarst Bosak, P., Ford, D.C., Horacek, I., Glazek, J. (eds., 1989): Paleokarst – a systematic and regional review. – Developments in Earth Surface Processes, 1, 726 pp., (Artia) Choquette, P.W., James, N.P. (eds., 1988): Paleokarst. – 416 pp., New York (Springer) Choquette, P.W., James, N.P. (1988): Introduction. – In: Choquette, P.W., James, N.P. (eds.): Paleokarst. – 1-21, New York (Springer) Ford, D.C. (1988): Characteristics of dissolutional cave systems in carbonate rocks. – In: Choquette, P.W., James, N.P. (eds.): Paleokarst. – 25-57, New York (Springer) Ford, D.C. (1995): Paleokarst as a target for modern karstification. – Carbonates and Evaporites, 10, 138-147 Ford, D.C., Williams, P.W. (1989): Karst geomorphology and hydrology. – 601 pp., London (Unwin) Playford, P. (2002): Palaeokarst, pseudokarst, and sequence stratigraphy in Devonian reef complexes of the Canning Basin, Western Australia. – In: Keep, M., Moss, S.J. (eds.): The sedimentary basins of western Australia, 3. – Proceedings of the Petroleum Exploration Society of Australia Symposium, Perth, 763-793 Reinhold, C. (1998): Ancient helictites and the formation of vadose crystal silt in Upper Jurassic carbonates (southern Germany). – Journal of Sedimentary Research, 68, 378-390 Schönlaub, H.P., Klein, P., Magaritz, M., Rantitsch, G., Scharbert, S. (1991): Lower Carboniferous paleokarst in the Carnic Alps (Austria, Italy). – Facies, 25, 91-118 Smart, P.L., Whitaker, F.F. (1991): Karst processes, hydrology, and porosity evolution. – In: Wright, V.P. (ed.): Palaeokarsts, and palaeokarstic reservoirs. – Postgraduate Research Institute of Sedimentology, Occasional Publication Series, 2, 1-55 Wright, V.P. (1991): Palaeokarst. Types, recognition, controls and associations. – In: Wright, V.P. (ed.): Palaeokarsts, and palaeokarstic reservoirs. – Postgraduate Research Institute of Sedimentology, Occasional Publication Series, 2, 56-88 Further reading: K031. An ‘Atlas of Speleothem Microfabrics’ by L. Bruce Railsback, including a glossary of relevant terms and a bibliography, can be found in the Internet (www.gly.uga.edu/railsback/speleoatlas).
15.3 Travertine, Calcareous Tufa and Calcareous Sinter Travertine and calcareous tufa are freshwater carbonates. The criteria and origin of travertines (layered carbonates deposited from mineralized waters at thermal springs) and tufas (porous carbonates deposited from cool running waters at springs, in creeks, in zones of ascending groundwater) were discussed in Sect. 2.4.1.6. Calcareous sinter is characterized by well developed lamination and a lack of porosity. These freshwater carbonates were already known in the Precambrian, but the number of Phanerozoic records is low, except for the Quaternary and Tertiary. Techniques used to study travertines and tufa include
Travertine, Tufa, Sinter
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(a) petrographical investigation of thin sections and polished slabs combined with staining and XRD analysis to determine mineralogical composition, (b) rare element analysis (e.g. Sr in order to check diagenetic alterations), and (c) stable isotope analysis, to distinguish between biologically and inorganically controlled precipitation. Diagnostic criteria The following overview of diagnostic criteria is based mainly on Schweigert (1996), who studied and illustrated the microfacies, and the sedimentological and paleontological criteria of Tertiary limestones formed in terrestrial and various freshwater settings. Many calcareous tufa and travertine deposits as well as lacustrine carbonates are modified by pedogenic processes. Therefore, microfacies types of non-marine limestones can vary within wide boundaries. Travertine Ancient travertine deposits are characterized by massive or coarse-bedded, layered carbonates. Biota are restricted to cyanobacteria and bacteria. Fossils enclosed within the travertine (e.g. terrestrial snails, vertebrates, plants) are allochthonous. Common features are laminar crusts consisting of large crystals, banding expressed by changes in color and texture, cyclic bedding caused by seasonal climatic fluctuations, tepee structures, pores filled with vadose cements, desiccation cracks, and brecciation. Microfacies types are controlled by distinct microbial associations, and include: • Laminar cementstones (calcareous sinter) consisting of coarse feather-like arrangements of calcite or aragonite crystals (ray crystal layers; Chafetz and Folk 1984). • Laminar microbial bindstones with a planar upper surface, consisting of very thin micritic and sparry, often discontinuous and broken laminae. • Micritic layers with abundant bushy structures consisting of branched filaments, interpreted having a cyanobacterial origin (Dichothrix morphotype; Guo and Riding 1994). They are often associated with laminar bindstone crusts (Pl. 2/1). • Micrite layers, sometimes with peloidal and clotted fabrics. • Microcrystalline and fine-sparitic layers composed of fibrous cement crystals enclosing large and small open or closed voids (‘sickle cell’ limestones; Fig. 15.6). • Layers with globular or vertically elongated open or closed voids interpreted as gas bubbles (Pl. 2/2). • Detrital layers intercalated between the other microfacies types, formed by reworking of microbial bindstones forming an intraclastic texture (Pl. 2/5).
Fig. 15.6. ‘Travertine’ fabric. ‘Sickle-cell’ limestone exhibiting microcrystalline sheets of inclusion-rich, fibrous cements spanning over lenticular voids. Highly porous limestone characterized by irregular, wavy laminae enclosing lenticular to sickle-shaped voids. These structures are only known from spring mounds in highly alkaline lakes. Miocene: Wallerstein castle, Ries, Germany. After Pache et al. (2001).
• Pisolitic layers composed of densely packed pisoids of different sizes. Inter-pisoid pores are partly or completely filled with aragonite cement (Fig. 15.7). Pisoid accumulations may be intercalated with micrite layers. Calcareous tufa Freshwater tufas are characterized by highly porous fabrics because of the great number of enclosed plant remains. Depending on the plants, a variety of facies types may develop, including leaf tufas, reed tufas, moss tufas (characterized by peloidal textures and sparry fabrics), characean tufas, Vaucheria tufas (tufa formed by xanthophycean green algae), and Cladophorites tufas. Microbial activity is very important for the precipitation of the carbonates but of minor importance for the development of the pores. Plant remains have sheetlike or tubular shapes and are commonly coated by micritic calcite crystals. Algal filaments vary in size from 50 to 100 µm. Different parts of mosses range in size from a few microns to several millimeters. Euhedral calcite crystals growing on the surface of mosses have similar size ranges (Irion and Müller 1968). Calcareous sinter Sinters differ from typical travertines and tufas in that they are predominantly inorganic in origin. They are controlled by high water temperatures in thermal springs, high ion content in mineral springs, high running rates of water, and by the absence or rareness of light leading to microbial settling in caves. This type of sinter occurs in different settings (e.g. thermal travertines, carbonate geyserites, and speleothems).
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Karlsbad Pisoids
Fig. 15.7. Pisolitic travertine. One of the most spectacular pisolitic travertines is the ‘Karlsbad Sprudelstein’, a pea-stone composed of densely packed, well-developed spherical and egg-shaped pisoids and composite pisoid grain aggregates. The pisoids consist of extremely fine concentric laminae arranged in layers that differ in thickness and color. The laminae consist of tangentially arranged aragonite crystals with submicroscopic micropores in between. A conspicuous feature of most pisoids is color banding caused by different iron compounds and organic impurities. The nuclei of the pisoids are very small rock and mineral fragments. The pisoids are poorly sorted. Some samples may be longer than 15 mm in size, but most lie between > 0.5 mm and < 2 mm. The pisoids occur in alluvial deposits and originated from hot, mineralrich springs situated along tectonic faults in a basement of crystalline and volcanic rocks. A: Pisoids differing in size and shape. Note the large composite grain near the bottom. B: Concentric pisoids connected by circumgranular aragonite meniscus cement. C: The cement fringes consist of densely spaced acicular aragonite needles. Fringes are almost uniformly thick (200 to 250 µm). Pores are still open. Karlovy Vary (Karlsbad), a famous spa in Bohemia, Czech Republic. Scale is 1 mm.
Characteristic criteria are layers consisting of calcite or aragonite crystals growing vertically to the substrate. The fabric is characterized by long-prismatic crystals or corresponds to bundles of fibrous crystals (Pl. 2/4). Crystal layers are separated by thin partings of microcrystalline carbonate. Basics: Travertine, calcareous tufa and calcareous sinter Arp, G., Wedemeyer, N., Reitner, J. (2001): Fluvial tufa formation in a hard-water creek (Deinschwanger Bach, Franconian Alb, Germany). – Facies, 44, 1-22 Ferreri, V. (1985): Criteri di analisi di facies e classificazioni dei travertini pleistocenici dell’Italia meridionale. – Rendiconti, Accademia delle Scienze Napoli, Dis. e Mat., 52, 1-47 Irion, G., Müller, G. (1968): Mineralogy, petrology and chemical composition of some calcareous tufa from the Schwäbische Alb, Germany. – In: Müller, G., Friedman, G.M. (eds.): Recent developments in carbonate sedimentology in central Europe. – 157-171, Berlin (Springer) Koban, C.G., Schweigert, G. (1993): Microbial origin of travertine fabrics - two examples from Southern Germany (Pleistocene Stuttgart travertines and Miocene Riedöschingen travertine). – Facies, 29, 251-263 Perry, C.T. (1994): Freshwater tufa stromatolites in the basal Purbeck Formation (Upper Jurassic), Isle-of-Portland, Dorset. – Geological Journal, 29, 119-135 Platt, N.H. (1992): Freshwater carbonates from the Lower Freshwater Molasse (Oligocene, western Switzerland): sedimentology and stable isotopes. – Sedimentary Geology, 78, 81-99 Schweigert, G. (1996): Vergleichende Faziesanalyse, Paläoökologie und paläogeographisches Umfeld tertiärer Süßwasserkarbonate auf der westlichen Schwäbischen Alb und im Hegau (Baden-Württemberg). – Profil, 9, 1-100 Walter, M.B., Desmarais, D., Farmer, J.D., Hinman, N.W. (1996): Lithofacies and biofacies of mid-Paleozoic thermal spring deposits in the Drummond Basin, Queensland, Australia. – Palaios, 11, 497-518 Further reading: K029
Lacustrine Carbonates
15.4 Lacustrine and Palustrine Carbonates Lacustrine carbonates (Sect. 2.4.1.7) originate in freshwater lakes and salt lakes. Palustrine limestones (Sect. 2.4.1.2) are formed through pedogenic modifications of lacustrine freshwater carbonates. The history of lakes is studied by means of sedimentological, geochemical and biological methods. Mineralogical data of carbonate and non-carbonate lake sediments, and trace elements and stable isotopes of lacustrine limestones allow seasonal productivity, temperature changes and lake paleohydrology to be reconstructed (see reviews by Kelts and Talbot 1990; Talbot and Kelts 1990). The isotopic composition is used to distinguish between open and closed lakes (Talbot 1990; Anadón and Utrilla 1993) and between lacustrine and palustrine carbonates (Platt 1992). The intensity and type of bacterial degradation of organic matter is reflected by the δ13C signal of early diagenetic cements. The δ18O value of lacustrine carbonates reflects the composition of the bottom waters from which they were formed (Oberhänsli and Allen 1987; Bellanca et al. 1992; Platt 1992). Biological data relevant for facies analysis of ancient lacustrine carbonates are the composition and zonation of the biota and the biotic controls on carbonate production. Epilithic and endolithic organisms con-
Fig. 15.8. Lacustrine microfacies types. The pictures show microfacies types of Early Permian lacustrine limestones of the Saar-Nahe region in southwestern Germany. The limestones comprise only a minor part of the predominantly siliciclastic sedimentary sequence and occur as thin units called ‘Flöz’ because of their association with coal beds. Carbonate beds may have a thickness of only a few tens of centimeters. Nevertheless, the microfacies types reflect distinct differences in depositional settings (see Fig. 15.10). A: Laminated mudstone (LMF 2) characterized by dark and light layers that can be better recognized at lower magnification. The laminae consist of carbonate and quartz grains. Shallow eutrophic lake. Odenbach Flöz, Lauterecken beds. B: Bivalve floatstone (LMF 9). Leached and recrystallized shells of anthaconaiid bivalves were accumulated by bottom currents. These shells bear some similarity to trilobite fragments and phylloid algae. The former are differentiated by calcium phosphatic shells (see Pl. 94), the latter by thicker plates with relicts of cortical pores (Pl. 58). Odenbach Flöz. C: Micro-oncoid packstone (LMF 10B). Note the distinct differences in size, shape and core/coating ratios. Wiesweil bed. D: Intraclast wackestone (LMF 4) consisting of reworked mudstones (irregularly shaped white grains). Arrows point to corroded fish scales. Odenbach Flöz. After Clausing (1990).
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Microfacies of Lacustrine Carbonates
Fig. 15.9. Lacustrine microfacies types from the Early Devonian Old Red Sandstone (Woodbay Formation) in central North Spitsbergen. A: Strongly bioturbated mudstone-wackestone (LMF 1). Sand-sized quartz grains and various skeletal grains (ostracods, charophycean oogonia, fish fragments) are embedded within a micritic matrix. B: Bioclastic ostracod packstone (LMF 6). Densely packed, thin ostracod valves forming a coquina. The accumulation is interpreted as population spread under favorable conditions. C: Bioclastic gastropod grainstone (LMF 8). The gastropods are associated with thick-shelled ostracods. Littoral lake environment. D: Lithoclastic rudstone (LMF 13). Most grains are very well rounded. The granule- to pebble-sized mudstone clasts were accumulated by storms. Scale is 1 cm. After Blomeier et al. (2003).
tribute to carbonate formation and destruction. Carbonate formation in lakes by algae and cyanobacteria is triggered by the trapping and binding of particles, cementation by and encrustation with CaCO3, and secretion of CaCO3 within algal mucus. Cyanobacteria and algae assist in the growth of oncoids, stromatolites and reefs. In addition, accumulations of benthic and planktonic algae as well as invertebrate shells contribute to the formation of lacustrine carbonates. Destruction of carbonates occurs by means of endolithic algae and bacteria loosening carbonate surfaces and producing finegrained sediment.
fossils. In addition, accumulations of lithoclasts and intraclasts may be common. Fossils include cyanobacteria, algae, gastropods and bivalves, ostracods, fish bones, scales and teeth, and amphibian remains. Algae are represented by green algae (Cladophorales) and occasionally also by calcareous algae known predominantly from marine environments (Pl. 131/4). Saltwater lake carbonates are poor in fossils but may yield ostracods and some remains of crabs. Other common criteria include stromatolites, oncoids and ooids (Pl. 13/1). Common fabrics of lacustrine carbonates are lamination and varves, nodular structures and bioturbation.
15.4.1 Microfacies of Lacustrine Carbonates
Stromatolites Cyanobacteria contributing to the formation of stromatolites are strongly salinity-controlled. Stromatolites originate in marine and marginal-marine environments of normal and hypersaline salinities as well as in nonmarine environments including freshwater and brackish lakes, and salt and soda lakes (Riding 2000).
15.4.1.1 Microfacies Criteria Common microfacies criteria of freshwater lake carbonates are micrite, peloids, radial ooids (see Boxes 4.15 and 4.16), oncoids and stromatolites, and various
Lacustrine Microfacies Types
Lacustrine stromatolites are already known from Archean and Proterozoic sediments. Common features of Phanerozoic stromatolites are a significant morphological variety (undulations, columns, domes, arborescent growth forms), light and dark macrolaminae composed of many dark and light micritic and microsparitic laminae, and radial fibrous textures reflecting growth patterns of cyanobacterial and algal threads. A survey of ancient lacustrine stromatolites was given by Stapf (1989). Oncoids Freshwater oncoids have been discussed in Sect. 4.2.4.1. The subspherical or concave-convex oncoids are scattered or densely packed within a micrite matrix or in marls. Sizes vary between less than 1 mm and several centimeters (see Fig. 15.8C; Pl. 131/5). Very
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large oncoids (up to 20 cm) occur at the margins of fluvial channels (Anadon and Zamarreno 1981). Many oncoids are irregularly shaped. The cores may be very small and formed by sedimentary grains or rock fragments, or smaller oncoids, or they may consist of gastropods and bivalve shells. The cores are surrounded by often densely spaced, irregularly concentric, up to 1 mm thick macrolaminae. These macrolaminae are composed of dark micritic and bright microsparitic microlaminae. The laminae are arranged in parallel, often crinkled, dome-shaped or columnar layers. Spongy and thrombolitic fabrics are common. Radial, fan-shaped microstructures are caused by the growth patterns of cyanobacteria (see Monty and Hardy 1976 for illustrations). Microstructures of lacustrine oncoids are often better preserved than marine oncoids.
Box 15.1. Common lacustrine microfacies types (LMF). LMF 1: Lime mudstone consisting of very fine-grained micrite crystals. Can originate from cyanobacterial and algal blooms or by abrasion of limestones. Sometimes strongly bioturbated (Fig. 15.9A). Occurs in deeper and protected lake parts, but the sediment may be reworked by storms and redistributed also in shallow areas. LMF 2: Inhomogeneous laminated lime mudstone with very rare or no fossils (Fig. 15.8A). The laminated fabric is caused by the alternation of light and dark layers, the latter often rich in organic matter, plant debris and/or cyanobacteria forming very thin laminae that can be recognized by means of fluorescence microscopy. Clay and fine, silt-sized terrigenous quartz may be common and contribute to the lamination. LMF 2 characterizes a relatively deep quiet-water basin with stratified water column and seasonal variations. Poor oxygenation is indicated by the lack of benthic organisms and burrowing. Fossils may be represented by rare fish remains, ostracods and conchostrakes. LMF 3: Lime mudstone and wackestone with leaves of plants accumulated on bedding planes. The leaves may be encrusted by carbonate. This microfacies is common in Late Paleozoic and Cenozoic lacustrine sediments (e.g. Potomogeton layers, Miocene). LMF 3 characterizes a quiet-water far-shore or near-shore environment. LMF 4: Inhomogeneous fine-bedded intraclast wackestone (Fig. 15.8D) characterized by reworked laminated mudstone; LMF 2) and nektonic fossils. Frequent in nearshore and shore lake-floor areas affected by waves. LMF 5: Densely packed nodular, peloidal wackestone or floatstone with reworked lacustrine grains and with black pebbles of lacustrine limestone, sometimes also with extraclasts. Common near the shoreline of large lakes. Formed by storm events. LMF 6: Ostracod mudstone, wackestone, packstone or grainstone characterized by double-valved or disarticulated thin or thick ostracod shells, often monospecific (Pl. 130/ 2; Fig. 15.9B). This type occurs in different near-shore, far-shore and basinal parts of lakes. LMF 7: Charophycean mudstone or wackestone (Pl.
130/4; Pl. 65/1, 3) with abundant oogonia and stem remains. Often associated with ostracods or ostracod debris. This microfacies characterizes the shore region of many humid and semiarid lakes and brackish environments of coastal regions. LMF 8: Gastropod packstone and grainstone (Pl. 130/ 3; Fig. 15.9C). Often composed of monospecific communities. The gastropods may be associated with bivalves and ostracods. LMF 9: Bivalve floatstone (Fig. 15.8B) composed of centimeter-sized, often parallel arranged shells. LMF 9 occurs on near-shore lake floors and is rarely affected by waves. Coquinas may represent storm deposits. LMF 10: Oncoid floatstone, packstone and wackestone. Two subtypes: Large oncoids (commonly porostromate or agglutinated oncoids; Pl. 131/5; LMF 10A) occur either associated with stromatolites or separately. Microoncoids with submillimeter sizes (Fig. 15.8C; LMF 10B) accumulate in packstones. These oncoids have nuclei of quartz or silt grains, surrounded by only a few micritic laminae. The large oncoids exhibit various microfabrics reflecting different growth forms of different cyanobacteria and algae. The micro-oncoid packstone facies occurs in near-shore and shore areas. Floatstones with centimeter-sized oncoids are common in higher-energy fluviallacustrine and brackish-water settings. LMF 11: Stromatolite bindstone (Pl. 131/1, 3). Very variable in geometry, shape and internal fabrics (see above and Sect. 9.1 for more information). Contribute to the formation of ‘algal reefs’. LMF 12: Green algal bafflestone and wackestone (Pl. 130/1). Tubular green algae (e.g. Cladophorites) living in near-shore and shore-settings contribute to the accumulation of sediment and the formation of lacustrine ‘algal reefs’. LMF 13: Lithobioclast rudstone (Fig. 15.9D) composed of reworked, poorly sorted and sometimes strongly rounded carbonate lithoclasts and worn fossil fragments. Storm deposits.
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The shape of freshwater oncoids seems to be controlled to a lesser degree by environmental factors (e.g. water energy, terrigenous input) than texture and microfabrics (Abell et al. 1982). The best way to gain a high paleoenvironmental resolution is the taxonomic differentiation of cyanobacterial and algal associations (see Pl. 8/4). Cyanobacteria communities in lacustrine oncoids and stromatolites have remained stable at least since the Cretaceous (Monty and Mas 1981).
15.4.1.2 Microfacies Types of Lacustrine Limestones Thirteen common and widely distributed microfacies types can be recognized when the facies of lacustrine limestones of different age are compared (see Box 15.2). These LMF Types are listed in Box 15.1 and shown in Pl. 130, Pl. 131 and in Figs. 15.8 and 15.9.
Lacustrine Limestones
15.4.1.3 Distribution of Lacustrine Microfacies Types Fig. 15.10 is a rough sketch showing the occurrence of more frequent microfacies types of freshwater carbonates. More facies types exist, of course, depending on the climatic setting, the size and depth of the lake basin, and the topography of the surrounding regions. A strict actuogeological approach to the environmental interpretation of ancient lake sediments is difficult, in that most present-day lakes are shaped by glaciers and therefore provide poor analogs for ancient lacustrine sedimentation. This was predominantly controlled by tectonic movements, and only few modern lakes have an areal extent and a sediment thickness comparable with those of ancient counterparts, particularly in the Paleozoic.
Plate 130 Lacustrine Limestones: Common Microfacies Types The plate displays Tertiary and Cretaceous lake carbonates. The Tertiary samples (–> 1-3) are taken from the Miocene Ries Lake, which formed within a large impact crater. Bioherms were built by green algae and cyanobacteria, and multiple lacustrine-vadose sinter veneers at the margin of the humid, hydrologically closed soda lake basin (Arp 1995). The lacustrine sediments are overprinted by pedogenic and early diagenetic processes modifying the primary depositional fabric and implying an adaption of traditional limestone classifications (see Sect. 8.4.2; Pl. 48). Algal limestone, ostracod grainstones, wackestones and packstones, as well as gastropod grainstones and wackestones are common microfacies types in lacustrine sediments. Unlike marine limestones, gastropods and ostracods are represented by very low-diversity associations, often by just one or two species. Other common type of lacustrine limestones formed in oligotrophic or brackish lakes are wackestones and mudstones with charophycean algae (–> 4). Lacustrine microfacies types (LMF) are defined in Box 15.1. 1 Green algal (Cladophorites) bafflestone: LMF 12. The basic elements of the algae are erect and branched tubes, arranged in cushion- or tuft-like thalli that sometimes coalescence forming meter-sized structures. Rhythmic growth patterns indicate seasonal controls. The tube walls consist of micrite. Cladophorites (C) is regarded as a green alga identical with or closely related to the modern genus Cladophora. Note associated gastropods (G) and charophycean oogonia (CH). Tertiary (Miocene): Hainsfarth, Ries Basin, southern Germany. 2 Ostracod grainstone: LMF 6. Some valves are stacked into each other (white arrows). The ostracods comprise 40–50% of the very porous rock. The shell interior is filled with sinter cement and meteoric-phreatic sparite. The shells are covered by a clear dolomite cement (black arrows). Calcite appears black due to staining with Alizarin-S. Paleoenvironment: Wave-exposed eulittoral zone, sands near the algal bioherm. Note that comparable microfacies types occur in shallowmarine hypersaline environments as well. Same locality as –> 1. 3 Gastropod grainstone with some ooids: LMF 8.The gastropod shells (Hydrobia) are leached and preserved as micrite envelopes only. The envelopes are covered by thin layers of microcrystalline cement. Hydrobia was the only aquatic gastropod that flourished in the soda lake. Some calcitic ooids exhibit a radial structure (white arrow), other ooids are dolomitized and appear black (black arrow). Paleoenvironment: Wave-exposed shallow eulittoral zone near the algal bioherms. Same locality as –> 1. 4 Charophycean algal wackestone: LMF 7. Most algal fragments are internode parts (arrows), see Pl. 65, Fig. 10.20, and Sect. 10.2.1.8. Thin shells are ostracod fragments. Paleoenvironment: Brackish-water lagoon. Cretaceous: Spain. –> 2 and 3: Arp (1995)
Plate 130: Lacustrine Limestones
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Palustrine Carbonates
Box 15.2. Selected case studies of ancient lacustrine and palustrine carbonates. Case studies on freshwater oncoids and stromatolites are listed in Box 4.8. References dealing with salt and soda lakes are marked by +, those on palustrine carbonates by *. Tertiary: Alonso-Zarza and Calvo 2000*; Anodon and Utrilla 1993; Anodon and Zamarreno 1981; Anodon et al. 2000; Arenas et al. 1992; Armenteros et al. 1997*; Arp 1995; Bellanca et al. 1992; Bradley 1929; Bustillo et al. 2002; Cabrera and Saez 1987*; Calvo et al. 1989, 1995, 2000; Fouch et al. 1976; Herbig 1988*, 1994*; Leinfelder and Hartkopf-Froder 1990; Krylov 1982; Nickel 1983; Platt 1992+; Reggiani 1989; Riding 1979; Rouchy and Monty 1981+; Schweigert 1996; Stanley and Collinson 1979*; Straccia et al. 1990; Surdam et al. 1975+; Surdam and Wray 1976; Szulc et al. 1991+; Szulc 1992; Trappe 1992+; Wells 1983*; Willems and Wuttke 1987; Wright 1990*. Cretaceous: Alonson et al. 1991; Bell 1989+; Buatois et al. 2000; Carozzi 1989; Fouch 1979; Fregenal-Martinez and Melendez 1994; Freytet 1984+; Freytet and Plaziat 1982*; Martin-Chivelet and Gimenez 1992*; Misik and
15.4.2 Palustrine Carbonates Palustrine limestones assist in recognizing seasonal wetland environments in the geological record (Wright and Platt 1995). These limestones (Sect. 2.4.1.2) exhibit criteria that also occur in pedogenic carbonates (pseudo-microkarst associated with fenestrae, circumgranular cracks and peloidal fabrics; Plaziat and Freytet 1978) and caliche horizons (rhizolites, rhizoconcre-
Sykora 1980; Monty and Mas 1981; Normati and Salomon 1989*; Platt 1989. Jurassic: Bell 1989+; Leinfelder 1985. Triassic: Carozzi 1964; Clemmensen 1978; Gore 1989; Milroy and Wright 2000; Olsen 1986; Seegis and Goerigk 1992*; Tucker 1978; Wright and Mayall 1981. Permian: Clausing 1990, 1992; Clausing et al. 1992; Freytet and Lebreton 1992; Gebhardt 1988; Gebhardt et al. 2000; Guy-Ohlson and Lindström 1994; Schäfer and Stapf 1978; Schneider and Gebhardt 1993; Stapf 1973, 1989; Warren 1983. Carboniferous: Buatois and Mangano 1994; Driese 1985; Gebhardt 1988; Valero-Garcés et al. 1994*. Devonian: Blomeier et al. 2003, Demicco et al. 1987; Donavan 1975; Duncan and Hamilton 1988.
tions, coated grains; Seegis 1993). The main difference, however, is the close association of these features with lacustrine sublittoral limestones. Palustrine carbonates are paleosols reflecting pedogenic effects, such as the repeated drying of the sediment in carbonate swamps bordering lakes, or in episodically drying shallow lakes. Examples are shown in Pl. 48/2, 3. The carbonate beds are commonly associated with claystones, marls, sandstones, cherts and coals. Stable isotopes of the micritic
Plate 131 Lacustrine Carbonates from the Early Permian (Rotliegend) of Germany The plate displays the microfacies of lowermost Permian lacustrine limestones, formed in intracontinental basins filled predomiantly with fluviatile-deltaic and lacustrine siliciclastic sediments. Limestones are rare. Most carbonates are of algal and microbial origin or formed by the accumulation of ostracods and mollusks. Sporadic subaerial exposure is indicated by pedogenic horizons rich in caliche pisoids (Pl. 14/6) at the top of stromatolites. Other facies types are shown in Fig. 15.8. Facies-diagnostic criteria of these carbonates are various types of biogenic laminations (–> 2), stromatolites (–> 1, 3) and oncoids (–> 5) formed in shallow marginal parts of the lakes. 1 Lacustrine columnar stromatolites: LMF 11. These structures occur in limestones and marls with a thickness of up to several meters. The columns are distinctly laminated. Larger columns consist of small-scaled stromatolites (black arrow). Early Permian (Lower Rotliegend): Altenglan, Saar-Nahe Basin, Southwestern Germany. 2 Irregularly laminated mudstone and peloid mudstone. Note the small-scale alternations of syngenetically deformed layers. These limestones yield ‘marine’ algae (–> 4). Early Permian (Lower Rotliegend): Birkigt near Freital, Döhlen Basin, Thuringia, Germany. 3 Stromatolite boundstone: LMF 11. Note the infilling of micro-oncoids (arrows) between the columns, indicating a sometimes turbulent shoreline/nearshore environment. Early Permian (Upper Kusel Group): Saar-Nahe Basin, Germany. 4 ‘Marine’ green and red algae can occur in lacustrine sediments as seen in examples from Thuringia (Schneider and Gebhardt 1992). The fossil was assigned to the dasyclad genus Macroporella. The occurrence of these algae has been explained by wind drift of algal cysts or acinetes from marine environments to relatively saline lakes. Same locality as –> 2. 5 Lacustrine oncoid with thrombolite (T) microstructure: LMF 10. Thrombolites are characterized by clotted structures and the absence of conspicuous lamination. The clots are interpreted as being formed by coccoid and filamentous microbes. Early Permian (Upper Rotliegend): Dannenfels, Saar-Nahe Basin, Germany. –> 1, 5: Stapf 1989; 3: Clausing 1990; 2, 4: Courtesy of J. Schneider (Freiberg)
Plate 131: Early Permian Lacustrine Carbonates
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Lacustrine Carbonates: Microfacies Distribution
Fig. 15.10. Schematic distribution of common microfacies types of lacustrine freshwater limestones. Based on Clausing (1990) and supplemented after Stapf (1989), Arp (1995), Schweigert (1996) and other sources (see Box 15.2). Dark bars indicate main distributions.
matrix and of carbonate cements offer a possibility for recognizing pedogenic overprinting associated with subaerial exposure of ephemeral lake deposits (Platt 1992). Common microfacies types are • mudstone with detrital quartz floating within a micrite carbonate matrix; some quartz grains may be surrounded by calcitic aureoles; • nodular dolomitic mudstone with irregular tube-like millimeter-sized voids, sometimes filled with sediment or quartz grains; common on top of lacustrine limestones; • mudstone with micritic nodules and voids, curved planes surrounding nodules. Basics: Lacustrine and palustrine carbonates Allen, P.A., Collison, J.D. (1986): Lakes. – In: Reading, H.G. (ed.): Sedimentary environments and facies. 2nd edition. – 63-94, Oxford (Blackwell) Anadon, P., Cabrera, L., Kelts, K. (eds., 1991): Lacustrine facies analysis. – International Association of Sedimentologists, Special Publication, 13, 336 pp., Amsterdam (Elsevier)
Arp, G. (1995): Lacustrine bioherms, spring mounds, and marginal carbonates of the Ries-Impact-Crater (Miocene, Southern Germany). – Facies, 33, 35-90 Blomeier, D., Wisshak, M., Dallmann, W., Volohonsky, E., Freiwald, A. (2003): Facies analysis of the Old Red Sandstone of Spitsbergen (Wood Bay Formation): reconstruction of depositional environments and implication of basin development. – Facies, 49, 151-174 Clausing, A. (1990): Mikrofazies lakustriner Karbonathorizonte des Saar-Nahe-Beckens (Unterperm, Rotliegend, SW-Deutschland). – Facies, 23, 121-140 Friedman, G.M., Krumbein, W. (eds., 1985): Hypersaline ecosystems. – Ecological Studies, Analysis and Synthesis, 35, 500 pp. Gebhardt, U., Merkel, T., Szabados, A. (2000): Karbonatsedimentation in siliziklastischen fluviatilen Abfolgen. – Freiberger Forschungshefte, C490, 133-168 Gierlowski-Kordesch, E.H., Kelts, K.R. (eds., 2000): Lake basins through space and time. – American Association of Petroleum Geologists, Studies in Geology, 46, 648 pp. Katz, B. (ed., 1990): Lacustrine basin exploration – case studies and modern analogues. – American Association of Petroleum Geologists, Memoir, 50, 340 pp. Matter, A., Tucker, M. (eds., 1978): Modern and ancient lake sediments. – International Association of Sedimentologists, Special Publication, 2, 290 pp. Monty, Cl., Mas, J.R. (1981): Lower Cretaceous (Wealden)
Peritidal Carbonates
blue-green algal deposits of the Province of Valencia, Eastern Spain. – In: Monty, Cl. (ed.): Phanerozoic stromatolites. Case studies. – 85-120, Berlin (Springer) Picard, M.D., High, L.R. (1972): Criteria for recognizing lacustrine rocks. – In: Rigby, J.K., Hamblin, W.K. (eds.): Recognition of ancient sedimentary environments. – Society of Economic Paleontologists and Mineralogists, Special Publication, 16, 108-145 Stapf, K.R.G. (1989): Biogene fluvio-lakustrine Sedimentation im Rotliegend des permokarbonen Saar-NaheBeckens (SW-Deutschland). – Facies 20, 169-198 Talbot, M.R. (1990): A review of the palaeohydrological interpretation of carbon and oxygen isotope ratios in primary lacustrine carbonates. – Chemical Geology (Isotope Geoscience Section), 80, 261-279 Talbot, M.R., Kelts, K. (eds., 1989): The Phanerozoic record of lacustrine basins and their environmental signal. – Paleogeography, Palaeoclimatology, Palaeoecology, Special Issue, 70, 304 pp. Wright, V.P. (1990): Lacustrine carbonates. – In: Tucker, M., Wright, V.P. (eds.): Carbonate sedimentology. – 164-189, Oxford (Blackwell) Wright, V.P., Platt, N.H. (1995): Seasonal wetland carbonate sequence and dynamic catenas: a re-appraisal of palustrine limestones. – Sedimentary Geology, 99, 65-71 Further reading: K025, K026, K027, K176, K188
15.5 Peritidal Carbonates Peritidal (literally:‘around tides’) carbonates are shallow-subtidal, intertidal and supratidal sediments formed in marginal-marine and shoreline depositional environments. Sediments formed in intertidal-supratidal flats are summarized as tidalites (see Alexander et al. 1998). Major criteria were discussed in Sect. 2.4.2.2 and Sect. 14.1.2.1. These carbonates are easy to recognize in outcrops and in thin sections. They are important paleobathymetric indicators, reflect sea-level fluctuations, and form the basis for the evaluation of sedimentary cycles and stratigraphic sequences (see Sect. 16.1). The facies types are vertically arranged in (regressive) shallowing-upward successions consisting of shallow-marine sediments overlain by intertidal and supratidal carbonates that are subject to periods of subaerial exposure. The subtidal zone is permanently submerged. The intertidal or littoral zone between normal low- and high-tide levels is submerged on a diurnal or semidiurnal basis. Tidal flats are repositories of allochthonous sediment transported by currents and storms from the subtidal carbonate factory. The Holocene tidal and storm deposits along the western shores of Andros Island in the Bahamas have become a diagnostic guideline for a humid tidally influenced carbonate environment (Shinn et al. 1965, 1969). The tidal flats, sabkhas and lagoons of Abu
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Dhabi in the Arabian Gulf serve as recent models for the accumulation of supratidal, intertidal and shallow subtidal sediments in an arid climate (Evans et al. 1969; Schneider 1975). The following section summarizes diagnostic microfacies criteria, and describes a case study from the Devonian.
15.5.1 Criteria of Peritidal Limestones This section lists the criteria of peritidal and shallowsubtidal carbonates and discusses the merits of the criteria in the context of facies analyses. Unlike marine limestones, the recognition and interpretation of ancient marginal-marine limestones is fundamentally dependent on comparison with modern transitional-marine carbonate environments. This approach is very successful, but should be applied cautiously. Fine-laminated carbonates do not necessarily originate in tidal zones, and spar-filled voids in micritic rocks could have been formed in tidal zones as well as in various subtidal environments. The best way to overcome these difficulties is to use the whole set of facies criteria and take into consideration the changes in microfacies and the sedimentary structures within vertical sequences. Ancient carbonate rocks are traditionally studied by dividing peritidal environments into subtidal, intertidal and supratidal zones, despite the problems involved. Whether an area is subtidal, intertidal or supratidal depends on the tidal range, mean sea level and its topographic position relative to sea level. To avoid these difficulties, the exposure index (percentage of the time that a particular area is exposed) proposed by Ginsburg et al. (1977) should be used (see Wright 1990 for discussion).
15.5.1.1 Definitions Tidal flat carbonates are predominantly micrites, often peloidal micrites. Common limestone types are mudstones, wackestones and bindstones. Intercalations of grainstones may be present, e.g. in tidal channels, or as storm deposits. Supratidal This is the most diagnostic zone in humid and arid tidal flats and represents the splash zone above high tide; it is flooded only a few times each month by spring tides or storms. Most of the sedimentation occurs above
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normal high tide during storm flooding. Diagnostic criteria for supratidal environments are • evidence of subaerial exposure and pedogenic influence, • evidence of cementation in the vadose zone. Intertidal This zone is alternately flooded and exposed. Intermittent exposure is indicated by desiccation features, muddy surfaces with animal tracks and trails, and pores filled with vadose cement. Characteristic features of the intertidal regime are alternating erosion and deposition and rapid changes in current and wave velocity. These processes result in discontinuous sedimentation associated with scour-and-fill and the formation of channels, accumulation of reworked sediment (e.g. intraformational conglomerates), and substantial changes in grain size from lamina to lamina and bed to bed. Subtidal This zone below the intertidal zone is permanently submerged. Water depth is a few meters. The subtidal zone is conventionally subdivided into a shallow and a deeper subtidal zone. The shallow subtidal zone is characterized by calcareous algae and invertebrates adapted to the phytal (e.g. foraminifera, bryozoans, worms; see Voigt 1956). Epiphytes are common. Important indicators are the occurrence of dasyclad algae and encrusting organisms. The lower boundary, after comparison with recent zonations, can be drawn at approximately 30 m. The deeper subtidal zone is characterized by high-diversity benthic organisms. Calcareous algae still occur (red algae). Dasyclads, however, are absent.
15.5.1.2 Major Facies Criteria of Peritidal Carbonates Bedding Lamination (Pl. 18/2; Pl. 123/1): A very common feature of tidal carbonates is fine-scale lamination, usually on a millimeter scale (Sect. 5.1.3). Lamination often occurs associated with fenestral pores. The sediment may be trapped and bound by microbial mats, giving rise to the alternation of fine-grained sediment layers and irregular crinkled layers. Many ancient peritidal limestones exhibit a lamination consisting of millimeter micrite or dolomicrite layers with thin organic or bituminous seams, and thicker, up to several centimeters, layers composed of grainstone or packstone with intraclasts and peloids. The latter represent storm deposits.
Peritidal Carbonates: Criteria
Lamination should be studied with regard to the (1) lithologic composition (only limestone or dolomite, or limestone-dolomite couplets), (2) geometry of the laminae (undulation; parallel laminae; cross-bedded laminae), (3) extension of laminae (continuous, disrupted), (4) microfacies of the laminae (micrite, pelmicrite, pelsparite), (5) type and frequency of associated fenestral pores. See Read (1975) for a classification of lamination types. –>Note: Parallel lamination occurs in subtidal, intertidal and supratidal environments, but is especially abundant in the intertidal zone. Stratification: Bedforms and stratifications made by waves and currents (ripples, planar and cross stratification, winnowed sand lenses; flat and cross-bedded calcarenites) exhibit a wide variety of types. Further information can be found in Demicco and Hardie (1994). –> Note: Common in the shallow subtidal and lower intertidal zone. Desiccation features Desiccation and shrinkage cracks seen on bedding surfaces are considered as one of the most useful sedimentary structures for identifying subaerial exposure and ancient carbonate tidal flats. Desiccation cracks formed by drying out of subaerially exposed mud have different forms depending on the rate of drying, presence or absence of protecting microbial mats, exposure time and bed thickness, and comprise three types: • Mud cracks (Pl. 25/5) and mud polygons. • Prism cracks: Polygonal network of cracks, essential normal to bedding, that breaks up the sediment into vertical prisms. • Sheet cracks: Planar to gently undulatory cracks, commonly parallel to the bedding and filled with sparry calcite or micrite. –>Note: Desiccation mud polygons are common in the supratidal and upper intertidal zone. Because mud cracks can also form by substratal and subaqueous processes, much care must be taken in interpreting mud crack structures in ancient carbonates (see Demicco and Hardie 1994). Tepees A special form of polygonal structures are tepees (similar to inverted depressed Vs resembling Red Indian tents) characterized by antiformal structures with sharp apices. Tepees originate from desiccation, but the formation of large peritidal tepees is also controlled by fluctuating groundwater conditions. They are associated with fractured and bedded tidal carbonates. The fractures are filled with marine-vadose travertines and/ or terra rossa.
Peritidal Carbonates: Criteria
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Other processes creating tepees are the crystallization of minerals (e.g. evaporites) and thermal contraction and expansion. See Assereto and Kendall (1977) and Kendall and Warren (1987) for details. –>Note: Tepees are common features of supratidal environments. Peritidal tepees (Assereto and Kendall 1971) are strandline indicators.
distribution of the laminoid fenestrae is closely related to that of microbial mats. Note that the often used term loferite (Fischer 1964) is not fully identical with birdseye limestone. Loferite refers to bedded or massive micritic limestones containing millimeter-sized pores that correspond to birdseyes as well as shrinkage pores, sheet cracks and tubular fenestrae.
Intraclasts and breccias Intraclasts (Pl. 16/7): Millimeter and centimetersized, and angular to rounded clasts of lime mudstones, packstone, bindstones and grainstones are common constituents of tidal carbonates. –> Note: Rip-up intraclasts are abundant in the intertidal zone and the lower part of supratidal zones. Storms affecting supratidal flats and coast lines produce intraclast rudstones and floatstones containing reworked indurated carbonate crusts and black pebbles. Flat-pebble and edgewise conglomerates: Many tidal and shallow subtidal carbonates contain beds composed of flat-lying and imbricated carbonate rounded clasts exhibiting different microfacies types. The clasts are produced by storms eroding the shallow sea bottom (see Sect. 5.3.3.3). –> Note: Flat-pebble conglomerates occur in supratidal, intertidal and subtidal settings but are more common in the intertidal zone. Black pebbles (Pl. 16/3): Millimeter- to centimetersized black lithoclasts are a valuable tool for identifying partial or complete subaerial exposure (see Sect. 4.2.8.1). Breccias: Solution-evaporite-collapse breccias (Sect. 5.3.3.3) may be common in supratidal limestones formed under arid or semiarid conditions.
Microbial/algal mats Algal-laminated (‘cryptalgal’) structures characterized by crinkled or flat micritic or pelmicritic laminae (Pl. 18/4; Pl. 123/2), and relicts of cyanobacterial or algal filaments are common criteria of ancient tidal carbonates (see Sect. 9.1.3.1). These bindstones and peloidal packstones are associated with fenestral limestones and lime mudstones. A characteristic feature is a frequent and abrupt interruption of continuous laminations and mats by lenses, channels or beds of coarsegrained sediment (intraformational conglomerate, ooid sand, skeletal sand). Because mats are often formed on more frequently exposed parts of tidal flats, the association of mats, fenestral pores, mud cracks and evaporites is a common feature. Microbial and algal mats may indicate destruction. Criteria for the existence of former biofilms and microbial mats are • local doming of sedimentary surfaces; • fenestral pores; • clasts derived from microbial mats; • textures indicating baffling or trapping effects (selective accumulations of fine sediment between coarser sediment; selective accumulations of coarser grains; orientation of grains parallel to the surface of microbial mats); • enrichment of terrigenous clay minerals in layers; • enrichment of Fe, S, P, Corg and various other elements in layers. Element mapping offers a good method for recognizing former microbial mats (Kropp et al. 1996). –> Note: Microbial mats occur in intertidal and supratidal settings, but may also form in subtidal environments if high salinities or other factors prevent grazing organisms from destroying the mats. Grazing organisms, particularly cerithid gastropods, are common constituents of well-laminated sediments consisting of smooth mats with laminoid fenestrae which are distinct features of the lower intertidal zone. The middle and the upper tidal zones are characterized by poorly laminated sediments with an irregular fenestral fabric lying below pustular mats. Records of microbial mats are abundant in the carbonates formed in the upper intertidal and the supratidal zone, probably because grazing is limited by prolonged exposure.
Fenestral structures Fenestral cavity structures (birdseyes; Sect. 5.1.5; Pl. 20/7; Fig. 14.25) are voids in peloidal and micritic mudstones and wackestones, and in packstones and grainstones. The voids are larger than the intergranular pores and partially or completely filled with sediment or cement. According to Groover and Read (1978), three types can be distinguished: • Irregular fenestrae (Pl. 50/5; Pl. 123/3), typical birdseyes: equidimensional to irregular in shape, several millimeters across. • Laminoid fenestrae developed in laminated sediments (Pl. 41/2), particularly microbial carbonates. • Tubular fenestrae, formed mainly by burrowing organisms or plant rootlets (Pl. 20/6). –> Note: Abundant in intertidal and supratidal environments. Irregular and laminoid fenestrae are characteristic of the upper intertidal and supratidal zones. The
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Stromatolites Spectacular stromatolites (Sect. 9.1.4; Fig. 9.2) consisting of laterally linked hemispheroidal growth forms and stacked hemispheroids are known from recent and ancient intertidal settings, but the existence of stromatolites does not necessarily indicate intertidal paleoenvironments. Many ancient stromatolites, particularly those with vertically pronounced reliefs, were formed in subtidal settings and are also known in modern subtidal environments.
Peritidal Carbonates: Criteria
versity is generally high. Low faunal diversity may indicate stressed environmental conditions. Most common invertebrates and many calcareous algae live in the illuminated, shallow upper part of the subtidal zone.
Biogenic structures Root structures: Submillimeter- and millimetersized, spar-filled branched tubular structures, different from fenestral structures, may represent relicts of plant roots (Pl. 20/6). –> Note: Root molds of land plants are common in supratidal environments, but morphologically similar molds are produced by subtidal aquatic plants. Structures interpreted as molds of plant roots, are no definitive characteristics for subaerial or subaquatic settings. Burrows (Pl. 19) are common in all marine settings, but show differences with regard to shape, size and distribution patterns in different parts of shelf environments (Sect. 5.1.4). –> Note: Variously arranged burrows are rare in supratidal environments. The lower intertidal zone is commonly heavily burrowed and often characterized by millimeter-sized vertical burrows. Note that strong burrowing occurs also in shallow and deeper parts of the subtidal zone.
Sea grass Today, sea grass meadows are located in the intertidal and the shallow subtidal zone. Since the vegetation determines the nature and distribution of carbonate sediment types, each zone has its own distinctive skeletal communities (Cann and Gostin 1985) and is reflected by calcareous epiphytes and organisms living in protected niches below and between the aquatic plants. • The subtidal Posidonia meadows occurring in shoreline-attached platforms or offshore banks can be recognized by foraminifera, coralline algae, bivalves and gastropods living on the blades. The corresponding sediment is a poorly sorted skeletal packstone. • Intertidal flats may be covered by meadows of the sea grass Zostera, that grows in loose communities in lower intertidal settings and in dense meadows in the upper intertidal zone. The sediment is a poorly sorted mixture of skeletal sand with molluscan infauna, and terrigenous mud and sand. These sand flats merge landward into the • upper intertidal mangrove zone, characterized by laminar mudstones, pseudobreccias and ‘root rocks’ (Galli 1991). The latter exhibits a dense, irregular network of circular voids and vertical cavities exhibiting downward bifurcations.
Biota The biotic diversity and the composition of benthic assemblages offer the best criteria for distinguishing and subdividing tidal and subtidal carbonates. • Supratidal environments are characterized by very low diversity, rare well-preserved fossils and the dominance of only a few organism groups (cyanobacteria, ostracods, some foraminifera, a few mollusks). Plant litter may be common. • Intertidal environments are characterized by low diversity organisms that adapted to rapidly changing stress conditions. More frequent organisms are (cerithid) gastropods, ostracods, benthic foraminifera and bivalves. Plant litter is common. Cyanobacteria are abundant (coastal and interior algal/microbial marshes in humid climate). Calcareous algae are represented by dasyclads. Storm layers may contain transported subtidal organisms. • Subtidal environments are characterized by normal marine biota including skeletal invertebrates and shell debris as well as trace fossils, especially burrows. Di-
Diagenetic features Cementation: Some early diagenetic criteria occur preferentially in tidal limestones (see Sect. 7.4.2 and Sect. 7.4.3). These criteria include: • Gravitational (microstalaktitic) cements, e.g. dripstone cements in intertidal birdseyes (Pl. 33/4). • Meniscus cement between grains of supratidal beachrock (Pl. 33/4). • Meteoric-vadose pisoids (Pl. 14/1; Pl. 126/1, 2) . • Calcite or aragonite sinter crusts formed on intertidal and supratidal surfaces. • Irregular carbonate encrustations on substrates affected by surf sprays; associated with pisoids (coniatolites; Purser and Loreau 1973; Pl. 14/2). Dolomitization: Intertidal and supratidal carbonates may be characterized by very early diagenetic fine-crystalline dolomicrite formed under conditions described by the sabkha evaporation model and the mixing zone model (see Sect. 7.8.2). Early diagenetic origin is indicated by dolomite chips occurring in tidal breccias. Often associated with fenestral voids (Pl. 39/2).
Criteria of Tidal Carbonates
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Table 15.1. General summary of criteria used in differentiating supratidal, intertidal and shallow subtidal carbonates. The relative frequencies of diagnostic features should not be taken as the only possible truth! Supratidal
Intertidal
Shallow subtidal
Sedimentary structures Fine-scale lamination Desiccation structures Tepees Flat-pebble conglomerates Collapse breccias
common abundant abundant rare abundant
abundant abundant rare common abundant
rare absent absent rare absent
Biogenic structures Microbial/algal mats Stromatolites Burrows
abundant rare rare
abundant abundant abundant
rare common abundant
Fenestral fabrics Irregular and laminoid birdseyes Tubular root voids
abundant abundant
abundant common
rare rare
Sedimentary grains Peloids Fecal pellets Benthic peloids Mud intraclasts Black lithoclasts Non-skeletal oncoids Skeletal oncoids Ooids Cortoids Aggregate grains
common rare rare common common rare rare absent absent absent
abundant abundant common abundant common common common common common common
common abundant absent common absent common common to abundant common abundant abundant
Biota: Characteristic groups Benthic foraminifera Ostracods Gastropods Bivalves Calcareous algae Cyanobacteria
very rare rare rare absent absent common
rare to common rare common to abundant rare rare abundant
common to abundant common common to abundant common common to abundant common
Biota: Diversity
very low
low
high
Diagenetic criteria Gravitational cement Meniscus cement Vadose pisoids Carbonate surface crusts Fine-grained dolomite Gypsum/anhydrite-carbonate associations
abundant common abundant abundant common abundant
common common rare common common abundant
absent absent absent absent common rare
Pedogenic structures
abundant
rare
absent
Evaporites Gypsum, anhydrite and halite occurring in association with peritidal carbonates are common in arid and semiarid areas. Upper intertidal and supratidal sabkha sediments are characterized by gypsum, landward areas by anhydrite nodules and chicken-wire anhydrite, often hydrated to gypsum. See Pl. 125/2, 4. Pedogenic structures Supratidal conditions characterized by prolonged subaerial exposure are reflected by dissolution and soil
formation connected with karstification in humid zones and calichification in arid and semiarid zones (see Sect. 15.1 and 15.2).
15.5.1.3 Synopsis of Diagnostic Criteria Lists of diagnostic peritidal features provided by various authors including Shinn (1983), Hardie and Shinn (1986) and Demicco and Hardie (1994), and the information in the preceding section show that it is difficult
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Devonian Peritidal Limestones
Plate 132 Peritidal Limestones: A Case Study from the Devonian The plate shows common facies-diagnostic criteria of cyclic peritidal and subtidal limestones forming shallowing-upward cycles. Important criteria of the intertidal and supratidal conditions are laminites and laminated microbial mats (–> 1), stromatolites (–> 5), fenestral fabrics (–> 6), reworked sediment and intraclasts (–> 4), and desiccation structures, e.g. mud cracks (–> 1). Synsedimentary reworking by tidal currents and by storms is a common feature (–> 2, 4). Intertidal mudstones and wackestones are frequently strongly burrowed (–> 3). The subtidal parts of the cycle can be recognized by fossils: Green algae (–> 7), ostracods (–> 8) and gastropods (–> 9) are common constituents of wackestones and packstones formed in the shallow subtidal environments. The taxonomic composition of the brachiopod fauna characterizes somewhat deeper parts of the subtidal zone. All samples are from the Eifelian Wojciechowice Formation of the Skaly quarry of the Holy Cross Mountains in central Poland. Supratidal and intertidal 1 Fine-laminated lime mudstone (bindstone) with small vertical desiccation cracks (arrows) indicating the shallowest supratidal parts of the cycle.The microfacies corresponds to Standard Microfacies Type SMF 19. 2 Intraformational breccia from topmost parts of intertidal laminites. The mixing of biotic elements from different habitats (ostracods in the micritic clasts, white arrow; brachiopod shells in the micritic matrix, black arrow) suggests a tempestite origin rather than an interpretation as tidal-channel deposit. 3 Transgressive basal breccia occurring at the sole of the subtidal part of the cycles. Micrite lithoclasts within a coarsegrained, dolomitized gastropod packstone matrix. Arrows point to burrows. The breccia is interpreted as surf-zone deposit formed during the initial phase of a transgressive episode. 4 Tidal-channel deposit. Interfingering of an intraformational tidal-channel breccia (BR) with intertidal laminated sediments (LS). Topmost part of intertidal laminites. 5 Detail of a hemispheroidal stromatolite occurring in well-defined horizons. Note the crinkled laminae consisting of micrite and peloids. 6 Fenestral algal bindstones occur in the lowermost part of the tidal sequence. The bindstone exhibits subhorizontal sparfilled voids bridged by microbial mats. The arrow points to planar fenestral structure. Note the oblique position of parts of the mat indicating incipient reworking. Intertidal. Shallow subtidal 7 Algal wackestone. The tubular structures are palaeoberesellid green algae regarded as dasyclad algae or as an independent green algal group (palaeosiphonoclads, see Fig. 10.17 ). This Devonian and Carboniferous group was usually restricted to calm, lagoonal shelf environments. 8 Ostracod wackestone. Non-disarticulated smooth-shelled ostracods (arrows) are abundant. Dissolution of the thin shells would result in the formation of ‘mold peloids’ (see Fig. 4.11). 9 Gastropod wackestone, slightly dolomitized. The shallower part of the subtidal unit of the cyclothem is dominated by gastropods in association with amphiporids (stromatoporoids, see Pl. 80/2). Devonian amphiporid stromatoporoids are abundant in black, bituminous limestones. The organisms were adapted to tranquil, reef-lagoon and back-reef habitats with diminished oxygen in bottom waters. Deeper subtidal 10 Brachiopod wackestone. Some Devonian brachiopods (e.g. Bornhardtina, shown here) and also Mesozoic bivalves (Fürsich 1993) were adapted to euryhaline conditions. Dense packing and stacking of shells is caused by high-energy events. Note the difference of the matrix type as compared with –> 9 (homogeneous micrite and calcisiltite). Deeper part of the cyclothem. –> 1–10: Skompski and Szulczewski 1994
Fig. 15.11. Idealized shallowing-upward succession in Middle Devonian platform interior carbonates showing main characteristics of intertidal-supratidal and shallow and somewhat deeper subtidal carbonates. See Pl. 132 for characteristic microfacies types. After Skompski and Szulczewski (1994).
Plate 132: Devonian Peritidal Limestones
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to distinguish between intertidal and supratidal carbonates if only a few facies criteria are used. You must try to find as many criteria as possible. Table 15.1 summarizes the main criteria of peritidal carbonates and gives an idea of the relative abundance of these criteria.
Basics: Peritidal Carbonates
15.5.2 Case Study: Middle Devonian Peritidal Carbonates from Poland
and fossils. Most peritidal carbonates are fine-grained bedded dolomites, differing in their yellow to brown rock color from the dark gray subtidal carbonates. This difference is also known from recent tidal carbonates, both in humid and arid settings. Facies-diagnostic criteria are small-scale parallel and crinkled lamination, hemispheroidal and isolated domal stromatolites, intraformational breccias, polygonal networks of up to 20 cm deep desiccation cracks, and small-scale channel-fill cross bedding.
Meter- to tens-of-meter-thick shallowing-upward cyclothems are common features of platform carbonates. Shallowing is indicated by changes in fossils, sedimentary structures and microfacies types. The case study from the Devonian of the Holy Cross Mountains in Poland (Skompski and Szulczewski 1994), documented in Pl. 132, describes a typical facies pattern showing some of the diagnostic facies criteria that allow subtidal and intertidal-supratidal carbonates to be differentiated. The Eifelian Wojciechowice Formation consists of lagoonal dolomites and limestones. Cyclic sedimentation is evident both in the limestone and dolomite part of the section studied (Fig. 15.11). The succession was formed in low-energy environments, as seen in the dominance of lime mudstones, wackestones and bindstones, and the rarity of grainstones and packstones. Rare high-energy episodes are indicated only by a brachiopod coquina (interpreted as a storm deposit), a basal breccia of one of the cycles, and few intraformational breccias at the top of intertidal laminites. Subtidal environments are derived from the composition and diversity of the biota comprising brachiopods, massive stromatoporoids, dendroid amphiporids, gastropods, ostracods, dasyclad calcareous algae as well as very rare trilobites and solitary rugose corals. Shallow subtidal and deeper subtidal assemblages have been differentiated by looking at the inferred ecological requirements of the biota. Shallow subtidal assemblages are characterized by amphiporids, gastropods, ostracods and calcareous algae. Deeper subtidal assemblages consist of large brachiopods (Bornhardtina, Emanuella) and massive stromatoporoid heads, and only subordinate gastropods and amphiporids. Six microfacies types can be distinguished based on the dominant fossils: (1) Brachiopod-stromatoporoid mudstone/wackestone, (2) brachiopod-peloid packstone, (3) Amphipora floatstone, (4) gastropod wackestone/packstone, (5) algal wackestone, and (6) ostracod wackestone. All the microfacies types are strongly burrowed. Carbonates interpreted as intertidal-supratidal are distinguished from the subtidal limestones by texture
Basics: Peritidal carbonates Alexander, C.R., Davis, R.A., Henry, V.J. (eds., 1998): Tidalites: processes and products. – Society of Economic Paleontologists and Mineralogists, Short Course Notes, 61, 171 pp. Fischer, A.G. (1964): The Lofer cyclothems of the Alpine Triassic. – In: Merriam, D.F. (ed.): Symposium on cyclic sedimentation. – State Geological Survey of Kansas, Bulletin, 169, 107-169 Ginsburg, R.N. (ed., 1975): Tidal deposits: a case book of Recent examples and fossil counterparts. – 428 pp., New York (Springer) Ginsburg, R.N., Hardie, L.A., Bricker, P.P., Garrett, P., Wanless, H.R. (1977): Exposure index: a quantitative approach to defining position within the tidal zone. – In: Hardie, L.A. (ed.): Sedimentation on the modern carbonate tidal flats of Northwest Andros Island, Bahamas. – 711, Baltimore (John Hopkins University Press) Galli, G. (1991): Mangrove-generated structures and depositional model of the Pleistocene Fort Thompson Formation (Florida Plateau). – Facies, 25, 297-314 Joachimski, M.M. (1994): Subaerial exposure and deposition of shallowing upward sequences: evidence from stable isotopes of Purbeckian peritidal carbonates (basal Cretaceous), Swiss and French Jura Mountains. – Sedimentology, 41, 805-824 Laporte, L.F. (1967): Carbonate deposition near mean sealevel and resultant facies mosaic: Manlius Formation (Lower Devonian) of New York State. – American Association of Petroleum Geologists, Bulletin, 51, 73-101 Hardie, L.A. (1986): Stratigraphic models for carbonate tidal deposition. – In: Hardie, L.A., Shinn, E.A. (eds.): Carbonate depositional environments, modern and ancient. Part 3: tidal flats. – Colorado School of Mines, Quarterly, 81, 59-74 Hardie, L.A., Shinn, E.A. (eds., 1986): Carbonate depositional environments, modern and ancient. Part 3: tidal flats. – Colorado School of Mines, Quarterly, 81, 74 pp. James, N.P. (1984): Shallowing-upward sequences in carbonates. – In: Walker, R.G. (ed.): Facies models. – Geological Association of Canada, Geoscience Reprint Series, 1, 213-288 Kendall, C.G.St.C. (1987): A review of the origin and setting of tepees and their associated facies. – Sedimentology, 34, 1007-1027 Klein, G. deV. (1971): A sedimentary model for determining paleotidal range. – Geological Society of America, Bulletin, 82, 2585-2592 Pratt, R., James, N.P., Cowan, C.A. (1992): Peritidal carbonates. – In: Walker, R.G., James, N.P. (eds.): Facies mod-
Carbonate Platforms and Ramps
els. Response to sea level change. – 303-322, Geological Association of Canada Sadler, P.M. (1994): The suspected duration of upward-shallowing peritidal carbonate cycles and their terminal hiatuses. – Geological Society of America, Bulletin, 106, 791802 Shinn, E.A. (1983): Tidal flat environments. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists, Memoir, 33, 171-210 Strasser, A. (1988): Shallowing-upward sequences in Purbeckian tidal carbonates (lowermost Cretaceous, Swiss and French Jura). – Sedimentology, 35, 369-383 Skompski, S., Szulszewski, M. (1994): Tide-dominated Middle Devonian sequence from the northern part of the Holy Cross Mountains (Central Poland). – Facies 30, 247266 Wright, V.O. (1984): Peritidal carbonate facies models: a review. – Geological Journal, 19, 309-325 Wright, V.O. (1990): Peritidal carbonates. – In: Tucker, M.E., Wright, V.O. (eds.): Carbonate sedimentology. – 137-164 Further reading: K020, K177
15.6 Carbonate Platforms and Ramps Facies models and microfacies types of rimmed carbonate platforms and carbonate ramps were discussed in Chapter 14. The objective of this section is to demonstrate the potential of microfacies analyses for differentiating and interpreting platforms and ramps by using selected case studies.
15.6.1 Ecological Controls on Platforms and Ramps The geometry and accumulation of sediment on carbonate platforms depend strongly on the biota living on the platforms, and the interplay of benthic organisms and physico-chemical environmental factors (Fig. 15.12, Fig. 15.13): • Ecologic factors control the location and rate of organic sediment production. The depth of the main carbonate production and its position in relation to platform margins affects the geometric evolution of platforms. Both parameters depend on the ecology of carbonate-producing organisms as reflected by their life habit. Carbonate production depth is controlled by illumination, and the lateral distribution of organisms across the platform is dependent on hydrodynamics, substrate and nutrients. • The presence or absence of a biogenic framework influence the stabilization and geometry of platform margins, and the composition and location of facies
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units. The presence of frameworks and reefs assists in coping with a rapidly rising sea level. The absence of frameworks may promote high rates of sediment reworking and export. • Sea-level fluctuations influence growth potential and the abundance of biogenic carbonate production on platforms. The frequency and growth potential of sediment-producing organisms may be different during transgressive and regressive phases. During transgressions large areas are flooded and biogenic production is high. During regressive phases, when large areas have emerged, the productive depth range is small. • The composition of skeletal grain associations depends strongly on the relative frequency and the taphonomic stability of skeletal grains. • The amount of biogenic sediment production influences the growth rates of platforms. Sediment produced on the platform is either deposited on the platform or exported. Maximum sediment production will normally occur, where the environmental parameters meet best the ecologic requirements of sediment producers and where environmental parameters remain constant over a longer time. • The early diagenesis of platform sediments is influenced by biofabrics that control the primary permeability and porosity, and by the mineralogy of skeletal grains. Sediments consisting mainly of aragonite and High-Mg calcite have a higher diagenetic potential than sediments dominated by Low-Mg calcite. Resulting differences in dissolution and reprecipitation will be particularly conspicuous in platforms and inner ramps affected by meteoric diagenesis. The Pomar Model: Carbonate-producing organisms and depositional profiles Facies models of platforms and ramps describing the depositional zones of shallow-marine to deep-marine transects were discussed in Chap. 14. A new model, developed by Pomar (2001a), emphasizes the interaction between the different sediment types being produced by organisms and the hydraulic energy acting in the diverse loci of production (Fig. 15.13). The basic ideas of the model are: (1) The differences between rimmed platforms, nonrimmed platforms, distally steepened ramps and homoclinal ramps, can be considered as a balance between (a) the type (basically the grain size) of the sediment produced, (b) the locus of the sediment production, and (c) the hydrodynamic energy. • The carbonate production (sediment supply) of platforms and ramps depends mainly on biological sys-
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Ecological Controls on Platforms
Fig. 15.12. Controls on platform development. Platform evolution and facies distribution depend to a large extent on the environment. The environmental controls are modulated by the ecosystems on the platform. Their impact affects parameters that are documented by microfacies criteria. The parameters emphasized in this sketch are derived from an analysis of Late Cretaceous platform carbonates, but are valid for most other tropical platforms. Modified from Stoessel (1999).
tems, and consequently on ecological factors governing the existence and distribution of organisms. • These factors predominantly include light, nutrient availability, water energy, water temperature and salinity. • Sediment dispersal including the export and redistribution of mud-, silt-, sand- and gravel-sized sediment, depends on the interaction between the sediment type (relative density, size and shape of skeletal grains) and the hydrodynamic energy existing instead of sediment production. • Sediment dispersal is frequently modified by biological processes (e.g. binding and baffling) and cementation.
shelves similar to modern reef platforms (Fig. 15.13A). • Gravel-producing biota living on soft substrates in the shallow euphotic zone create flat-topped open shelves (Fig. 15.13B). • Gravel-producing biota such as some larger foraminifera and red algae living in the oligophotic zone generate distally steepened ramps (Fig. 15.13C). • Mud-producing biota living in either euphotic or oligophotic zones generate homoclinal ramps. • Photo-independent carbonate-producing biota living above the wave-base give rise to open non-rimmed shelves and ramps, depending on the grain size, but may produce mounds if carbonate production occurs below wave base or sweeping current.
(2) Carbonate-producing biota differ strongly in their dependence on light: • Euphotic biota (autotrophs and mixotrophs) need good light conditions and live in shallow wave-agitated environments of the euphotic zone (see Sect. 12.1.4). Recent examples are green algae and scleractinian corals. • Oligophotic biota (autotrophs and mixotrophs) live under poor light conditions, either in shaded shallowwater zones or deeper on the shelf. The oligophotic zone on the shelf is characterized by a decrease in light and water temperature. Red algae and a number of larger foraminifera are characteristic of this zone. • Photo-independent biota (heterotrophs) are not controlled by light. These organisms can live in any environment and at all water depths, depending on the temperature, salinity and water energy, but are concentrated in the photic or aphotic zone according to the availability of nutrients and food. Examples are bryozoans, mollusks, crinoids, brachiopods, and sponges.
(4) Sediment dispersal on the shelf tends to form a sediment equilibrium profile. • The shelf equilibrium profile (Swift and Thorne 1991) is a balance between fluid power and sediment input. Fluid power depends on the spectrum of wave amplitudes that operate during fair-weather and stormweather conditions. • Sediment dispersal produces textural gradients and facies differentiation that can be visualized in terms of ‘hydraulic competence’ corresponding to the ability of water to transport detritus in terms of particle size rather than quantity. • Hydraulic competence is measured as the diameter of the largest transported particles.
(3) Euphotic, oligophotic and photo-independent biota produce different depositional profiles: • Euphotic framework-producing biota create rimmed
The advantage of the Pomar Model is that it relates microfacies data (texture, grain size, types of skeletal grains) to sediment production and dispersal, and to the major control mechanisms (sea level and accommodation). The usefulness of the concept has been demonstrated in the study of Miocene ramps and rimmed shelves (Pomar 2001b).
Pomar Model
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Fig. 15.13. Ecological controls on the depositional styles of carbonate platforms and ramps exemplified by the Pomar Model. This model explains the different kinds of platforms and ramps as a function of the type of sediment that was produced (reflected basically by grain size), the place (locus) of sediment production, and the hydraulic energy. The type of sediment and the locus of production depend on the carbonate-producing organisms whose distribution is governed by light and nutrients. Figs. A, B and C show which sediment types are produced (expressed by grain size) and by what processes and where the sediments accumulate contributing to the creation of different depositional slopes. Note the differences in the carbonate production curves and the curves of hydraulic competence (showing which maximum detrital grain sizes are transported). ‘Gravel’ as used in this figure refers to granule-, pebble- and cobblesized particles (2 to 256 mm). A: Euphotic framework-producing biota, e.g. modern corals, create rimmed shelves. Large and encrusting skeletons form shallow-water reefs in the zone with the highest water energy. Non-cemented fine-grained sediments are shed seaward, accumulating as talus, or back into the backreef area. Oligophotic biota may also contribute to the building of slope deposits or even basinal deposits, if the basin is shallow enough for light to reach the sea floor. Note that the figure does not consider the backreef lagoonal areas. B: Euphotic soft-substrate dwelling organisms (e.g. algae) produce gravelsized skeletons that create a flat-topped carbonate platform. Finer bioclasts are shed off the platform, creating a talus, whereas gravel-sized skeletons remain in situ, producing a zone of hydraulicenergy dissipation that may stabilize or baffle the sediment. C: Oligophotic organisms, such as some larger foraminifera and red algae, living in deeper commonly non-wave agitated areas, produce gravel-sized sediments and distally steepened ramps. A talus slope forms at certain water depths from increased sediment accumulation, in-situ gravel-sized skeleton production and fine-grained sediment swept from the shallower euphotic zone into deeper areas. Modified from Pomar (2001a).
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15.6.2 The Response of Carbonate Platforms to Drowning Shallow-marine platforms and reefs are drowned when rising sea level or tectonic subsidence outpaces carbonate accumulation and benthic carbonate production ceases (Schlager 1981). Drowned platforms present a fundamental paradox, because the growth potential of healthy carbonate platforms is thought to be possibly one or two orders of magnitude greater than long-term rates of sea-level rise. Prolonged periods of platform drowning events over a geographically restricted area result in drowning unconformities (Schlager and Camber 1986; Erlich et al. 1990) that can be recognized from geologic and/or seismic data. Drowning unconformities are related to an abrupt shift from shallowwater carbonate sedimentation to deeper-water clastic or carbonate sedimentation. A specific type of drowned platforms are pelagic carbonate platforms (PCPs; Catalano et al. 1977) corresponding to submarine intrabasinal highs, which are drowned fragments of ancestor shallow-marine platforms and ramps (Sanantonio 1994). These platforms are common at rifted continental margins bordered by synsedimentary faults. PCPs are sites of condensed and discontinuous pelagic sedimentation. They are recognized by a pile of shallow-water limestones capped by thin condensed pelagic sequence and bounded by paleofaults. Other terms used for these depositional sites are seamount, swell, plateau, guyot or ‘haut-fond pélagique’. The terms seamount and guyot should be restricted to submarine highs built on oceanic crust (see Sect. 15.7).
Drowning mechanisms The death of carbonate platforms is caused by different mechanisms, including long- and short-term sealevel fluctuations, tectonic processes, climate and related environmental changes. Long-term geodynamic processes triggering sea-level history may lead to the drowning of platforms and reefs, if the reduction of growth with increasing time is largely caused by environmental factors (Schlager 1999). In many cases, drowning might have resulted from an interplay of differently scaled factors as shown by the study described in Sect. 15.6.2.2. • The increased and rapid pulses of relative sea-level rise caused by tectonic plus eustatic events (Schlager 1981; Dominguez et al. 1988; Purdy and Bertram 1993). This is the simplest explanation and may explain worldwide drowning events.
Platform Drowning
• The death of the platform by repeated emergence and submergence, indicated by criteria pointing to subaerial exposure (Schlager 1998; Wilson et al. 1998). • Environmental deterioration (e.g. loss in water quality; decrease or increase in water temperature, nutrients, salinity, water energy; see Erlich et al. 1990). This explanation has the advantage that at least some of the environmental factors can be traced by paleontological and geochemical proxies; changes in the environment may be recorded in depositional textures of limestones, paleocommunities and in changes in microfacies types. Drowning induced by environmental factors can lead to the decrease or loss of the growth potential of reef and platform organisms. Nutrient-rich surface waters of the equatorial upwelling zone may have been responsible for the ‘death-in-the-tropics’ of Cretaceous platforms (Wilson et al. 1998). Changes in nutrient conditions affecting the conditions favorable for carbonate producing organisms may reduce the aggradation potential of the carbonate factory and lead to an increase in accommodation space. The sedimentary signature of reduced carbonate production and the increase in accommodation space is the formation of sedimentary discontinuities separating neritic from deeper water sediments. The increase in nutrient levels and the decrease in sea-surface temperatures associated with an acceleration of the biological pump and atmospheric CO2 draw-down during oceanic anoxic events have been invoked to explain the drowning of Cretaceous platforms (Hallock and Schlager 1986; Rougerie and Fagerstrom 1994). • Oversteepening and self erosion of platform margins. Drowning may occur by backstepping of the platform rim and related partial multiple drownings (Erlich et al. 1990; Betzler et al. 1995). Sediment bypass from the platform to deeper water will result in a sediment accumulation on the platform that is lower than the sealevel rise, thus terminating platform growth. • Burial of carbonate platforms by prograding siliciclastics. Platforms may be suffocated by shelf sands or prodelta shales (Schlager 1989).
15.6.2.1 Microfacies Signals of Drowning History Capped by shales or tight limestones and often exposed and leached prior to drowning, drowned platforms carbonates are stratigraphic traps for hydrocarbons (e.g. Miocene platforms, Pearl River Mouth Basin, offshore China). Understanding the origin of drowning allows us to make better reservoir and seal predictions. Microfacies analysis of core material offers a possibility to
Microfacies of Platform Drowning
describe drowning effects in different parts of platforms and slopes and to distinguish unconformities caused by drowning from those caused by subaerial exposure. A discussion of discontinuity surfaces, including submarine condensation surfaces that may be associated with the drowning of platforms, can be found in Sect. 5.2. Indications of drowning • Successions of shallow-marine neritic sediments rapidly passing upward into deep-marine deposits. • Vertical replacement of shallow-water grain assemblages by deeper-water grain assemblages. • Hardgrounds with crusts of oxide, phosphate and glauconite divide neritic and deep marine deposits, and indicate a period of non-deposition. • Burrowed discontinuities. • Reefs: Increasing reworking of platform-margin reef material leading to the destruction of protecting rims. The loss of protecting elevated margins results in stronger reworking of upper slope sediments and platform sediments. This may be reflected by an increase in the abundance of reef- and platform-derived material in slope sediments. • Reefs: Significant changes in dominant growth forms of reef-building organisms, e.g. from massive to low-laminar, or massive to high-dendroid growth forms, indicate adaption to changes in water energy and depth. • Abrupt vertical changes in dominating microfacies types. • Grainstones and packstones with abundant shallowmarine biota, intercalated in slope- and basinal sediments. • Grainstones and rudstones with abundant platformderived rounded lithoclasts, intercalated in slope sediments (see Sect. 15.7). • Increase in microborings and the frequency of cortoids (see Sect. 15.7.1). • Increase in biogenically encrusted grains, indicating longer periods of ceased sedimentation. • Decrease in thin-section diversity. Environmental stress related to drowning may contribute to the development of monomict biotic associations. • Mineralized microbial crusts. • Fe-, Mn- and P-enrichments are common in crusts of drowning and post-drowning carbonates and can be related to a primary higher nutrient content in the water column (Föllmi et al. 1994; Drzwiecki and Simo 1997) reflecting changes from oligotrophic to eutrophic conditions. • Increase in suspension-feeders and grazers, e.g. echinoderms and gastropods.
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• Increase in open-marine biota, associated with shallow-marine constituents. • Increase in ‘deep-water’ non-photic organisms, e.g. sponges. • A shift of the benthos/plankton ratio towards the plankton. Many carbonate platforms suffered subaerial exposure prior to drowning by renewed flooding. Exposure phases are indicated by vadose-meteoric cement types, pedogenic structures and paleokarst (e.g. Bernecker et al. 1999; see also Sect. 15.1 and Sect. 15.2).
15.6.2.2 Case Study: Platform Drowning Reflected by Microfacies The Jbel Bou Dahar in the Atlas Mountains of Morocco represents an Early Jurassic carbonate platform that was drowned at the beginning of the Toarcian. The pre-drowning phase (Late Sinemurian to Late Pliensbachian), drowning phase (Late Pliensbachian to Early Toarcian) and post-drowning phase (Early Toarcian to Aalenian) were studied by Blomeier and Reijmer (1999) with regard to facies changes reflected by qualitative and quantitative microfacies criteria. Pl. 133 and Fig. 15.14 summarize the development of the platform over time. Methods: The thin sections were studied with respect to sediment composition and fossils. Quantitative analyses were carried out with point-counting, using 200 points. A total of 25 categories were differentiated, and subsequently grouped into 8 point-count groups comprising matrix (micrite and cement), shallow-marine biota (living on the platform top and the platform margin), suspension feeders and grazers (echinoderms and brachiopods characteristic of the drowning phase), open-marine biota (marine organisms characterizing deep-marine environments; e.g. ammonites, siliceous sponges, filaments), encrusted grains (including zoogenic oncoids, micro-oncoids and aggregate grains), non-skeletal grains (ooids and litho- and extraclasts), peloids, and terrigenous input (mainly detrital quartz). Every phase is characterized by an association of characteristic microfacies types: Pre-drowning phase During this phase, the platform interior was characterized by restricted-marine lagoonal environments with limestone/marl alternations separated by distinct exposure surfaces, and protected by scattered buildups and cemented debris at the platform margin. Common
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Drowning Stages
A
B
C
Case Studies: Platform and Ramps
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<<< Fig. 15.14. Microfacies reflecting drowning history of a carbonate platform. Case study Jbel Bou Dahar, High Atlas (Early Jurassic). Modified from Blomeier and Reijmer (1999). A: Changes in the sediment composition within different stages of drowning and in different platform sections. The changes in the diagnostic criteria are derived from quantitative microfacies analysis. Note the distinct differences in shallow-marine biota, frequency of suspension feeders and grazers (mainly echinoderms and brachiopods), open-marine organisms (pelagic filaments, ammonites, siliceous sponge spicules) and the frequency of lithoclasts (intra- and extraclasts). B: Platform morphology and distribution of microfacies types within different phases of pre-drowning, drowning and postdrowning. SMF: Standard Microfacies Types (see Box 14.1 for definitions), MF: local microfacies types. Arrows mark the main directions of sediment transport. C: Sea-level variations and drowning stages. 1a: Pre-drowning phase and sea-level highstand. Carbonate production takes place in the flooded platform interior. 1b: Pre-drowning phase and sea-level lowstand. Subaerial exposure of the platform top, the carbonate production has shifted to the platform margin and upper slope. 2: Renewed flooding of the platform top at the beginning of the drowning phase. 3: Gradual transition to deep-marine sedimentation during the post-drowning phase.
microfacies types include wackestones (pelmicrites and biopelmicrites corresponding to SMF 22; Pl. 133/1), mudstones (micrites, dismicrites; SMF 21), fenestral microbial mats (SMF 19), and intercalated grain- to packstones (pelsparites; SMF 17). The upper and midslope were dominated by thick-bedded, coarse-grained, poorly sorted limestones, deposited through debris flows during sea-level lowstands. Microfacies types are wackestones/floatstones and grainstones/rudstones dominated by shallow-marine bioclasts, peloids and encrusted grains (Pl. 133/2). Sea-level highstand deposits occurring at the toe-of-slope consist of mediumbedded marls and turbiditic limestones (grain- to packstone with abundant, highly diverse shallow-marine skeletal grains and peloids; SMF 4) marls and hemipelagic limestones (wackestones characterized by bioturbated biopelmicrites with tiny grains; SMF 9; Pl. 133/3). Drowning phase This phase is characterized by a condensed section and a fundamental environmental change along the entire platform expressed by a distinct facies change. The base of the drowning succession on the platform top is marked by the abrupt occurrence of red-brownish carbonate beds, unconformable overlying sediments of the pre-drowning phase, and showing totally different microfacies. The microfacies differs in the large quantities of fragments of suspension feeders and grazers (mainly echinoderms) and recrystallized bioclasts of shallow-marine platform biota (Pl. 133/4). These particles were deposited in three pulses with decreasing absolute amounts. Densely packed skeletal grainstones represent the switch to high-energy environments within the platform interior, causing erosion of former pre-drowning lagoonal sediments. The erosional products were rede-
posited on the platform slope. The drowning succession of the mid-slope consists of thin- to medium-bedded, red-brownish carbonates and marls, unconformably onlapping at a low angle the thick limestone beds of the pre-drowning phase. The most common microfacies are packstones and wackestones (biomicrites and biopelmicrites; Pl. 133/5). The composition of these limestones differs drastically from that of the carbonates of the pre-drowning phase. Non-skeletal grains, especially lithoclasts, dominate the grain spectra (Pl. 133/6). In contrast to the predrowning phase, encrusted grains are almost completely missing. Intercalations of brachiopod and oyster coquinas are common. Platform interior and slope successions exhibit cyclic variations in sediment composition, possibly reflecting small-scale sea-level fluctuations. Post-drowning The transition to the post-drowning phase is conformable and gradual. During this phase, thick-bedded marls and thin-bedded limestones (wackestones and mudstones; biopelmicrites and micrites; SMF 3; Pl. 133/7) were deposited elsewhere on the platform, followed by the deposition of calcisiltites (SMF 2; Pl. 133/ 8). Common skeletal grains are derived from suspension feeders and grazers, and open-marine biota. Shallow-marine biota are almost completely absent. The facies types indicate uniform, deep-marine and quietwater sedimentation below the storm-wave base.
15.6.3 Case Studies: Platform and Ramp Carbonates Many carbonate platforms exhibit recurring microfacies patterns that reflect the hydrodynamic conditions in specific parts of the platforms and the varying compo-
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sition of the sediment. Distributional patterns of calcareous algae and benthic foraminifera as well as the occurrence of particular grain types (e.g. ooids, oncoids) are important facies-diagnostic criteria. In addition, cement types and other diagenetic features can be used in recognizing subaerial exposure and sea-level falls. The following three case studies on platforms of different types and ages are based on the investigation of geological sections or core sections and the interpretation of microfacies data derived from thin-section analyses.
Platform Drowning: Morocco
15.6.3.1 A Late Jurassic Bahamian-Type Carbonate Platform from the Northern Calcareous Alps The example shown in Pl. 134 comes from Sulzfluh limestone at the Austrian/Swiss border (Ott 1969; Flügel 1979), a Late Jurassic (predominantly Tithonian) non-attached platform resting on tectonic highs within a deep ocean. Sulzfluh limestone is a massive and thickbedded white limestone exposed in spectacular mountain cliffs. The limestone is generally very poor in nonText continued on p. 768
Plate 133 Microfacies of Platform Drowning: Early Jurassic of the High Atlas (Morocco) Drowning of carbonate platforms is reflected by drastic changes in microfacies and in dominant grain types. The carbonates of the Jbel Bou Dahar platform in the High Atlas of Morocco documents environmental changes in platform interior areas and various parts of the slope (Blomeier and Reijmer 1999). Profiles comprise all phases of platform growth and demise. Bed-per-bed studies and microfacies data combined with the investigation of large-scaled platform geometries and depositional features show that the ultimate drowning resulted from a series of disadvantageous processes starting with (a) the destruction of former reef communities related to subaerial exposure caused by high-frequency sea-level fluctuations at the end of the pre-drowning phase, followed by (b) local environmental effects during the renewed flooding of the platform top, hampering the re-establishment of a flourishing carbonate-producing community, (c) prevailing high-energy conditions during almost the entire drowning phase, resulting in continuous transport of eroded platform-top material across the margins to the slope, (d) competition between better adapted echinoderms and mollusks with the carbonate-producing organisms, and (e) last but not least regional/global changes in ocean circulation patterns, producing unfavorable conditions for platform growth during the drowning phase. The plate shows characteristic microfacies types of the pre-drowning, drowning, and post-drowning phases (see Fig. 15.14A). The drowning phase might have lasted several hundred thousands of years. Pre-drowning phase 1 Platform interior. Laminated pelmicrite. Standard Microfacies Type SMF 22. 2 Upper slope. Coarse, poorly sorted grainstone/rudstone. The recrystallized skeletal grains are encrusted by foraminifera and bryozoans. Arrows point to composite grains. SMF 5. 3 Lower slope. Fine-grained, well-sorted packstone. Grains are peloids, terrigenous quartz, sponge spicules, and thin-shelled mollusks. White arrows point to burrowing. Drowning phase 4 Platform top. Bioclastic grainstone. Most grains are echinoderms and recrystallized shells. The black grain is an extraclast. Characteristic for the drowning phase is the abundance of suspension feeders (echinoderms). Note the clear syntaxial cement enclosing the echinoderm grains. These cements are common in grainstones deposited on the top of the platform during the drowning phase. SMF 12. 5 Mid-slope. Coarse-grained, poorly sorted rudstone consisting of large rounded extraclasts and smaller mud peloids. The extraclasts are a characteristic criteria of the drowning phase. Note the isopachous cement crusts on the lithoclasts (arrows). This cement type occurs exclusively within the platform interior and slope sediments of the pre-drowning and drowning phases. 6 Mid-slope. Densely packed wackestone (biopelmicrite). Grains are peloids, micritic extraclasts, some shells and echinoderm fragments embedded within a heterogenous microspar matrix. Post-drowning phase 7 Mid-slope. Bioclastic wackestone (biomicrite). The microfacies is characterized by pelagic filaments associated with phosphatized and bored echinoderm fragments. Note the absence of allochthonous shallow-marine biota. SMF 3. 8 Basin. Densely packed, quartz-rich, well-sorted calcisiltite, deposited adjacent to the lower slope during the last postdrowning phase. The arrows point to burrowing. SMF 2. –> 1–8: Blomeier and Reijmer (1999).
Plate 133: Microfacies of Platform Drowning
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Microfacies of a Late Jurassic Platform
Plate 134 Microfacies of a Late Jurassic Bahamian-Type Platform: Sulzfluh (Graubünden, Switzerland) The plate displays characteristic microfacies types of the platform limestones and microfacies criteria useful for paleoenvironmental interpretation. Mudstones are represented by fenestral mudstones (–> 4, platform interior). Peloidal and bioclastic wackestones occur in protected platform interior settings and open-marine parts of the platform. Packstones, grainstones and rudstones are most common in open platform environments with good water circulation. They comprise at least five types of skeletal grainstones, grainstones with aggregate grains and/or coated and micritized grains (–> 1), and ooid grainstones forming platform margin shoals (–> 7). Small reefs were formed within the sand flats. Frame-building organisms are corals (–> 2), chaetetid coralline sponges and encrusters (foraminifera, various microproblematica, e.g. Bacinella and Lithocodium). Most sediments were deposited at very shallow depths, as indicated by the abundance of dasyclad algae (–> 1), microborers and meteoric-vadose cement (–> 7), infillings of crystal silt in sheet cracks (–> 8) and fossil molds. Facies-diagnostic fossils are calcimicrobes (–> 9), dasyclad green algae (–> 5) and larger foraminifera (–> 6). Characteristic microfacies types 1 Poorly sorted intraclastic grainstone (intrasparite). Open-marine part of the platform. Larger black grains are aggregate grains, smaller grains are peloids. Many grains are reworked cyanobacterial (C) and algal fragments. CL: encrusted dasyclad alga (Clypeina). Open platform interior. Facies Zone FZ 7 (see Fig. 14.1). The sample corresponds to Standard Microfacies Type SMF 11. 2 Coral framestone. Patch reef. Meter-sized coral patch reefs occur on the open platform. Corals and encrusting organisms (Bacinella, arrows) form a rigid framework. Bacinella acted as stabilizers for the framework (see Sect. 10.2.6.2). 3 Peloid grainstone (pelsparite) with oncoids. Open platform. Reworking is indicated by poor sorting. Reworking may have taken place by bioturbation, tidal currents, or stirring of the sediment by storms. 4 Laminoid fenestral mudstone (dismicrite). Restricted platform interior. The fabric characterizes a tidal flat facies. The thin irregular fenestrae, some of which resemble plant roots, point to supratidal muds. Facies Zone FZ 8. Standard Microfacies Type SMF 21-FENESTRAL. Facies-diagnostic criteria 5 Calcareous green algae: Clypeina (see Pl. 62/5, 6). The distribution of algae is facies-controlled, allowing the platform to be differentiated internally (see Fig. 10.19). Restricted lagoonal environments of the platform are characterized by wackestones with abundant Clypeina. The occurrence of Clypeina in grainstones of the open platform indicates transport of lagoonal grains to the open platform parts. 6 Benthic foraminifera. Larger foraminifera are strongly substrate-controlled and limited to distinct parts of platforms. Example: Kurnubia (Textulariina; see Pl. 69/3), characterizing high-energy open parts of platforms. 7 Ooids forming oolite shoals are common in platform-margin position but also occur in high-energy areas on the platform. Example: Grainstone consisting of tangentially structured micritized ooids. The irregular shape of some ooids results from biogenic encrustations, indicating only weak reworking of the grains. The grains are surrounded by symmetric circumgranular calcite cements, confirming meteoric-phreatic overprint. Platform margin. SMF 15-C. 8 Early diagenetic sheet crack within grainstones, outlined with scalenohedral calcite. The remaining cavity is filled with crystal silt (CS), indicating meteoric-vadose diagenesis possibly caused by subaerial exposure. Platform interior. 9 Skeletal cyanobacteria. Calcification intensity and abundance of cyanobacteria vary in different parts of inner platform environments. Example: ‘Cayeuxia’, characterized by parallel arrangement of tubes subsequent to their bifurcation. Compare Pl. 53/1-3. Fig. 15.15. Microfacies of a Jurassic Bahamian-type platform. Generalized facies zones of the Late Jurassic Sulzfluh platform. The sketch is based on paleontological and microfacies data in the SulzfluhWeißfluh area near the SwissAustrian boundary. Lithofacies and biofacies are similar to facies types of protected and open interior parts of the Holocene Bahama platform (compare Pl. 3). Not to scale. Modified from Flügel (1979).
Plate 134: Microfacies of a Late Jurassic Platform
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Microfacies of a Devonian Ramp
Plate 135 Microfacies of Ramp Environments: Middle Devonian of Graz (Austria) The ‘Barrandei Limestone’ outcropping near Graz in Styria (Austria), named after a heliolitid tabulate coral species (see Pl. 83/5, 6) is an example of carbonates deposited in near-coastal, subtidal ramp environments influenced, but not controlled by siliciclastic influx. The up to 40 m thick sequence consists of dark-blue to black limestones that alternate with calcareous marls and shales. Both limestones and marls, yield abundant fossils, predominantly crinoids, brachiopods, tabulate and rugose corals, and stromatoporoids. The limestones represent a low-energy mud facies with mudstones, wackestones and packstones (–> 3), a fossiliferous high-energy mud facies (–> 1, 2, 4–7), a high-energy debris facies with tempestites, and a reefoid facies with tabulate corals, stromatoporoids and udoteacean green algae. Microfacies, fossils, and geochemical data indicate sedimentation on the inner part of a ramp with shallow shelf lagoonal and somewhat deeper open-marine environments (Hubmann 1993). The input of clay, silt and landderived organic matter (Hubmann 2000) is not reflected by differences in the development of microfacies types over time, but by the changes in the quantitative composition of skeletal grains and deviations of the normal growth patterns of favositid tabulate corals. The crinoid columnals exhibit broad variations with regard to the number and shape of central canals (–> 4–7) allow morphotypes to be differentiated. They have become increasingly more important both in facies analyses and biostratigraphy (Le Menn 1987; Ebert 1994). All samples shown are from dark, bedded Middle Devonian (Eifelian) limestones exposed at the Fürstenstand on the Plabutsch Mountain west of Graz, Austria. This locality is characterized by a significantly high abundance of fossils (Heritsch 1943). From the Middle Ages on, Devonian limestone has been used for building material in the city of Graz. 1 Bioclastic coral-brachiopod wackestone/floatstone. The limestone is characterized by diverse thick-shelled brachiopods (B), dendroid tabulate corals, and crinoids (C). The black arrow points to a cross section of a brachiopod shell cut oblique to the hinge line and showing the interior plates. The shell was infilled with sediment and coral fragments. The white arrow points to infilling with large peloids (fecal pellets). Dendroid tabulate corals are represented by cross-sections of Thamnopora (TH) and Syringopora (SY). The inhomogeneous texture of the micritic-siltitic matrix indicates strong burrowing. High-energy mud facies. This sample corresponds to Standard Microfacies SMF 9 and Facies Zone FZ 8 in the Wilson Model, and to the Ramp Microfacies RMF 3 common in mid-ramp and outer-ramp settings (see Fig. 14.30). The biotic assemblage of this facies type is different than the standardized Middle and Late Devonian shallow-marine carbonates (Machel and Hunter 1994; Fig. 14.9 and Fig. 14.10). 2 Detail of the sample shown in –> 1. Tabulate corals (cross and oblique sections of Thamnopora, TH), abundant crinoid fragments (C) and brachiopod shells (B). The micritic matrix contains quartz silt. 3 Calcisphere packstone. Circular and elliptical spar-filled sections are calcispheres, most probably algal spores. Note the highly variable shapes and the wide size range of the spheres. The interpretation as algal spores is supported by the common occurrence of udoteacean green algae in the studied section. The arrow points to a recrystallized udoteacean thallus. Low-energy facies. In the standardized Devonian reef model (Machel and Hunter 1994), this facies type would be assigned to facies I-f, characterizing sediments deposited in quiet water below the wave base in an off-reef setting. Using a ramp model, this microfacies would be interpreted as a sediment deposited in an inner ramp, lagoonal environment (Ramp Microfacies RMF 17; see Fig. 14.30). This interpretation is in agreement with the geological and geochemical criteria of the Barrandei limestones. Using a platform model, the microfacies would correspond to SMF 18, characterizing a bank interior environment (FZ 8 of the Wilson model). 4 Characteristic fossils. Cross-section of Thamnopora boloniensis (TH), crinoidal columns and trilobites (TR). The black arrow points to a thin-shelled brachiopod. The section resembles cross sections of vermetid gastropods and is reminescent of a cross section of an ammonoid shell. However, these shells have different microstructures. This identification as a brachiopod shell is based on the occurrence of fine pores within the shell. Note the strong biogenic encrustation of the brachiopod by cyanobacteria (Girvanella). The white arrow points to the double-hole structure within a crinoid columnal (C). High-energy facies. 5 Characteristic fossils. Cross sections of crinoid columnals exhibiting dumbbell-shaped central structures. The variations in the central structures allow columnal morphotypes to be distinguished. Types and abundance of these morphotypes vary throughout the Paleozoic (Stukalina 1988). Tabulate corals are represented by two genera (Thamnopora, TH; and longitudinal and cross sections of Syringopora, SY). Note the difference in the size of the corallites of Thamnopora as compared with –> 4, indicating the existence of two species (–> 4: Thamnopora boloniensis; –> 5: Thamnopora reticulata). 6 Differentiation of crinoid elements. Crinoid columnal exhibiting the canal pattern of Cupressocrinites characterized by four peripheral canals separated from the central lumen. Cupressocrinites is a widely distributed genus known from the Late Emsian to the Early Frasnian (Ures et al. 1999). 7 Crinoid packstone. Accumulation of small crinoidal columnals with two small, closely set canals of equal size.
Plate 135: Microfacies of a Devonian Ramp
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Microfacies of Ramp Carbonates
Plate 136 Microfacies of a Late Jurassic to Early Cretaceous Ramp: Subsurface of Bavaria (Germany) The plate exhibits the microfacies of ramp carbonates in the Kinsau well near Augsburg in southern Bavaria. Fig. 1 on Pl. 136 is the lowermost, Fig. 10 the topmost sample. The sequence starts with Tithonian shallow-marine, sometimes lagoonal carbonates (–> 1, 3–5), and coral reef limestones (–> 2; maximum size of the patch reefs in outcrops 20 m), overlain by ramp deposits representing the Early Cretaceous ‘Purbeck facies’ (Berriasian; –> 6–10). This facies known from England, France and Switzerland, occurs in southern Germany and northern Austria in the subsurface. Purbeck sediments formed in strongly varying environments, and comprise marine, brackish, freshwater and hypersaline sediments (see Fig. 4.26 for an example). Other microphotographs can be found on Pl. 16/7 (tidal intraclasts), Pl. 29/3 (fabric-selective oomoldic porosity), Pl. 65/5, 6 (charophycean brackish-water algae), and Pl. 93/5 (Favreina-peloid sand). The shallowing-upward sequence consists of outer, mid- and inner ramp deposits (Fig. 15.16). Outer ramp sediments, deposited below the normal storm wave base, comprise mud-rich subtidal carbonates and marls (–> 1), small patch reefs (–> 2), and shallow subtidal limestones with red algae (–> 3 and 5). The existence of midramp sediments is indicated by Larger foraminifera (compare Pl. 69/1) and micrite oncoids (–> 6). Low-energy inner ramp deposits are indicated by oncoid grainstones with porostromate cyanobacteria (–> 4). High-energy inner-ramp sediments are represented by sand shoals and sheets (–> 7–9) controlled by tidal currents. For a coastal sabkha see –>10. Porosity and reservoir qualities are controlled by primary depositional facies (geometry and grain composition of high-energy shoals; micritic seals) and facies-selective diagenesis (freshwater dissolution of ooids, repeated dolomitization). 1 Deeper outer ramp (A in Fig. 15.16): Fine-grained bioclastic packstone. Grains are echinoderms (white), red algal clasts (black) and coated shell fragments. The packstone is intercalated within mud- and wackestones, pointing to downslope sediment transport. 2 Patch reef (B in Fig. 15.16): Framestone consisting of intergrown corals (C), chaetetid coralline sponges (CH) and foraminiferal crusts (F). Arrows point to boring. Outer ramp. Ramp microfacies RMF 12 (see Fig. 14.30 for definitions). 3 Deeper outer ramp (C in Fig. 15.16): Red algal packstone composed of rhodoids formed by Marinella Pfender. This alga has millimeter-sized thalli and extremely fine threads (< 5 to 12 µm). The alga is known from the Late Jurassic to the Oligocene, and is common in Jurassic and Early Cretaceous platform and reef carbonates (Leinfelder and Werner 1993). At smaller magnifications, the micritic thalli can be mistaken for micritic clasts. Note intergranular dissolution (white) due to strong compaction. 4 Inner ramp (D in Fig. 15.16): Oncoid rudstone. Larger grains are filamentous cyanobacteria (C) and gastropods (G), small grains are algal peloids. Blocky cements and infill of intergranular voids with crystal silt (arrow) indicate meteoric overprint. Restricted intertidal environment. RMF 21. 5 Outer ramp (C in Fig. 15.16): Rhodoid packstone with Marinella (M). The facies interfingers with facies D. Low-energy subtidal environment. RMF 9. 6 Protected lagoonal environment (E in Fig. 15.16): Oncoid wackestone with micrite oncoids and large lituolinid foraminifera (Pseudocyclammina, P), indicating a mid-ramp position. The micritic matrix is selectively dolomitized. RMF 13. 7 Inner ramp (F in Fig. 15.16): Ooid grainstone with simple and compound, tangential ooids. Intergranular pores are occluded by calcite cement, some of which exhibits gravitative textures indicating vadose meteoric diagenesis. High energy inter- and subtidal shoals. RMF 29. 8 Intertidal shoals on the inner ramp (G in Fig. 15.16): Fine-grained peloid grainstone. Most grains are fecal pellets as indicated by their rodlike shape. 9 High-energetic shoals of the inner ramp (H in Fig. 15.16): Favreina grainstone consisting predominantly of fecal pellets of decapod crustaceans (see Sect. 10.2.4.7). RMF 27. 10 Back-ramp sabkha (I in Fig. 15.16): Evaporitic argillaceous mudstone composed of alternating mudstone and layers with anhydrite (white). Supratidal environment. Standard Microfacies Type SMF 25.
Fig. 15.16. Generalized facies distribution of a Late Jurassic/Early Cretaceous carbonate ramp based on data from the subsurface of southern Bavaria, Germany.
Plate 136: A Late Jurassic/Early Cretaceous Ramp
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carbonate residues. The platform exhibits two major facies, mud-dominated and grain-dominated limestones. Maximum thickness is 150 m. The size of the platform is difficult to evaluate owing to the strong tectonic faulting in the whole region, but might be within the range of several kilometers. Fig. 15.15 shows the distribution of facies and subfacies types representing Facies Zones FZ 9A (tidal flat), FZ 8 (lagoon), FZ 7 (open platform) and FZ 6 (platform margin sands) of the Wilson Model. Other photographs of Sulzfluh limestone samples can be found on Pl. 53/3 (porostromate cyanobacteria), Pl. 62/ 2, 3, 6 (dasyclad algae), Pl. 69/4, 6, 10 (foraminifera) and Pl. 127/3 (microfacies). The facies and biota of the isolated Sulzfluh platform are similar in many respects to the Late Jurassic Plassen platforms of the Northern Calcareous Alps, although the Sulzfluh and Plassen platforms originated in different paleo-oceans. 15.6.3.2 A Middle Devonian Ramp from Graz, Southern Austria Many Devonian shallow-marine limestones are amazingly uniform all over the world with regard to their gross biotic composition, which includes predominantly tabulate and rugose corals, stromatoporoids, brachiopods and crinoids. These organisms occur in platform and ramp settings as well in large reefs formed at the margins of attached and isolated platforms and in smaller reefs formed on platforms. Pl. 135 exhibits some common and widely distributed types of Devonian limestones formed in ramp settings. Ramp deposits can be differentiated in more detail using the taxonomic composition of fossil assemblages.
15.6.3.3 A Late Jurassic/Early Cretaceous Ramp from the Subsurface of Southern Bavaria, Germany Late Jurassic/earliest Cretaceous homoclinal carbonate ramps covered by Tertiary sediments of the Molasse trough have been studied in the course hydrocarbon exploration in southern Germany. The main targets were reef structures and shallow-marine sand bodies. Common microfacies types are shown on Pl. 136. Basics: Carbonate platforms and ramps Blomeier, D.P.G., Reijmer, J.J.G. (1999): Drowning of a Lower Jurassic carbonate platform: Jbel Bou Dahar, High Atlas, Morocco. – Facies, 41, 81-110 Erlich, R.N., Barrett, S.F., Ju, G.B. (1990): Seismic and geologic characteristics of drowning events on carbonate plat-
Basics: Carbonate Platforms and Ramps
forms. – American Association of Petroleum Geologists, Bulletin, 74, 1523-1537 Erlich, R.N., Longo, A.P., Hyare, S. (1993): Response of carbonate platform margins to drowning: evidence of environmental collapse. – American Association of Petroleum Geologists, Memoir, 57, 241-266 Flügel, E. (1979): Paleoecology and microfacies of Permian, Triassic and Jurassic algal communities of platform and reef carbonates from the Alps. – Bulletin, Centres des Recherches Exploration-Production Elf-Aquitaine, 3, 569-587 Grötsch, J., Flügel, E. (1992): Facies of sunken Early Cretaceous atoll reefs and their capping Late Albian drowning succession (Northwestern Pacific). – Facies, 27, 153-174 Handford, C.R., Loucks, R.G. (1993): Carbonate depositional sequences and system tracts - responses of carbonate platforms to relative sea-level changes. – American Association of Petroleum Geologists, Memoir, 57, 3-41 Hubmann, B. (2000): Die Rannach- und Schöckel-Decke des Grazer Paläozoikums. – Mitteilungen der Gesellschaft der Geologie- und Bergbaustudenten in Österreich, 44, 1-44 Pomar, L. (2001a): Types of carbonate platforms: a genetic approach. – Basin Research, 13, 313-334 Pomar, L. (2001b): Ecological control of sedimentary accommodation: evolution from a carbonate ramp to a rimmed shelf. Upper Miocene, Balearic Islands. – Palaeogeography, Palaeoclimatology, Palaeoecology, 175, 249-272 Schlager, W. (1981): The paradox of drowned reefs and carbonate platforms. – Geological Society of America, Bulletin, 92, 197-211 Schlager, W. (1989): Drowning unconformities on carbonate platforms. – In: Crevello, P.D., Wilson, J.L., Sarg, J.F., Read, J.F. (eds.): Controls on carbonate platforms and basin development. – SEPM Special Publication, 44, 15-25 Schlager, W. (1998): Exposure, drowning and sequence boundaries on carbonate platforms. – In: Camoin, G., Davies, P. (eds.): Reefs and carbonate platforms of the W Pacific and Indian Oceans. – International Association of Sedimentologists, Special Publication, 25, 3-21 Schlager, W. (1999): Scaling of sedimentation rates and drowning of reefs and platforms. – Geology, 27, 183-186 Stoessel, I. (1999): Rudists and carbonate platform evolution: the Late Cretaceous Maiella carbonate platform margin, Abruzzi, Italy. – Memorie di Scienze Geologiche, 51, 333-413 Swift, D.J.P., Thorne, J.A. (1991): Sedimentation on continental margins, I: a general model for shelf sedimentation. – In: Swift, D.J.P., Oertel, G.F., Tillman, R.W., Thorne, J.A. (eds.): Shelf sand and sandstone bodies. – International Association of Sedimentologists, Special Publications, 14, 3-31 Szulczewski, M., Belka, Z., Skompski, S. (1996): The drowning of a carbonate platform: an example from the Devonian-Carboniferous of the southwestern Holy Cross Mountains, Poland. – Sedimentary Geology, 106, 21-49 Wilmsen, M. (2000): Evolution and demise of a mid-Cretaceous carbonate shelf: the Altamira Limestones (Cenomanian) of northern Cantabria (Spain). - Sedimentary Geology, 133, 195-226 Wilson, P.A., Jenkyns, H.C., Elderfield, H., Larson, R.L. (1998): The paradox of drowned carbonate platforms and the origin of Cretaceous Pacific guyots. – Nature, 192, 889-894 Further reading: K177, K190
Transects: Platform to Basin
15.7 Platform-Slope-Basin Transects Deep-marine carbonates comprise two major sedimentary groups differing in depositional criteria and diagenetic development: • Pelagic carbonates characterized by fine-grained sediments formed by pelagic planktonic and nektonic organisms (Sect. 2.4.5.3), and • Resedimented allochthonous carbonates (Sect. 2.4.5.4), whose constituents are derived from platform and slope settings and exported to slopes and into basins. Microfacies criteria of the first group are discussed in Sect. 15.8. This section deals with the second group, specifically with carbonates produced by gravity mass transport and consisting of resediments. The first part gives a summary of the origin, the composition and the diagnostic criteria of carbonate slopes. The second part describes case studies and demonstrates the potential of microfacies analyses for understanding slope history and slope geometries. Significance: Recent research on fossil and recent submarine slope systems has provided new insights into the relationships between sediment composition and slope morphology. Declivity and stabilization stages are reflected in the depositional texture, grain types, and biota, and revealed by microfacies features, including compositional data, diagenetic criteria and paleontological proxies. Sea-level fluctuations leading to different kinds of carbonate shedding are reflected by grain compositional logs that can be successfully used in correlating platform-, slope- and basin sediments. Carbonate slopes pass seawards from shallow-water to deeper-water environments. Their development is controlled by extrinsic and intrinsic factors. Extrinsic factors are sea level, tectonics, oceanography, siliciclastic input and climate. Intrinsic factors are biological productivity, sediment transport and erosion, biological stabilization and cementation, the growth of mounds and reefs, and bioerosion and dissolution. A concise synthesis of these controls and an overview of the large-scale patterns of carbonate slopes was given by Coniglio and Dix (1992). Modern slopes: Carbonate slopes are deep-water depositional environments. They exhibit a variety of depositional angles and sediment types that are different for erosional slopes, bypass slopes, and accretionary or depositional slopes. Modern carbonate slope models have been intensively studied in the Bahamas
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and Florida (Mullins and Neumann 1979; Mullins et al. 1980, 1984, 1986; Mullins and Cook 1986) and off Jamaica (Land and Moore 1977), Belize (James and Ginsburg 1979) and Grand Cayman (Rigby and Roberts 1976). The accretionary Bahamian slopes are differentiated into an upper gullied slopes and lower nongullied slopes or a basin-margin rise that passes laterally to the flat ocean floor. The sediment of the upper slope consists of periplatform oozes composed of platform-derived material and pelagic debris. Cementation and hardground formation takes place in the shallow part of the upper slope. The oozes are strongly bioturbated in the deeper part. Differences in the composition and the dominant mineralogy of periplatform ooze layers are explained by changes in the submergence and exposure of the platform, and by differences in the dissolution of aragonitic grains. Lower slope deposits are gravity deposits, interbedded with periplatform ooze. Mud- and grain-supported debris flows and turbidites contribute to the formation of aprons near the base of the slopes. Particular elements of the lower slope are deep-water reefs (lithoherms) with ahermatypic corals (see Sect. 2.4.4.3).
15.7.1 Types and Composition of Carbonate Slopes Slope types and facies models of slopes have been summarized by (1) Cook et al. (1972), with regard to the sediment wedge underlying the slopes and forming aprons adjacent to the shallow-water platform; (2) James and Mountjoy (1983) and McIlreath and James (1984), distinguishing bypass and depositional slopes based on the nature of the platform margin and the relief between platform and basin; (3) Mullins and Cook (1986), emphasizing similarities and differences between siliciclastic fans and carbonate aprons, and (4) Coniglio and Dix (1992). Commonly used classifications are those by McIlreath and James, and Coniglio and Dix. The models described by the latter are based primarily upon the geometry of shelf deposits and the adjacent platform margin, but also consider the essential factors listed below: • The intensity and location of the carbonate production on the platform. • The type of the carbonate production and the morphology of the platform margin. • The feeding areas of the exported carbonates (line source and a multitude of feeding channels). • The line source of platform-derived sediments originates from numerous channels dissecting oolitic and bioclastic shoals, reefs and islands along margins of
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rimmed platforms (depositional slope). Downslope sediment transport occurs predominantly via nonchannelized sheet flow. Carbonate slope sediments extend without interruption from the basin along gentle gradients (< 4°) up to a shallow platform margin. • The existence of erosional slopes. Steep to vertical escarpment walls can be formed during sea-level lowstands and by tectonics. The upper part of these walls may be settled by luxuriant reef-building communities. The lower part of the cemented rock can exhibit ledges and ridges acting as a base for sediment deposition and the formation of microbially induced micrite crusts that can also be recognized in ancient slopes (Brachert and Dullo 1994). • Incised parallel gullies, oriented perpendicular to the platform margin, and formed by erosion associated with downslope gravity flows. These gullies are a line source of sediment supply from the platform and upper slope to the lower slope. • The angle between the platform margin and the slope controlling the facies belts parallel to the platform margin. Slope angles of modern slopes may be low (5°– 15°), but are often relatively high (<40°) and may even be considerably steeper. • Changes in dip from upper to lower slopes (steepened in mid-slope or lower slope settings).
Platform Margin Types
Platform-margin to slope transitions and facies patterns of carbonate slopes: Fig. 15.17 shows stratigraphic profiles describing three generalized types of buildup and geometry in Phanerozoic platform margins. Differences in the size, position and composition of these margins are important for understanding depositional slopes because the sediments produced on the platform and at the platform edge are the main source for carbonate slopes. Types I and III correspond to rimmed carbonate platforms, type II to a homoclinal ramp. Note that Fig. 15.17 documents only the uppermost parts of slopes adjacent to the platform-margins. The composition of slope sediments: Microfacies studies are essential for understanding the composition of slope sediments, differentiating allochthonous and autochthonous material, and recognizing sediments formed by specific depositional processes (e.g. turbidites, debrites, slope breccias). Carbonate slope sediments consist of allochthonous and autochthonous sediments and include: • Allochthonous shallow-marine platform material composed of mud- and sand-sized particles (disintegrated calcareous algae, various skeletal grains, bioerosional chips, peloids, composite grains), and sand-,
Fig. 15.17. Basic types of platform margins after Wilson (1974) distinguished by the stratigraphic profile, shape and position of sand bars or different kinds of reefs along the profile slope, and diagnostic facies patterns. This classification calls particular attention to the geometry of rock bodies and is a useful approach in hydrocarbon and mineral exploration. A: Type I (downslope mud accumulations) is characterized by linear belts of downslope lime mud accumulations forming foreslope mud mounds below the tidal wave base. Crestal deposits are sand bars or sand islands. B: Type II ((knoll reef ramps) includes linear ramps or non-rimmed platforms with patch reefs (knoll reefs), commonly formed on gentle slopes at the outer edges of shelf margins. Bioclastic interreef material is much greater in volume than the patches of in-situ framework. Crestal deposits are sand shoals, bars or mud flats behind the wide, gently sloping platform. C: Type III (frame-built reef rims) is a platform margin created by an organic reef rim of resistant framework built into the active wave base and forming the topographic crest of the profile. Types I and II appear to have been more common in the geological record than Type III.
Allochthonous Carbonates
gravel- and boulder-sized lithoclasts derived from lithified platform or platform-margin carbonates. Submergence of the platform may result in an increased contribution of platform-margin material to slope deposits, leading to a higher input of lithoclasts. • Slope material: Boulders, blocks and giant carbonate bodies deposited on the slope or at the toe-of-slope are eroded from platform margins or from higher parts of the slope. Sand- to boulder-sized carbonate lithoclasts may also have their source in eroded rocks from higher, lithified slope parts. • Fine-grained shallow-marine material intermixed with pelagic carbonate grains forms periplatform carbonates that can accumulate as periplatform sands at the foot of steep platforms, as fine-grained periplatform turbidites interbedded with pelagic oozes in basins, or as soft periplatform ooze and mud deposited in deep basins. The emergence of platforms shut down the source area of platform material, and more pelagic sediment is deposited than peripelagic ooze. • Hemipelagic terrigenous material, usually claysized particles, is transported across the shelf and deposited on slopes or in the basin. The material is intermixed with carbonate sediment, forming a mixed siliciclastic-carbonate facies or occurs in discrete clayrich marly and argillaceous beds sandwiched between fine-grained limestone beds. • Autochthonous carbonate on slopes is produced by organisms able to induce carbonate precipitation (e.g. microbes) or secrete calcareous skeletons (various invertebrates including foraminifera, sponges, corals, mollusks, serpulids). Some of these organisms contribute to the formation of extended reef structures, including mud mounds, reef mounds and coral reefs.
15.7.2 Allochthonous Slope and Basin Deposits: Diagnostic Criteria Depositional processes and products: Sediment is transported to and accumulated in slope and basin settings by • Settling of fine-grained platform-derived material from suspension (often forming black, very fine crystalline and often laminated, rhythmically bedded limestones and shales). • Subaqueous rockfall and talus (e.g. below steep escarpments). Carbonate rockfall sediments are characterized by clast- or mud-supported, poorly sorted sediments with angular clasts of different size. Rockfall produces forereef breccia, escarpment breccias, and various types of slope breccias.
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• Sediment gravity flows (mass flows) characterized by a mixture of sediment and fluids comprise turbidites, grain and debris flow deposits, breccias, and conglomeratic deposits. • Submarine sliding (movement of a rigid, internally undeformed mass along a discrete shear surface) is triggered by slope failures. • Downslope creeping (slumping) is a characteristic long-term process within non- or semilithified sediment piles, caused by bedding-parallel translocation along well-defined surfaces or a diffuse shear zone. The sediment becomes simultaneously deformed and may develop distorted bedding and synsedimentary folds (e.g. slump folds). The availability and volume of detrital material transported by gravity flows from neritic shelf environments to basins depend on an interplay between the production of biogenic material, climate, sea-level changes and tectonics. In carbonate systems the sea-level related timing of shedding from platforms and climate is the essential control (Droxler and Schlager 1985; Sarg 1988; Schlager et al. 1994). Climate can favor increased volumes of specific shelf carbonates (e.g. foramol/ heterozoan carbonates during relatively cold phases) thus providing more source material for turbidites. One example is the Early Tertiary Zumaya series of the Gulf of Biscay, as shown by Gawenda et al. (1999).
15.7.2.1 Submarine Rockfalls Rockfalls are characterized by free falling, and often large rock fragments that are transported across steep and moderate angles over short distances. The material is eroded from cliffs, escarpments or rocky shores (see Sect. 5.3.3.3). Important criteria are: • Depositional units characterized by distinct boundaries. • Clasts are angular to subangular, uniform or polymict, and usually very closely packed. • Sorting is poor to very poor. No grading. • The size of the clasts varies between a few millimeters to tens of meters. Large clasts occur together with very small clasts. • The fabric is usually clast-supported (e.g. at the base of oversteepened platform rims). • The matrix consists of poorly sorted rock debris, fine-grained carbonate, or argillaceous material. Microfacies: The clasts may exhibit different lithologies, different microfacies types and rocks of different
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stratigraphic age depending on the composition of the eroded rock sequences. Clasts occur together with isolated fossils (e.g. corals or rudists transported from reefs), or are mixed with exotic material brought in by rivers from the hinterland. Ancient carbonate rock fall deposits are known from near-coast and transitional environments (e.g. Thiedig 1975), rocky coasts, seamount cliffs (Seyfried 1980), and steep forereef slopes. 15.7.2.2 Breccias and Megabreccias Sedimentary breccias are common constituents of carbonate slopes (Sect. 5.3.3). The breccias are characterized by • discontinuous lenses (meters to tens of meters) with sharp and undulated boundaries at the top and the bottom, • polygenic compositions, yielding shallow-water and deep-water elements, and consisting of bioclasts and lithoclasts, • variously shaped and usually poorly sorted clasts, • clasts floating within a fine-grained matrix (composed partly of comminuted breccia material) or exhibiting fitted contacts, • clasts sometimes showing a preferred orientation of longer clasts. Megabreccias are matrix-supported breccias consisting of well-defined blocks, from meters to more than several hundred meters in size. Seismic shocks and gravity collapses of the outer parts of high-angle outer carbonate platforms may contribute to the development of huge breccia masses deposited in upper and midslope settings. Scalloped embayments and large-scale erosion features depicted in seismic reflection profiles record the catastrophic removal of many cubic kilometers of sedimentary units and explain the presence of major megabreccia sheets at the toes of slopes (Mullins et al. 1986; Mullins and Hine 1989; Hine et al. 1992; Stewarts et al. 1993; Sano and Tamada 1994). An excellent overview of the importance of megabreccias in the context of sequence stratigraphy was provided by Spence and Tucker (1997). Common features of megabreccias are the • close association of breccias with various types of mass-flow deposits, • occurrence of megablocks consisting of reworked limestone breccias together with breccias of variable size, embedded within fine-grained matrix (see Sect. 15.7.3.1 and Sect. 15.7.3.2). • mixture of intrabasinal and extrabasinal blocks in size up to several hundreds of meters, • disorientation of geopetal fabrics within individual
Megabreccias
blocks, indicating alternations of transport and resting phases, • rounding of smaller and larger blocks caused by dissolution and biogenic encrustation, • occurrence of cryptic, cavity-dwelling organisms and specific low-growing encrusters, • joint occurrence of blocks exhibiting different diagenetic histories (e.g. blocks derived from subaerially exposed platform parts and blocks eroded in deep-water environments). The development of breccias often starts with submarine rockfall deposits and avalanches, followed by the formation of megabreccias, and finally the deposition of turbidites. 15.7.2.3 Debris-Flow Deposits Debris flows composed of clasts and supported and carried by a mud-water mixture lead to the deposition of sediments that have variously been called debrites, debris sheets or mass breccia flows (Hiscott and James 1985). Many so-called debrites have a granular matrix that is noncohesive. In these cases, transport by turbulence and grain interactions is more likely. Debris flows are capable of traveling over very gentle slopes and may represent a pre-phase of turbidite sedimentation. Debrites range in thickness from a few decimeters to several tens of meters. Most debrites form sheetlike and lenticular bodies with conformable, sometimes also erosional contacts with the underlying fine-grained sediment. Upper contacts are sharp or the debrite bed passes upward into turbidites (Krause and Oldershaw 1979). Many debrites are coarse-grained breccias or conglomerates characterized by poor sorting, lack of stratification and often random or chaotic clast fabrics. Clast packing is variable and shows matrix-support as well as clast-support. The source area of clasts can only be inferred by comparing clast microfacies with microfacies of various parts of platform and slope carbonates (Nebelsick et al. 2001). The common criteria of debrites in microfacies studies are: • Sedimentary units are generally massive or coarsebedded, often with significantly thick beds. Irregular top surface. • Large clasts may project above from the bed. • No preferred depositional fabric except for crude grading in the basal part. • Limestone clasts are very poorly sorted and of variable size, usually of sand-grade or finer size. • Angular or rounded clasts, or of mixtures of both types occur.
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Debris- and Grain-Flow Deposits
• Clasts are supported by fine-grained sediment consisting of micrite, calcisiltite of argillaceous material. • Larger boulders or pebbles may be embedded within debris flow deposits. A specific type of debris flow deposits are disorganized limestone conglomerates, consisting of shallowand/or deep-water clasts forming extended sheets and channels. Common criteria are: • Sand- to boulder sized clasts; very poor sorting. • Angular, subrounded or rounded clasts. • Randomly distributed, often densely packed clasts. • Monomict or polymict clasts, possibly derived from shallow-marine, slope or deep-marine sediments. • Lime mud or argillaceous mud matrix. • Internal structures within the beds rare or absent. • Normal grading absent; inverse grading in the basal zone of the unit. Examples of disorganized limestone conglomerates and breccias have been described from the Cambrian (Keith and Friedman 1977; McIlreath 1977; Cook and Taylor 1977), Devonian (Cook et al. 1972), Cretaceous (Cook and Enos 1977), Pleistocene (Crevello and Schlager 1980).
15.7.2.4 Grain-Flow Deposits Grain flows are gravity flows in which material of different grain sizes is supported within the flow, mainly by the strength of a fluid matrix consisting of water and clay minerals. Identification of grain flow deposits is difficult because most of the features are also produced by other transport processes. Very often grain flow deposits are associated with turbidites (e.g. Stauffer 1967; Eberli 1987). Common criteria are: • Thick massive beds. • Deposits draping over irregularities and thinning up flanks of channels. • Flat tops and flat base. Sometimes flute marks. • Scours and injection structures at the base. • Partial to complete clast support. • Clasts floating in sandy or muddy matrix. • Poor sorting. • Grain orientation parallel to flow; faint, dish-shaped laminae. • No parallel lamination or cross lamination. • Normal grading absent or rare. Sometimes reverse grading near the base.
Calcareous debris flow sediments are predominantly deposited on deep-marine slopes or at the base of slopes, in contrast to turbidites that are deposited in proximal or distal positions depending on the position of the source areas. Grain flow deposits may pass rapidly into pelagic and hemipelagic sediments.
15.7.2.5 Turbidites Turbidites are sedimentary deposits laid down by turbidity currents and intercalated in fine-grained sediments (Edwards 1993). Turbidite successions vary greatly in sedimentary structures, bed thickness, and textural features. Variations are controlled by distance between the source and of the depositional area, point or line source, composition of the sediment available, topography of the depositional area and density of turbidity currents (low or high density suspensions). Common criteria of turbidites are: • Regular vertical sequence of units characterized by specific sedimentary structures (Bouma Sequence respectively Meischner Sequence). • Graded bedding and lamination common. • Bottom surface sharply defined, frequently with sole marks. Upper bed surface usually not well defined, showing transitions into the overlying beds. • Biota consisting of allochthonous, redeposited fossils derived from shallow-water and slope environments. Autochthonous fossils rare, apart from trace fossils. Pelagic and deep-water benthic fossils occur in intercalated shales and may form a small part of the fossils occurring within the turbidites. Bouma Sequence: An ideal turbidite sequence consists of a vertical succession of internal sedimentary structures. The Bouma Sequence (Bouma 1962), first studied in siliciclastic deposits, exhibits five intervals, from base to top: • graded or massive unit (division) A, • lower parallel laminated unit B consisting of thin laminae, • unit C characterized by current ripple lamination with several sets of unidirectional foreset bedding and/or convolute bedding, • upper parallel laminated unit D consisting of very thin, often obscure laminae within fine-grained sediment, • very fine-grained unit E without primary sedimentary structures. As a rule, complete Bouma Sequences should characterize proximal turbidites, incomplete Bouma Sequences distal turbidites (Walker 1967).
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Limestone turbidites, allodapic limestones and the Meischner Sequence Carbonate and siliciclastic turbidites: Limestone turbidites differ from siliciclastic turbidites in • the size of the bioclastic particles; it is predominantly controlled by ecological constraints in the source area and by taphonomic criteria, and not by the range of grain transport as for siliciclastic grains; • the abundance of lithified particles. Platform and shelf carbonates are rapidly cemented. Lithoclasts are therefore much more common than in siliciclastic turbidites; • the variability of grains contributing to limestone turbidites; it is much higher than the rather uniform composition of siliciclastic turbidites. Transport and settling of skeletal grains is influenced by differences in size, shape, microstructure, porosity and density. Hydraulic sorting is common. • Unlike siliciclastic turbidites, complete Bouma Sequences are rare in carbonate deposits. Some of these differences may be caused by the relatively weak thixotropy of calcareous muds as compared to clay muds. Criteria of limestone turbidites: Carbonate turbidites occur in deep-marine and shallow-marine, but also in lacustrine settings. Sequences containing turbidites are interpreted as fan deposits, or described by apron models. Common criteria are: • Thick or thin allochthonous limestone beds intercalated within micritic carbonates, marls and argillaceous sediments. • Graded calcarenites or calcisiltites; calcirudites. • Lower bed boundary sharp, but sometimes with bottom marks (groove and flute casts, erosional marks). • Gradual transition of top surfaces into overlying beds. • Vertical sequence consisting of sedimentary units (zones) characterized by specific sedimentary structures (Meischner Sequence). • Grading and lamination common. • Cross-bedding and convolute lamination relatively rare. • Internal size grading. • Sizes of lithoclasts range from silt to boulders, sizes of skeletal grains between coarse sand and silt. • Grain size decreases parallel to the decrease of bed thickness. • Sorting in coarse parts good, in laminated parts moderate to good. • Particles are skeletal grains, peloids and ooids as well as lithoclasts and sometimes also extraclasts.
Limestone Turbidites
• Lithoclasts are often concentrated at the base, bioclasts in higher parts of the turbidite bed. • Skeletal grains include benthic organisms derived from shallow-marine and slope environments, fragments of reef organisms, and some pelagic fossils. • Secondary silicification is common. Late diagenetic silification occurs along sedimentary structures due to the high porosity. Pore waters are enriched in silica because of the rapid burial of the sediment. Silicification is often preceded by calcite cementation (Hesse 1987). Source areas of limestone turbidites are shallowwater environments (platforms, platform-margins, banks, reefs) and slope environments. Limestone turbidite beds can be followed over several hundred meters to several kilometers. Bed thickness ranges between < 1 cm and several meters. Amalgamation of thick beds is common. Allodapic limestones: This term, suggested by Meischner (1964), denotes limestone beds consisting of graded carbonate debris, transported and redeposited by turbidity currents. The term ‘allodapic’ refers to detrital material ‘which originated elsewhere’. Originally coined for limestone turbidites yielding grains derived from platform rims or reefs, the term is now used more or less as a synonym for carbonate turbidites. • Internal sequence: The Meischner Sequence (Fig. 15.18) summarizes the criteria of an ideal allodapic bed. The composition and sequence of allodapic beds is determined by the amount of transported material, the distance from the source area, the rate of accumulation, the intensity of background sedimentation. • Geometry: Each bed forms a lenticular body. The coarser basal parts reach their maximum thickness upstream. The fine-grained upper parts have their maxima further down stream. • Fossils: The coarse-grained parts of allodapic limestones usually contain benthic biota, both macrofossils and microfossils. Mixing of ecologically different types may occur. Microfossils can exhibit hydraulic sorting resulting in a fractionation of foraminifera depending on size, density and settling potential (Herbig and Mamet 1994). Planktonic fossils occurring in fine-grained parts are sometimes strikingly well preserved. • Diagenetic overprint, e.g. compaction, can destroy primary depositional boundaries. Care must be taken in differentiating depositional lamination and lamination caused by stylolitization. Pressure solution
Meischner Sequence
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can destroy primary sedimentary structures on bed tops and bottoms and obliterate internal structures e.g. in turbidites. ‘Bedding’ can be created by diagenetic remixing of carbonates (Eder 1971). Meischner Model and Bouma Model: Zone 1a and Zone 1b of the Meischner Sequence correspond to Division A of the Bouma Sequence. Zone 1c exhibits criteria of Division B. Zone 2a and Zone 2b correspond to Division C. Zone 3 has criteria of Division D of the Bouma Sequence. The uppermost pelitic Division E of turbiditic and hemipelagic origin is not included within the Meischner Sequence. The models are similar, but exhibit some differences. In comparison with the Bouma Sequence, Zone 1 of the Meischner sequence is characterized by ruditic rather than arenitic allochthonous grains. Reverse grading near the base of Zone 1 seems to be more common than in Division A of the Bouma Sequence. The differentiation of lower planeparallel laminated (Division B), middle ripple or convolute laminated (Division C), and upper parallel laminated parts (Division D) of the Bouma Sequence is much less developed in Zone 2 of limestone turbidites. Differentiation of turbidites: Turbidites are differentiated into several depositional types: • Proximal turbidites, deposited relatively close to the source area are massive, relatively weakly graded, and exhibit only poorly developed tractional structures and little interbedded pelagic sediment. • Turbidites characterized by the Bouma Sequence, have distinct graded bedding, oriented erosion and fill markings at the base, interbedded pelagic sediment and a characteristic succession of sedimentary structures. • Distal turbidites, formed far away from the source region, are characterized by thin, fine-grained graded layers, well-developed cross-lamination and absence of massive intervals and parallel laminations. Some of the more important characteristics for the proximality and distality of limestone turbidites are summarized in Tab. 15.2. Ideally, proximity indicators work for turbidite sequences deposited only from a single source by longitudinal sheet flows. In turbidite fans, turbidites with proximal and distal features may be juxtaposed. Meandering of turbidity currents may produce strong deviations from general proximal-distal patterns. Some of the deviations are caused by differences in the sedimentation in deep and large versus shallow and small basins, point or line sources, and in the different settling behaviors of carbonate particles (Sarnthein and Bartolini 1973).
Fig. 15.18. The Meischner Sequence describing an ideal ‘allodapic’ limestone turbidite. The turbidite bed is intercalated between beds formed by the basinal background sediment (fine-grained ‘pelites’, i.e. marls or micrites containing nektonic and planktonic fossils). The ‘pre-phase’ (zone 0) is characterized by micrites or marls with scattered small lithoclasts and rather large rock fragments. The genesis of the pre-phase is controversial. Solution and reprecipitation of carbonate may be important. The contact to the underlying pelagic zone is sharp. The upper surface of the pre-phase may be wavy and irregular. The ‘main phase’ (zone 1) exhibits three parts: Part 1a is a distinctly graded limestone with shallow-water fossils and lithoclasts. Reverse grading and grain imbrication may occur. Sorting increases upward. Part 1b is a fine-grained micrite. The upper part 1c may be faintly laminated and includes angular limestone clasts and micrite pebbles. Zone 2 is a micrite characterized by planar bedding planes with densely spaced laminations (2a), overlain by a unit with current ripple lamination (2b), and sometimes also with convolute bedding. Zone 3 is represented by marls that may exhibit flaser textures. This zone merges gradually into the overlying marly or pelagic sediment. Bed thickness is about 1 m. Bed thickness is high in the near-source proximal position, and decreases in distal depositional areas. After Meischner (1964).
Fluxoturbidites: Turbidites can be underlain by beds characterized by large, very poorly sorted and inversely graded components forming breccias. These fluxoturbidites are overlain by normally graded turbidites. Fluxoturbidites contribute to a traction carpet, signifying the begin of allochthonous sedimentation in proximal depositional sites on slope or base-of-slope environments (Stanley and Unrug 1972; Sallenger 1979; Steiger 1981). Deep-water source turbidites: Limestone turbidites with grains derived not from shallow-marine environments but rather from upper slope or deep-water buildups exhibit features known in allodapic limestones, but also show significant differences when compared with the Meischner or Bouma Model. Differences involve the sequence of internal structures and the greater varia-
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Proximal and Distal Turbidites
Tab. 15.2. Criteria of proximal and distal limestone turbidites. The mean maximum grain size (decreasing in a downstream direction), the ABC-Index (percent of beds beginning with Division A plus half the percent of beds beginning with Division B), and the ratio of the thickness of Division A to the mean maximum grain size may also be used. Grading curves and proximality curves exhibiting the differences in maximum grain sizes at different vertical positions of the turbidite bed have also proved useful (Engel 1974; Steiger 1981). Proximal Field Aspects Distribution of turbidite beds
Distal
Densely spaced; isolated or amalgamated; background sedimentation limited Ratio of turbidite and pelagic sediment High Geometry of turbidites Lenticular beds of varying thickness; medium- to thick-bedded Associated sedimentary structures Base of turbidite beds Fluxoturbidite at the base of the bed Top of turbidite beds
Internal Sequence Complete Bouma/Meischner Sequence Graded part (Zone 1a) Grading Inverse grading Coarse-tail grading Base of basal detrital parts Top of basal detrital parts Lamination (Zone 1c and 2a) Ripple and convolute lamination Micritic upper parts of turbidite beds Lowermost units of turbidite beds Grain size Average grain size Significant vertical breaks in dominant grain sizes
Scattered, isolated; background sedimentation high Low Regular, planar parallel beds; thin-bedded (thinner than a few tens of cms) Slumping structures; breccias; conglomerates Breccias rare Sharp; sometimes scours, tool marks,washSharp outs, reworked pebbles, sometimes erosive Abundant Absent or rare Gradual Sharp and gradual Relatively common
Rare
Well-developed, thick Graded, poorly graded or not graded Graded zone thick Common Common Generally sharp Often sharp Less common, often restricted to thicker beds Restricted to thicker beds Bedding (Zone 2b) Thin or absent Zone 1a (Division A)
Decrease in thickness Grading common
Rudite and arenite Common
Fossils Microfacies of clasts
Well-developed Zone 1b or Zone 1c
Micrite Low
Lithoclasts, extraclasts, fossils, ooids, peloids Benthic shallow-water fossils derived from platforms and platform margins; slopederived fossils, rare pelagic fossils Variable to highly variable
tion in ecological types of fossils (Tucker 1969; Remane 1970; Steiger 1981). Crinoidal turbidites are characterized by the dominance of horizontal lamination, scarceness of the graded parts, the lack of the uppermost fine-grained unit, and a sharp contact with the shales above. Deep-water calciturbidites can originate from the transport of pelagic sediment, primarily deposited on the tops or flanks of a sea-floor relief, to
Less common to absent
Arenite and smaller Absent
Matrix Matrix between grains in detrital Sparry calcite parts of the turbidite Ratio of turbidite and pelagic sediment High Microfacies Clasts
Absent Absent Sharp Grades into finer sediment More common
Lithoclasts, fossils, ooids, peloids Benthic platform- and slope-derived fossils; autochthonous deep-water fossils Relatively uniform
intrabasinal depressions. A well-described example are the Late Jurassic Saccocoma calciturbidites from Poland (Matyszkiewicz 1996). Practical advice for studying limestone turbidites The investigation of allodapic limestones requires that field observations and data extracted from thin sections and fossils be well integrated. The following text
Limestone Turbidites
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Fig. 15.19. Turbidite or tempestite? Turbidites and storm-induced tempestites may exhibit similar structures (see Fig. 15.20).This sample comes from limestone beds deposited in a base-of-slope position in a shallow inter-reef environment close to the margins of coral reefs (Flügel 1975). The stratigraphic section consists of bedded micritic limestones alternating with abundant intercalations of bio- and lithoclastic beds exhibiting coarse- and fine-grained skeletal grains. The thickness of the allochthonous beds varies between a few millimeters up to 30 centimeters. Skeletal grains were derived from shallow-water environments (dasyclad algae, various porostromate algae; lituolid foraminifera). The source of coral debris and shells of bivalves, gastropods, brachiopods and echinoderms were small patch reefs. The frequency of skeletal grains varies within the allochthonous beds. Some beds yield up to 70% porostromate calcimicrobes, some beds 40% bivalve shells and about 25% micritic intraclasts, peloids and oncoids, and other beds are predominantly composed of oncoids, peloids and intraclasts. The sequence shown in the figure starts with an erosional surface at the base separating the underlying micrite from the bioclastic grainstone. The grainstone unit, consisting of bioclasts (predominantly calcimicrobes, foraminifera, mollusk shells) peloids and some lithoclasts, is graded. Grading is reflected by the upward decrease in grain size. The upper surface of the graded unit is sharp. The overlying sediment consists of packstone in the lower part, grading into micritic and fine-peloidal mudstone in the upper part. Lamination is indicated by thin dark laminae and slightly thicker light laminae. Skeletal grains in the lower part are disoriented. The tiny shell fragments occurring in the upper part of the sample have a parallel orientation. The sample is a turbidite, as indicated by the erosive base of the graded unit (corresponding to the boundary between Zone 0 and Zone 1a of the Meischner Sequence), the normal grading (corresponding to the transition from Zone 1a to 1b), the differentiated lamination of the upper part (corresponding to Zone 1c and Zone 2), and the difference in the background community that indicate the low-energy basinal depositional environment. Late Jurassic (Early Tithonian): Kapfelberg near Kelheim, Bavaria, Germany.
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Limestone Turbidites: Practical Advice
Top
Sedimentary
Turbidites
Tempestites
Wave ripples and ripple cross lamination
Absent
Common in proximal parts
Current ripples and current ripple bedding
Common
Less common than in turbidites Rare
structures
Convolute lamination
Common
in beds
Hummocky cross stratification
Absent
Common
Traction carpet with inverse grading
Common in proximal parts
Absent
Sole marks
Uni-directional
Often bipolar, irregular scouring, gutter cast channeling
Benthic background community
Deep-water facies
Shallow-water facies
Displaced fossils within event beds
Shallow- and deep-water species
Shallow-water species only
Autochthonous post-event fauna and bioturbation
Episodic colonization by specific fauna preceding return to background conditions
Fauna similar to pre-event fauna if similar substrata are available
Amalgamation
Less common
Very common and pronounced
Stratigraphic
Lateral continuity of single beds
Often over long distances
Usually limited
context
Thickness of sequence
Often large, commonly but not always associated with deep-water facies
Limited, associated with shallow-water facies
Bottom
Biofacies
Fig. 15.20. Turbidite and tempestite criteria. High-energy events can produce similar or identical textures and sedimentary structures. The vertical sequence of turbidites as well as that of storm deposits (see Sect. 12.1.2.1) is characterized graded and laminated structures. Differentiating turbidite and tempestite beds requires consideration of various aspects of sedimentary structures, biofacies and the stratigraphic context. See Fig. 15.19 for an example.
may act as a guide for sampling and studying limestone turbidites. Box 15.3 lists papers containing methodological advice on the advantages and disadvantages of using calciturbidites as facies indicators. How to investigate limestone turbidites? Use criteria displaying diagnostic features on different scales. Note similarities with tempestites (Fig. 15.19, Fig. 15.20). Sampling Sampling of turbidite beds must consider the great variation in turbidite sequences. Densely spaced samples should be taken from the base and the top of the bed, and within the depositional units of the sequence. Equal distance sampling within the units is recommended. Many calciturbidites are stylolitized. Take samples from individual stylobeds for recognizing significant changes in grain sizes. Additional samples must be taken from the background sediment. Try to investigate the total thickness of turbidite beds. This is easily done because many beds are only a few tens of centimeters thick.
Field aspects • Determine the geometry, lateral extension and thickness of the turbidite beds. • Evaluate the frequency of turbidite beds within a given section. Measure the distance between turbidite beds. Do turbidites prevail over beds formed by normal pelagic sedimentation (common if the turbidites have been formed in a position proximal to the source area) or are turbidite beds limited to only a few beds within a sequence consisting predominantly of basinal sediments (common in distal position)? • Look for indications of pressure solution that may strongly alter appearance of bedding (see Sect. 7.5.2). Bouma/Meischner Sequence units • Try to recognize the Zones 1 to 3 of the Meischner Sequence or the divisions A to E of the Bouma Sequence. Which zones are missing? Have a look at the boundaries between the zones, which may be sharp, gradual or faint. • Describe the units by means of texture considering the size of the dominant transported grains. Differentiate between textures made by silt-sized and sand-sized particles, or fine- and coarse-sized particles.
Limestone Turbidites: Practical Advice
• Measure the thickness of the units and compare the thickness with the thickness of the layers representing the background sedimentation. Microfacies • Use vertically and horizontally oriented thin sections. Vertical sections exhibit the composition and boundaries of individual internal units. Sections parallel to bedding planes allow sediment transport to be recognized. • Determine the type and frequency of skeletal grains, lithoclasts, extraclasts and other particles. Frequency should be determined by point counting (Sect. 6.2.1.1). Differentiate between particles that were already lithified at the time of erosion and redeposition, and particles that were not lithified. • Note the shape and roundness of litho- and extraclasts. Distinguish lithoclasts of different texture and microfacies. • Look for orientation patterns of particles indicating type of grain settling and current transport. Grain size and packing • Perform a thorough grain-size analysis. The maximum grain size of lithoclasts and bioclasts reflects the degree of transport energy. At least 50 grains should be averaged. Relate maximum grain size and grain size ranges with bed thickness and the thickness of sedimentary units within the turbidite bed. Consider differences in porosity and the density of grains that could influence settling behavior. Calculate grain-size parameters (Sect. 6.1), particularly the mean, standard deviation and skewness (see Table 6.1). Parameter diagrams indicate the type and strength of transport energy. Sect. 6.1.2.2 gives an example using Barmstein limestone samples. • Use packing indices (Sect. 7.5.1), e.g. the number of grain cross sections per square centimeter. A high grain number indicates the deposition of many small grains; low grain numbers indicate the predominance of coarse grains. Biota • Differentiate between benthic and planktonic fossils. Benthic fossils should be determined by their assignment to groups or taxonomic units. The occurrence of dasyclad green algae in calciturbidites should indicate a source area in very shallow, well-lighted, platform interior environments. Specific foraminifera can be used to distinguish source areas within different parts of reef complexes (see Sect. 14.2.2 and Pl. 111). Consider the assemblages of microfossils that may reflect the major source environments, e.g. platform interior,
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high- and low-energy reefs, or upper and lower parts of slopes (see Sect. 15.7.5). • Consider the preservation of fossil shells. Strongly bored and worn shells may indicate a source area with reduced sedimentation and high nutrient input (e.g. lagoons). • Look for burrowing at the top of the turbidite or within individual sedimentary units. The latter indicate an arrested deposition of turbidite material. 15.7.2.6 Sliding and Slumping A slide is the movement of a rigid, internally undeformed mass. A slump originates from mass sliding and the creeping of semi-consolidated sediment. The detachment surface of slides is called an intraformational truncation surface. The study of detached submarine slide masses aids in recognizing slope failures. Slide masses can be preserved either downslope or within the adjacent basins. The reasons for sliding are varied and interrelated (Vortisch 1964; Bernoulli and Jenkyns 1970; Schwarz 1975; Nardin 1979; Mills 1983; Eberli 1988; Coniglio and Dix 1992). Sediment overloading and seismic activity as well as excessive pore pressures are the most obvious reasons. Evidence of slide masses Box 15.3. Selected references on ancient carbonate turbidites. Many of these papers include descriptions not only of turbidites but also of other gravity flow deposits. Tertiary: Betzler et al. 2000; Engel 1974; Gawenda et al. 1999; Reijmer et al. 1992; Skaberne 1989; Westphal 1998. Cretaceous: Carrasco and Baldomero 1977; Cazzola and Soudt 1993; Colacicchi and Baldanza 1986; Everts 1991, 1994; Ferry 1979; Harloff 1989; Hesse and Butt 1976; Mutti et al. 1978; Sagri 1979. Jurassic: Bosellini 1967; Carozzi 1955; Eberli 1987, 1988; Ebli 1997; H.W. Flügel and Fenninger 1966; H.W. Flügel and Pölsler 1965; Garrison and Fischer 1969; Marcinowski 1970; Matyszkiewicz 1996; Misik and Sykora 1982; Mutti et al. 1978; Scheibner and Reijmer 1999; Schlager 1980; Steiger 1981. Triassic: Matzner 1986; Reijmer and Everaars 1991; Reijmer et al. 1994; Watts 1988. Permian: Brown and Loucks 1993, Flügel et al. 1991. Carboniferous: Davies 1977; Franke et al. 1975; Hemleben and Reuther 1980; Herbig and Bender 1992; Herbig and Mamet 1994; Izart et al. 1997; Meischner 1962, 1964; Yurewicz 1977 Devonian: Babek and Kalvoda 2001; Carozzi and Banaree 1984; Cook et al. 1972; Eder 1971; Junge 1992; Lütke 1976; Szulczewski 1968; Tucker 1969, 1974. Ordovician: Pohler and James 1989. Cambrian: Cook and Taylor 1977; Demicco 1985; Reinhardt 1977.
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includes hummocky surfaces created by creep lobes, initial dissection of strata, and compressional folds. Large-scale gravity slides are associated with extended megatruncation surfaces forming megabreccia sheets. Examples have been described from the Cambrian and Cambro-Ordovician (James 1981; Stewart et al. 1993), Silurian (Surlyk and Ineson 1992), Late Paleozoic (Franseen et al. 1989) and Triassic (Bosellini 1984, 1989; Heck and Speed 1987). Sliding and slumping of carbonate sediments occurs in deep-water as well as in shallow-water environments (e.g. tidal zones). Common features that can be studied on an outcrop scale are • deformed beds occurring as distinct intercalation between undeformed beds; • welded upper contact in the deformed bed, characterized by a depositional fit between the irregularities of the upper surface and the base of the overlying bed; • the deformed bed may be overlain by limestone consisting of graded or non-graded calcarenites or calcirudites; • deformation of beds as shown by folding patterns; • thickening or thinning of folded strata; • folded anticlines may be eroded at the upper surface of the deformed bed; • preferred orientation of folding axes unrelated to the tectonic strike; • deformed beds displaying a broad spectrum of structures comprising simple sliding of intact sets of layers, convolution and distortion of beds, rolled-up structures, and intensive shearing of dislocated beds. Most limestones with pronounced slumping structures are micrites. Usually no differences can be detected between the microfacies of slumping beds and the underlying beds. Some differences have been found in the insoluble residues of the limestones. Larger amounts of clay-sized particles in the non-carbonate constituents could have been responsible for a retardation of early diagenetic cementation and for changes in the thixotropic behavior of the sediment (Backhaus and Flügel 1971; Kenter and Schlager 1989).
15.7.3 Microfacies of Slope Carbonates: Case Studies The following case studies describe microfacies types of limestones formed in various parts of high-angle and low-angle slopes and in base-of-slope settings.
Permian Slope Carbonates
15.7.3.1 Permian of Sicily: Megablocks and Base-of-Slope Carbonates The case study reveals the complex history of allochthonous sedimentation (Fig. 15.21). The Late Permian of western Sicily is famous for its abundance of well-preserved fossils. The fossils are concentrated in large- and small-scale allochthonous carbonates represented by limestone breccias and calcarenites forming megablocks and megabreccias intercalated within basinal sediments. The breccias result from multi-phase debris flows and turbiditic sedimentation. Microfacies criteria and paleontological data provide evidence of long- and short-term erosion of Middle- to Late Permian platform-margin and upper-
Fig. 15.21. Facies model of Permian allochthonous slope carbonates. Factors that might have controlled the formation of the Permian Pietra di Salomone megablock in western Sicily (modified from Flügel et al. 1991). Phase A: Middle Permian shallow-marine platform with a reef zone and a periplatform talus zone. Platform- and reefderived material is deposited in periplatform position as debris flows and in parts of the pelagic basin as fine-grained calciturbidites. Phase B: Disappearance of the reef zone, possibly caused by sea-level fluctuations, and reworking of the periplatform talus, associated with input of platform sands to the basin. Phase C: Steepening of the slope and redeposition of slope breccias within depressions of the marly basin resulting in the formation of the multi-phase megabreccia now recorded by the Pietra di Salomone megablock embedded in Late Permian radiolarian-bearing argillaceous sediments. Phases A and B took place during the Wordian, phase C sometime in the Capitanian or later. This is a time range of approximately 10 ma.
781
Slope Carbonates: Permian Case Study
Fig. 15.22. Permian megabreccia and calciturbidites. A: The constituent of a megabreccia deposited at the base of a slope (see Fig. 15.21 for the complex history of the megabreccia). The poorly sorted, disorganized breccia consists of lithoclasts exhibiting different microfacies types and transported fossils. Most of the clasts are derived from platformmargin reefs, but some clasts record the erosion of slope material. Lithoclasts are packstones (PS) and boundstones (BS). The dark packstone clasts correspond to fine-grained sediment formed in an upper slope position. Boundstone clasts consist of Tubiphytes, sponges and organic crusts indicating reef facies. Smaller lithoclasts and fossils bound together by early lithification form a composite breccia (CB). Isolated fossils are represented by sclerosponges (S) associated with biogenic crusts. The arrow points to cement layers overgrowing Archaeolithoporella crusts. B: Bioclastic fine-grained grainstone consisting of dasyclad algae (EP: Epimastopora), fragments of organic crusts (white arrows) and shell fragments preserved due to thin micrite envelopes (black arrows). Minor constituents are benthic foraminifera (F). Note weak imbrication and parallel orientation of grains indicating some current transport. Distal calciturbidite. Both samples from the Late Permian of Pietra di Salomone near Palazzo Adriano, western Sicily, Italy. After Flügel et al. (1991). Scale is 2 mm for A and B..
stones and grainstones. The latter vary in predominant grain size and composition. Skeletal grains are derived from platforms and deposited in thin distal calciturbidite beds (Fig. 15.22B).
15.7.3.2 Triassic of the Southern Alps: Allochthonous Slope Sediments
slope reefs and platform carbonates. Subsequent to recurrent erosion, lithified carbonate material was transported downslope by gravity flow processes and finally deposited as fillings in depressions and channels incised in basinal sediments adjacent to the base-of-slope. The largest of the megablocks is the Pietra di Salomone block about 40 km south of Palermo. It is about 200 m long, up to 100 m wide and 30 m high. Textural types represented in the breccias (Fig. 15.22A) are, in decreasing order, lithoclastic and lithobioclastic rudstones, boundstones with sponges, Tubiphytes and biogenic crusts (e.g. Archaeolithoporella), bioclastic float-
Middle Triassic (Ladinian) and early Late Triassic (Carnian) sequences in the Dolomites are excellent examples of the export of platform and platform-margin carbonates to slope and basinal environments. The flanks show well-developed clinostratifications, with angles of 30° - 40°. These steep angles contrast significantly with the horizontal bedding of the platform interior sediments. The allochthonous material was transported by various processes, including collapse of platform margins, sliding, debris and grain flows, turbidity flows, and rockfalls. Progradation generated large-scale inclined beds (clinoforms) with depositional slope angles up to 40°. Allochthonous carbonates occur as isolated blocks (Fig. 15.23), extended megabreccias (Fig. 14.17), calciturbidites (Pl. 137) and debris flows (Pl. 115/2). Foreslope and slope facies types have been described by Brandner et al. (1991), Harris (1994) and Russo et al. (1998). Common sediment types are fine- to coarse-grained limestone breccias and rudstones, calcarenites (skeletal grainstone, lithoclast
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Triassic Slope Carbonates
Fig. 15.23. Platform-derived block within slope sediments. This block, exposed in the road cut at the Pordoi Pass, is a spectacular example of the deposition of lithified platformmargin material on carbonate slopes and in basinal sediments of the Triassic in the Dolomites. The gravity-displaced block is embedded within beds of the S. Cassiano Formation that consists of volcaniclastic sediments and limestone turbidites. Rotated geopetal fabrics within the boulders argue against it being an ‘in situ’ patch reef. The block is a boundstone with biogenic crusts, growth cavities and encrusting biota. The rounded block has a maximum height of 3 meters. Rounding of the block and of its constituent boulders is partly due to organic envelopes developed during resting phases in block transport. The platforms acting as a source of the allochthonous carbonates prograded some kilometers over the S. Cassiano Formation. Late Triassic (Carnian).
Plate 137 Microfacies of Base-of-Slope Carbonates: Triassic of the Dolomites (Italy) Excellent examples of ancient carbonate slopes are exposed in the Triassic of the Dolomites, where carbonate platforms prograded by the redeposition of shallow-water material on the slopes. Microfacies analysis reflects the source and depositional pattern of the allochthonous sediment.This plate shows a thin section (with enlarged details of it) of a single 6 cm thick limestone bed. Common criteria of carbonate slope and base-of-slope deposition and the distinct changes in platform-derived gravity flow sediment and pelagic sediment. Allochthonous sedimentation is represented by the turbiditic influx of shallow-water derived grains formed on a platform and resedimented in deep water (units A, C and D2). The pelagic background sedimentation and distal turbidites are recorded in unit B as well as in unit D1. Note the sharp boundaries between the major units. Middle Triassic (Buchenstein Formation, Ladinian): Seceda platform, Western Dolomites, Italy (–> 1–6). The basinal Buchenstein (Livinallongo) Formation consists of flat, medium- to thin-bedded marls and limestones that intergrade upslope with allochthonous limestones consisting of material shed from the margins of a platform. The carbonates of the Buchenstein Formation exhibit several microfacies types, including radiolarian packstones and wackestones, fine-laminated dark mudstones, and calciturbidites with shallow-marine and slopederived lithoclasts and fossils. The varying composition of the calciturbidites reflects platform progradation and high-frequency eustatic oscillations (Maurer et al. 2003). 1 Vertical sequence consisting of resedimented shallow-marine carbonate grains and pelagic deep-marine sediment. Unit A1: The basal lithobioclastic grainstone consists of lithoclasts, skeletal grains and peloids. Larger lithoclasts are peloid packstones (PP) as well as bindstones (B). Both microfacies are common in Middle Triassic platform interior and reef mound settings. Bioclasts are dasyclad green algae (D), foraminifera, ostracods, and rare echinoderm debris, ‘Tubiphytes’ (T) and small coated debris. Note fining-upward pattern. The unit corresponds to part of a proximal calciturbidite. Unit A2: The upper part of unit A differs from the lower part by the admixture of pelagic biota (‘filaments’, e.g. thin pelagic bivalve shells; see Pl. 113/2) and echinoderm fragments. Unit B: Radiolaria packstone. The abundant spumellarians occur with some broken filaments, a few monaxon sponge spicules (SS) and rare ostracods (O). Darker patches are caused by burrowing. Unit C: Lithobioclastic packstone. In comparison to unit A, lithoclasts (L) are smaller and better sorted. Bioclasts are represented by some mollusk shells (S) and rare echinoderms. Larger grains are strongly micritized bindstone clasts. Unit D1: This unit is characterized by a chaotic accumulation of densely packed pelagic bivalve shells (coquina floatstone). Platform clasts, including thick-shelled bivalves (B), bindstone lithoclasts and coated grains, are subordinate. Unit D2: Alternation of platform-derived material and pelagic grains. Platform-derived grains are bindstone lithoclasts and thick-shelled bivalves (B). Parallel alignment of filaments and sponge spicules point to the existence of weak bottom currents. A part of the variations in turbidite composition is related to hydrodynamic sorting. 2 Detail of part A1 (not shown in –> 1). Lithobioclastic grainstone. Grains are fragments of the enigmatic encrusting fossil ‘Tubiphytes’ (T; see Sect. 10.2.6), pelmicrite clasts and small echinoderms (top left corner). 3 Detail of part A2 (not shown in –> 1). Characteristic sections of ‘Tubiphytes’, variously sized peloids, filaments (black arrow) and a few radiolarians (white arrow). 4 Detail of part B. Radiolaria packstone. Densely packed, variously sized spumellarian radiolaria, rare ostracods (O) and monaxon sponge spicules (SS). 5 Detail of the upper part of unit C and the lower part of unit D1. Micritic platform clasts overlain by pelagic filaments. 6 Detail of the upper part of unit D1 and the lower part of unit D2. The irregular orientation pattern of the filaments indicates burrowing (BU).
Plate 137: Triassic Base-of-Slope Carbonates
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Slope Carbonates: Jurassic Case Study
Fig. 15.24. Microfacies of boulders of carbonate slope wedges (see Fig. 14.17). A: Boundstone consisting of microbial crusts. Growth cavities are lined by isopachous fibrous calcite cement (former magnesian calcite cement; white arrow) and filled with early marine botryoidal calcite (originally aragonite) cement. Note thrombolitic microfabric (black arrow) characterized by clots of small peloids surrounded by a thin calcite cement rim. The sample typifies a reef limestone formed at the transition of a platform margin to an upper slope. B: Polymict breccia consisting of poorly sorted carbonate clasts. Differences in shape, roundness and size of the clasts point to reworking and redeposition of lithified and semi-lithified material. The clasts are peloidal packstones eroded in platform interior environments. Arrow points to associated volcaniclastic particles. Rare fossils are an echinoid spine (E) and shells. Matching boundaries of some clasts point to in-situ brecciation and displacive cement growth subsequent to the mechanical integration of the allochthonous carbonate (see Blendinger 2001). Middle Triassic (Ladinian): Mahlknecht Cliff, Dolomites, Italy. After Brandner et al. (1991).
grainstone and packstone) and flat graded grainstones consisting of redeposited platform-interior and platform-margin carbonate sands at the toe-of-slope. Turbidites exhibiting units of the Bouma Sequence occur in a proximal and distal position. The latter are characterized by thin grainstone and packstone beds with skeletal grains and peloids. The samples shown in Fig. 15.24 are constituents of foreslope breccias forming platform margin to basin intercalations within slope sediments (See Fig. 14.17). Most of the allochthonous material was eroded from platform and platform-margin limestones. The eroded material was transported by debris flows and rockfalls to the slope. Erosion was triggered by sea-level fluctuations. See Pl. 115/2 for another picture of these allochthonous carbonates.
15.7.3.3 Jurassic of Morocco: Platform-Slope-Basin Transect A platform-slope-basin system formed during the Late Jurassic in the Tabas area of eastern central Iran is a good example of microfacies types of low-angle slope carbonates (Fürsich et al. 2003; Fig. 15.25). The export of platform material into the slope and basin is interpreted as highstand shedding related to relative sealevel fluctuations. The slope contains a mixture of allochthonous and autochthonous sediments. Autochthonous sediments are oncolitic floatstones and rudstones, and spiculitic wackestones. Mud flow deposits are characterized by large, fragmented or complete fossils floating in a wackestone to packstone matrix, and by bioclastic grainstones (debris flows).
Jurassic Slope Carbonates
785
Fig. 15.25. Microfacies of Jurassic slope carbonates. The samples come from the Jurassic Qal’ch Limestone Formation: east of Koron, Shotori Mountains (Tabas area; central Iran). The unit is Middle Callovian to Late Oxfordian in age. The unit includes slope-to-basin sediments adjacent to the Esfandiar platform, consisting of an extended platform interior and a small high-energy platform margin. The slope varies laterally from a low-angle depositional slope to a steeper bypass slope. Common autochthonous microfacies types are microbe-sponge carbonates with strong biogenic encrustations (–> A) formed on the upper and middle slope; packstones with encrusting Tubiphytes (–> B) occurring in close association with the microbialites; and oncolitic carbonates formed in protected areas on the upper slope (–> E). Allochthonous grains shed from the platform margin are ooids (–> C) and bioclasts (–> D). This material occurs in calciturbidites and debris flows intercalated in slope and basin deposits. The autochthonous background sedimentation is represented by spiculitic wackestones (–> F). A: Microbialite with serpulids (S). The black arrow points to smaller calcareous worm tubes (Terebella), the white arrow to an internal growth cavity filled with peloidal sediment. B: Skeletal packstone with Tubiphytes (T), crinoid fragments (white) and bryozoans (N: Neuropora). Neuropora is a common bryozoan in many Late Jurassic platform carbonates. C: Ooid packstone with aggregate grains (bottom), crinoid fragments and encrusted and coated shells. Calciturbidite. Highstand shedding. D: Poorly sorted, lithoclastic bio-rudstone. Calciturbidite. Mid-slope lowstand wedge. E: Oncoid within an oncolithic floatstone. Note the abundant bioclastic debris and foraminifera incorporated in microbial laminae. F: Fine-bioclastic wackestone with sponge spicules and some ostracods (arrows). Scale is 1 mm. All pictures from Fürsich et al. (2003).
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15.7.3.4 Jurassic of the Northern Calcareous Alps: Detailed Information from Limestone Turbidites on Source and Deposition Patterns Late Jurassic sedimentation in the Northern Calcareous Alps is characterized by carbonate platforms with reefs (Plassen Formation), and pelagic carbonates deposited in deep-marine basins (Oberalm Formation). Slope carbonates are characterized by a mixture of allochthonous and autochthonous sediments. Autochthonous sediments are formed by benthic organisms growing on lithified surfaces. Allochthonous carbonates consist of fine-grained and sand- and gravel-sized material exported from the platform and platform margins to the slope and the basin. One of the best studied examples of allochthonous carbonates is the Tithonian Barmstein Formation (Steiger 1981). The Barmstein limestones are exposed between Hallein near Salzburg, Austria, and Berchtesgaden, Bavaria. The thick-bedded carbonates are predominantly turbidites, sometimes with a fluxoturbidite at the base. Most of the coarseand fine-grained resedimented material was eroded from the Plassen platforms that originated on isolated highs. The limestones were investigated by means of qualitative and quantitative microfacies analysis, texture analysis describing the composition of the sedimentary sequences, grain-size analysis including the statistical evaluation of grain-size parameters, and paleontological analysis. The latter comprises a systematic treatment and paleoenvironmental interpretation of thin-section fossils. Microfacies data on rock fragments and lithoclasts, diagenetic criteria and the abundance, association and taphonomic features of fossils were used in reconstructing the facies types of the Plassen platforms acting as source areas. The microfacies of Triassic and Jurassic extraclasts from the section below the Late Jurassic platform carbonates reveals the intensity and course of submarine erosion over time. The destruction of the platforms and the increase in slope sediments was triggered by diapirism and related tectonics. The study shows the great potential of the combined study of microfacies, fossils and sedimentological criteria. Of particular interest is the differentiation of platform material that was lithified at the time of erosion and redeposition, non-lithified platform sediments and organisms living in various parts of the destroyed platforms, and limestone clasts representing input from subaerially exposed parts of the source area. The study of field logs and the consideration of the geometry and extension of the turbidites and their distribution within the small depositional basins (a few tens of kilometers in size) result in a conclusive model of deep-water sedi-
Slope Stability
mentation in basins with strong input from intrabasinal platforms and reefs. The methodological approach of the study and of this model is recommended.
15.7.4 Slope Stability Reflected by Texture and Microfacies Carbonate platform flanks are almost exclusively made up of sediments shed by the platform itself. The composition of this platform overproduction determines the angle that the sediment can maintain (Kenter and
Fig. 15.26. Slope angles of carbonate platforms are controlled by the dominant depositional fabric. Grainy, non-cohesive mud-free buildups form steeper slopes than muddy, cohesive sediments. Grain-supported fabrics with minor or no matrix (grainstones, rudstones, clast-supported breccias; e.g. Gosaukamm; Pl. 27/2) build slope angles varying up to about 40°. Grain-supported textures with matrix (packstones; e.g. Late Jurassic mounds in Germany) and mud-supported fabrics exhibit slope angles up to 15°. Mudstones (e.g. modern Little Bahama Bank or Late Paleozoic algal mounds) have low slope angles up to 5°, mud-supported wackestones and floatstones up to about 15°. Well-documented examples are marked by circles; examples lacking precise control of geometry by squares; and examples whose flanks are stabilized by microbial automicrite (e.g. Waulsortian mounds, or the Sella Platform in the Dolomites) and/or synsedimentary carbonate cements or organic framework by triangles. Early stabilization results in medium to steep slope angles. The range of fabric and slope angles is shown by bars for each data point. These data indicate that the physical behavior of the sediments in the gravity field, angle of shearing and readjustment processes are linked to the composition of the slope sediment. Textural composition appears to be the major control on slope angle and slope curvature besides windwardleeward setting, sea level and climate. Based and slightly modified from Kenter (1990), with additional data for the Gosaukamm (own observations) and for the Sella (Keim and Schlager 1999). A list of localities and references can be found in Kenter (1990).
Stabilization of Carbonate Slopes
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Fig. 15.27. Stabilization of steep carbonate slopes by microbial automicrite and early cement. The examples demonstrate the important role of microbially mediated, in-situ precipitated micritic carbonate in the early lithification of steeply dipping slopes. However, slope geometry is ultimately controlled by the angle of repose of layers differing in texture and grain size (see Fig. 15.26). Triassic (Late Ladinian to Early Carnian): Slope of the Sella Platform, South Tyrol, Italy. A: Planar beds on the slope. Mechanically deposited sediment. Fine-grained grainstone consisting of micritized peloids and skeletal grains. B: Microbial automicrite of steep slope beds. The microfabric differs significantly from that shown in –> 1. Note the very irregular outlines of the peloidal micritic aggregates and the sparry calcite cement areas. The aggregates consists of peloids forming a loosely connected framework. Bioclasts are very rare. Water depth is estimated at 30 m based on the vertical distance from correlative horizontal topset beds. Automicrite can be followed down to water depths of about 200 m. C: Automicrite layer of the topsets. Internal cavity alternately filled by cement and sediment. The inclined cavity fillings consisting of fine-grained peloidal sediment are interpreted as small foresets created by bedload transport. The growth of fibrous cement is interrupted by thin micrite linings (arrows). D: Breccia of the steeply dipping slope deposits. Slope breccias are cement-rich (up to 50% of the rock volume), and almost matrix-free. The sample shows the high amount of calcite cement, clotted micrite (M), various tilted generations of geopetals (arrows) as well as millimeter-sized cracks cutting through the fibrous cement. Note the different types of internal sediments. The last phase of internal sedimentation is an infilling of basinal sediment (B) into a cavity. The same sediment fills cracks (circle) cutting through the cement, demonstrating syndepositional cementation and fracturing. Cracks filled with white calcite (top left) represent the youngest tectonic events. A, B and C: Pordoi, southern margin of the Sella platform, Dolomites, Italy. D: Val de Mesdi, northern Sella. Scale is 2 mm. All photographs from Keim and Schlager (1999).
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Tracing Platform-Basin Transitions
mounds). Slopes between 5° and 25° are typical for transitional mixtures between muddy and granular fabrics.
Fig. 15.28. Western Sella Platform, Passo Pordoi, Dolomites. Bedding geometry at the platform margin. Horizontal topsets (bioclastic pack-, grain- and rudstones) showing planar bedding and bed thickness in the range of decimeters to one meter pass into the steeply dipping slope by thickening and steepening in a basinward direction. The slope consists of irregularly distributed breccias and blocks, and grain- and packstone beds stabilized by automicrite. Primary depositional dip angles reach 35°. The topsets of the Triassic (Late Ladinian to Early Carnian) slope of the Schlern Dolomite are overlain by the Carnian Raibl Formation, a siliciclastic unit recognizable over the Alpine Chains. The Norian Dolomia Principale is a peritidal complex. See Fig. 15.27 for pictures of the microfacies. After Keim and Schlager (1999).
Schlager 1987; Kenter 1990). Declivity and geometry of submarine slopes is controlled by the composition of slope sediments. Calcareous sediments build steeper slopes than siliciclastic sediment. Steep slopes may be subdivided into three parts: (1) An upper part dominated by avalanching of coarse gravelly sandy sediment, low in mud (lithoclastic rudstone with abundant marine cement, e.g. radiaxial cement), separated from the middle and lower slope by a distinct break. (2) Middle slope characterized by scouring and deposition of gravelly debris flow (poorly sorted lithoclastic rounded grain- to rudstone, gravel and sand mixture). (3) Lower slope with sedimentation of concave muddy deposits, muddy sands or mudfree sands (composite beds with mudstone and wackestones, and fine-grained lithoclastic sand. The controls on the stabilization of slopes are a crucial problem in facies interpretation, particularly of steep platform slopes or the slopes of mud mounds, which can dip down with angles of between 30 and 75 degrees (e.g. Middle Devonian mud mounds, Morocco; Wendt 1993). An important control on the stability of slope angles are sediment fabrics reflected by the texture, size and abundance of grains (Fig. 15.26). Grainsupported carbonates, devoid of fines, pile up to over 25° (e.g. Triassic of the Dolomites). Mud-supported fabrics build slopes up to 5° (e.g. Carboniferous algal
Other factors influencing slope stability are microbially triggered cementation and automicrite covers. Experiments with sand seeded with bacteria and fungi show that microbes influence the angles of avalanche and angles of repose and increase stability (Meadows et al. 1994). Bacteria bind particles together with extracellular polymeric material, while fungi bind particles by holding them together with a network of hyphal filaments. Microbial automicrite and synsedimentary cementation are essential controls on the stability and preservation of steep and high carbonate slopes together with textural composition (Keim and Schlager 1999). In the Triassic of the Dolomites (Fig. 15.28), automicrite formed by a combination of precipitation in microbes, microbially mediated and inorganic precipitation extend from the outer part of flat platforms down to over 200 m on 25°–35° dipping slopes (Figs. 15.27, 15.28). The microfacies of these automicrites is characterized by peloidal boundstones (Fig. 15.27B). The autochthonous micritic carbonate lithified almost instantly, and together with fibrous cement, intermittently stabilized the platform slope. Downslope transport of fragmented automicrite layers (Fig. 15.27D) forming slumps and debris flows contributed to the deposition of cipit boulders in distal toe-of-slope positions.
15.7.5 Tracing Platform-Basin Transitions Using Grain Composition Logs 15.7.5.1
Concept and Methods
Shedding of fine-grained carbonate material from platforms, platform margins and upper slopes to adjacent periplatform environments, connected with the transport of the grains by turbidity currents and debris flows lead to the formation of calciturbidites intercalated in basinal carbonates. Compositional grain analysis and statistical evaluation of the compositional data provide important information on the history of exposure and flooding of platforms, the timing of cyclic sedimentation, and stratigraphic platform-to-basin correlations. The interpretation of compositional logs requires knowledge of the depositional environments of carbonate grains and thorough numerical analyses of grain composition, frequency and association. Platformslope-basin correlations by compositional logs should be combined with other methods, e.g.
Platform-Basin Transitions
• correlation by physical tracing of laterally continuous bedding surface of platform tops and platform flanks; • biostratigraphic correlation, limited by ecological and preservation factors; • well-log marker, e.g. condensation horizons, light stable isotopes, element chemostratigraphy, foraminiferal abundance, organic carbon (see Kauffman 1988). The principal assumption in tracing platform-basin relationships by compositional logs is that compositional variations in redeposited shallow-water material within basinal sediments adjacent to platforms reflect variations in water depths at the top of the platform and sea-level fluctuations as demonstrated for Late Quaternary interglacial and glacial calciturbidites differing in grain composition (Reijmer et al. 1988; Haak and Schlager 1989). The abundance increase and decrease in specific grains may indicate evolutionary stages of platforms, e.g. drowning of platforms associated with a decrease in ooids, or prograding and aggrading of platforms by an increase in the abundance of coarse skeletal sands. The intensity of shallow-water sediment export may be different at different flanks of the source area, e.g. at the opposite sides of a reef (Betzler et al. 1995). The timing of strong shedding is controversial: • Shedding during sea-level highstands (Droxler and Schlager 1985; Reijmer et al. 1988; Haak and Schlager 1989; Glaser and Droxler 1991; Schlager 1992). • Shedding during the late highstand (Shanmugan and Moiola 1988). • Shedding during sea-level lowstand. The basic prerequisite for the success of the Grain Composition Log method is the knowledge of the specific occurrence of grain categories and fossils in specific source areas, e.g. specific platform and reef environments. This knowledge presupposes that the distribution patterns and their ecological controls of platform and reef organisms of a particular time slice are already well established. Since distribution patterns of benthic organisms have changed over time, compositional logs of calciturbidites will be only interpreted successfully if paleontologically well-founded zonation patterns for the specific time slice are available (Reijmer and Everaars 1991; Reijmer et al. 1991; Everts 1991, 1999). Caution: Similar or identical biota may indicate platform interior environments for one time slice, and marginal platform environments for another time slice. Some common fossil groups (e.g. echinoderms, brachiopods) can hardly be used as indicators
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of a particular depositional environment without taxonomic treatment or at least the consideration of ecomorphological groupings. Seek out cooperations with paleontologists! 15.7.5.2 Case Study: Late Triassic of the Gosaukamm Region, Austria The potential of microfacies analyses for interpreting platform-to-basin relations is exemplified by impressive case studies dealing with Late Triassic (Norian and Rhaetian) depositional environments in the Northern Calcareous Alps (Fig. 15.29). Shallow-marine environments are represented by the Dachstein Formation consisting of thick series of platform interior carbonates with marginal coralline sponge and coral reefs. The rhythmically bedded platform carbonates display the ‘Lofer cycles’ reflecting sea-level fluctuations of different orders (see Sect. 16.1). The platform margin passes into the deeper marine Hallstatt basin filled with marls and pelagic carbonates and gravity flow deposits including calciturbidites. The periplatform ‘Pedata/ Pötschen beds’, representing the basinal equivalent of the Dachstein Formation, are characterized by an abundance of calciturbidites and submarine slides. These formations are exposed in the Gosaukamm area about 60 km southeast of Salzburg, Austria. Microfacies and paleontological data of the Dachstein limestone are well known, allowing the paleoenvironmental significance and distribution of the shallow-marine biota in different parts of the platform and the reefs (Wurm 1982) to be evaluated. Reijmer and Everaars (1991) and Reijmer et al. (1991) studied a 95 m thick section of the Late Norian Pedata/Pötschen limestone exhibiting 810 calciturbidite beds. The Lacke section comprises mud-, pack- and grainstones, alternating with calcisiltites and marls. The mudstone to grainstone beds show grading and sharp contacts and correspond to turbidites. Calcisiltites and marls represent the pelagic background sediments. Point-counter analyses of samples taken from the top and bottom, in thicker beds also from the middle part of each turbidite, were the basis for the statistical treatment of the frequency data of grains and matrix. The constituents encountered in thin sections were summarized into facies-diagnostic groups and assigned to seven point-counter groups (clusters) that characterize particular paleoenvironments of the suggested source areas on the platform (e.g. platform interior), within the reef (e.g. shallow reef, deeper forereef) and in the open ocean. Summary statistics revealed a close relationship between the biota present on the Dachstein platform and the composition of the calciturbidites.
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Platform-Basin Transitions
Plate 138 Tracing Platform-Basin Relations by Statistical Clusters of Late Triassic Calciturbidites The plate displays statistically confirmed ‘clusters’ of calciturbidites occurring within the basinal Norian basinal Pedata/Pötschen Beds of the Lacke section near the Gosaukamm (Austria). The Pedata/Pötschen limestones are bedded micritic carbonates with abundant turbidites (packstones and grainstones) and submarine slumping and sliding structures. The clusters are defined by cluster analysis and the dominant grain composition and evaluated by factor analysis and correspondence analysis. The analysis of the total section resulted in the differentiation of seven clusters. Six clusters (shown in this plate) represent 744 of 747 samples (99.5%). The individual clusters can be defined by the relative amount of one or more point-count groups compared to the overall means. The source area of the turbidite grains is deduced from the distribution patterns of platform and reef biota of the GosaukammDachstein platform, particularly of foraminifera that are strongly facies-controlled (see Pl. 111). The grain composition of the calciturbidites indicate alternations of (1) turbidites with predominantly platform-top derived grains (shallow reef, platform interior), and (2) turbidites containing mainly basin-derived material, planktonic and pseudoplanktonic material in combination with micrite. Factor analysis show that three factors account for 88% of the total compositional variation. Factor 1 (50% of the total variables) is controlled by the variable groups representing the open ocean input versus shallow-marine biota and clasts. Factor 2, explaining 27% of the variation, is loaded with clasts and open ocean input. The variation in factor 3 is due to platform interior material combined with shallow reef biota and micrite. Factors 2 and 3 can be explained as dilution by mud related to changes in productivity or the supply of grains. The contrast between the clusters A and B and D, the clusters E and G are best explained as fluctuation between deep-water sedimentation and platform-derived input, corresponding to pronounced highstand shedding from platforms to deep basins. During the lowstand carbonate productivity was low, and the shallow reef and platform interior stopped producing carbonate. Sediment deposition was determined by open marine biota together with material exported from protected and/or somewhat deeper reefs. 1 Cluster A. Calciturbidite characterized by abundant filaments (open-marine biota) and increasing micritic matrix. High input of open oceanic biota and deep reef biota and low input of shallow reef and platform interior biota. Corresponds to Standard Microfacies Type SMF 3. 2 Cluster B, characterized by dominance of open-marine biota and low amount of micritic matrix. The arrow points to Variostoma crassum Kristan, a characteristic foraminifer of the basinal Hallstatt facies. Highest mean input of open ocean material combined with a minimum of shallow-reef material and non-facies-specific biota. Corresponds to SMF 3. 3 Cluster C. The sediment consists of open ocean originated grains and platform derived grains, embedded in micritic matrix. Open ocean biota are represented by thin shelled bivalves (‘filaments’, F) of the Halobia-Posidonia group and the planktonic alga Globochaete (most tiny globular structures). Platform derived grains are foraminifera (PF). Cluster C comprises the largest group (399 samples). These turbidites were deposited during the transition from A to B, and D to G. 4 Cluster D. The sediment is characterized by an increase in the input of shallow reef and platform interior material and a decrease in micritic matrix. Shallow reef material is represented by characteristic reef foraminifera (RF; see Pl. 111) and peloids reworked from reef cavities. These peloids are characterized by microbially induced calcite rims. Echinoderm (E) fragments might have had their source area in upper slope environments. Cluster D covers 228 samples. The clusters D, E and G correspond to SMF 4. Characteristic of high input of shallow-reef and platform interior material and low input of micritic sediment. 5 Cluster E. As compared with –> 4, the input of platform interior grains is much higher. Note the abundance of platformderived foraminifera (PF) and echinoderms (E) associated with shallow-reef foraminifera (RF) a few cemented reefderived lithoclasts (LC) and reworked reef crusts (RC). Clusters E and G are dominated by the high input of platformoriginated grains like clasts, platform interior grains and shallow reef biota. Cluster E is more pronounced in that respect than cluster G. 6 Cluster G. High input of platform derived grains (small peloids), platform interior material (foraminifera) and reef material (black fragments). Note the occurrence of fragmented reef biota comprising recrystallized coralline sponges and reef crusts (RC), and interreef material comprising pseudopunctate brachiopod shells (B) and echinoderms (E). The increase in grain size from cluster D to G (–> 4 to –> 6) reflects changes in high-energy conditions. –> 1–6: Reijmer and Everaars (1991)
Plate 138: Clusters of Late Triassic Calciturbidites
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792
Basics: Platform Margins and Slopes
Fig. 15.29. Strongly generalized facies distribution in the Late Triassic of the Northern Calcareous Alps. The bedded Dachstein limestone of the platform interior is separated from the Hallstatt Basin by a reef belt built by coralline sponges, corals and reef detritus. Reefs were formed in shallow and slightly deeper and protected environments. Foreslope sediments in the Gosaukamm area correspond to the forereef slope. The slope grades into the Pedata/Pötschen Beds, representing dark slope and basinal bedded limestones with cherts and marls. Sedimentary folding structures indicate submarine sliding. Calciturbidites are common and were the focus of the study described in Pl. 138. The gray and red Hallstatt Limestone is an open-marine pelagic and hemipelagic sediment.
Numerical treatment of grain associations of samples using cluster analysis and factor analyses, and evaluated by correspondence analyses display significant variations in platform- and reef-derived input of shallow-water sediment to the slope and basin during time. The compositional variations of the calciturbidites and cyclic patterns revealed by spectral analysis are conceived as a response to sea-level fluctuations that controlled sediment production on the platform and the intensity of shedding to the basin. Different methods of frequency spectral analyses must be applied in searching for possible relationships between rhythmicity and orbital forcing and distinguishing regular cyclicities from spurious peaks (Reijmer et al. 1994; Reijmer 1998). Basics: Platform margins and carbonate slopes Betzler, C., Pfeiffer, M., Saxena, S. (2000): Carbonate shedding and sedimentary cyclicities of a distally steepened carbonate ramp (Miocene, Great Bahama Bank). – International Journal of Earth Sciences, 89, 140-153 Biddle, K.T., Schlager, W., Rudolph, K.W., Bush. T.L. (1992): Seismic model of a progradional carbonate platform, Picco di Vallandro, the Dolomites, Italy. – American Association of Petroleum Geologists, Bulletin, 76, 14-30 Bouma, A.H. (1962): Sedimentology of some flysch deposits. – 168 pp., Amsterdam (Elsevier) Bouma, A.H., Stone, C. (eds., 2000): Fine-grained turbidite systems. – American Association of Petroleum Geologists, Memoir, 72, 342 pp. Brachert, T. C., Dullo, W.C. (1994): Micrite crusts on Ladinian foreslopes of the Dolomites seen in the light of a modern scenario from the Red Sea. – Abhandlungen der Geologischen Bundesanstalt, 50, 57-68 Brandner, R., Flügel, E., Senowbari-Daryan, B. (1991): Microfacies of carbonate slope boulders: indicator of the source area (Middle Triassic: Mahlknecht Cliff, Western Dolomites). – Facies, 25, 279-296 Coniglio, M., Dix, G.R. (1992): Carbonate slopes. – In: Walker, R.G., James, N.P. (eds., 1992): Facies models. Response to sea level change. – 349-374, Geological Association of Canada
Cook, H.E., Enos, P. (eds., 1979): Deep-water carbonate environments. – Society of Economic Paleontologists and Mineralogists, Special Publications, 25, 336 pp. Cook, H.E., Hine, A.C., Mullins, H.T. (eds., 1983): Platform and deep water carbonates. – Society of Economic Paleontologists and Mineralogists, Short Course, 12, 563 pp. Cook, H.E., Mullins, H.T. (1983): Basin margin environment. – In: Scholle, P.A., Bebout, D.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists, Memoir, 33, 539-617 Crevello, P.D., Harris, P.M. (eds., 1985): Deep-water carbonates: buildups, turbidites, debris flows and chalks. – Society of Economic Paleontologists and Mineralogists, Core Workshop, 6, 557 pp. Eberli, G. (1991): Calcareous turbidites and their relationship to sea-level fluctuations and tectonism. – In: Einsele, G., Ricken, W., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 340-359, New York (Springer) Edwards, D.E. (ed., 1993): Turbidity currents. Dynamics, deposits and reversals. – Lecture Notes in Earth Sciences, 44, 755 pp. Einsele, G. (1991): Submarine mass flow deposits and turbidites. – In: Einsele, G., Ricken, W., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 313-339, New York (Springer) Engel, W. (1974): Sedimentologische Untersuchungen im Flysch des Beckens von Ajdovscina (Slowenien). – Göttinger Arbeiten zur Geologie und Paläontologie, 16, 1-65 Everts, A.J. (1994): Carbonate sequence stratigraphy of the Vercors (French Alps) and its bearing on Cretaceous sea level. – Ph.D. Thesis, Vrije Universiteit, Amsterdam, 176 pp. Flügel, E., Di Stefano, P., Senowbari-Daryan, B. (1991): Microfacies and depositional structure of allochthonous carbonate base-of-slope deposits: The Late Permian Pietra die Salomone megablock, Sosio Valley (Western Sicily). – Facies, 25, 147-186 Fürsich, F.T., Wilmsen, M., Seyed-Emami, K., Axhairer, G., Majidifard, M.R. (2003): Platform-basin transect of a Middle to Late Jurassic large-scale carbonate platform system (Shotori Mountains, Tabas Area, East-Central Iran). – Facies, 48, 171-198 Haak, A.B., Schlager, W. (1989): Compositional variations in calciturbidites due to sea-level fluctuations, late Quaternary, Bahamas. – Geologische Rundschau, 78, 477-486 Herbig, H.-G., Bender. P. (1992): A eustatically driven
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Pelagic Carbonates
calciturbidite sequence from the Dinantian II of the Eastern Rheinisches Schiefergebirge. – Facies, 27, 249-262 Hohenegger, J., Yordanova, E. (2001): Depth-transport functions and erosion-deposition diagrams as indicators of slope inclination and time-averaged traction forces: applications in tropical reef environments. – Sedimentology, 48, 1025-1046 James, N.P., Mountjoy, E.W. (1983): Shelf slope break in fossil carbonate platforms: An overview. – In: Stanley, D.J., Moore, G.T. (eds): The shelf-break: Critical interface in continental margins. – Society of Economic Paleontologists and Mineralogists, Special Publication, 33, 189-206, Tulsa Keim, L., Schlager, W. (1999): Automicrite facies on steep slopes (Triassic, Dolomites, Italy). – Facies, 41, 15-26 Kenter, J.A.M. (1990): Carbonate platform flanks: slope angle and sediment fabric. – Sedimentology, 37, 777-794 Kenter, J.A.M., Campbell, A.E. (1991): Sedimentation on a Jurassic carbonate platform flank: geometry, sediment fabric and related depositional structures (Djebel Bou Dahar, High Atlas, Morocco). – Sedimentary Geology, 1-34 Kenter, J.A.M., Schlager, W. (1989): A comparison of shear strength in calcareous and marine sediments. – Marine Geology, 88, 145-152 Meischner, K.-D. (1964): Allodapische Kalke. Turbidite in riffnahen Sedimentationsbecken. – In: Bouma, A.H., Brouwer, A. (eds.): Turbidites. –Developments in Sedimentology, 3, 156-191 Mullins, H.T., Cook, H.E. (1986): Carbonate apron models: alternatives to the submarine fan model for paleoenvironmental analysis and hydrocarbon exploration. – Sedimentary Geology, 48, 37-79 Mutti, M., Ricci Lucchi, F. (1978): Turbidites of the northern Apennines: introduction to facies analysis. – Geology Review, 20. 125-166 Reijmer, J.J.G., Everaars, J.S.L. (1991): Carbonate platform facies reflected in carbonate basin facies (Triassic, Northern Calcareous Alps, Austria). – Facies, 25, 253-288 Reijmer, J.J.G., Schlager, W., Bosscher, H., Beets, C.J., McNeill, D.F. (1992): Late Pliocene platforn facies transition recorded in calciturbidites (Exuma Sound, Bahamas). – Sedimentary Geology, 78, 171-179 Reijmer, J.J.G., Sprenger, A., Ten Kate, W.G.H.Z., Schlager, W., Krystyn, L. (1994): Periodicities in the composition of Late Triassic calciturbidites (Eastern Alps, Austria). – In: De Boer, P., Smoth, D.G. (eds.): Orbital forcing and cyclic sequences. – International Association of Sedimentologists, Special Publication, 19, 323-343 Reijmer, J.J.G., Ten Kate, W.G.H.Z., Sprenger, A., Schlager, W. (1991): Calciturbidite composition related to exposure and flooding of carbonate platforms (Triassic, Eastern Alps). – Sedimentology, 38, 1059-1074 Russo, F., Mastandrea, A., Neri, C. (1998): The Cipit boulders from the surroundings of the Sella Massif. – Giornale di Geologia, Ser. 3, 60, Special Issue, ECOS-VII-Southern Alps Field Trip Guidebook, 108-115 Schlager, W., Camber, O. (1986): Submarine slope angles, drowning unconformities and self erosion of limestone escarpments. – Geology, 14, 762-765 Schlager, W., Reijmer, J.J.G., Droxler, A.W. (1994): Highstand shedding of rimmed carbonate platforms - an overview. – Journal of Sedimentary Research, B64, 270-281 Shanmugam, G. (2000): 50 years of turbidite paradigm (1950s-1990s): deep water processes and facies models -
a critical perspective. – Marine and Petroleum Geology, 17, 285-342 Spence, G.H., Tucker, M.E. (1997): Genesis of limestone megabreccias and their significance in carbonate sequence stratigraphic models: a review. – Sedimentary Geology, 112, 163-193 Stanley, D.J., Moore, G.T. (eds., 1983): The shelf-break: Critical interface in continental margins. – Society of Economic Paleontologists and Mineralogists, Special Publication, 33, 467 pp. Steiger, T. (1981): Kalkturbidite im Oberjura der Nördlichen Kalkalpen (Barmsteinkalke, Salzburg, Österreich). – Facies, 4, 215-348 Further reading: K178, K180, K192
15.8 Pelagic Deep-Marine Carbonates Deep-marine basins in open-marine and deep shelf settings are characterized by pelagic sediments. Carbonate pelagic sediments originate from particle-by-particle sedimentation of skeletal debris of planktonic organisms from suspensions. Today coccolithophorids (planktonic algae; Pl. 7/3) and pteropods (pelagic gastropods) make up the calcareous oozes distributed in deeper-water settings, including the ocean floor, seamounts, mid-oceanic ridges and outer continental shelves (Sect. 2.4.5.3). Pelagic carbonates have been reviewed by Scholle et al. (1983), Leggett (1985), Jenkyns (1986), and Tucker (1990). Pelagic limestones are well represented in the Mesozoic and Cenozoic, but they appear to be less common in the Paleozoic (Box 15.4). Most studies describing ancient ‘deep-marine’ limestones are focused on allochthonous carbonates deposited in slope or basinal settings. Microfacies studies of autochthonous pelagic carbonates make up less than 20% of the total of papers dealing with the lithofacies and biofacies of carbonate rocks. One of the reasons for this significant bias may be the rather simple and somewhat monotonous composition of pelagic limestones in thin sections. Pelagic carbonates must be studied using combined geochemical, sedimentological and paleontological methods.
15.8.1 Setting, Controls and Biota of Pelagic Limestones Setting: Pelagic carbonates are deposited over a wide range of settings from continental margins on drowned carbonate platforms to slopes and to basins (Fig. 15.30).
794
Box 15.4. Selected references on the criteria and microfacies of ancient pelagic and deep-water basinal carbonates. References on ancient hardgrounds are listed in Box 5.5. Cretaceous and Tertiary Chalk: Bromley and Ekdale 1987; D’Heur 1984; Feazel et al. 1985; Feazel and Farrell 1988; Hakansson et al. 1974; Hardman 1982; Hardman and Kennedy 1980; Kennedy 1980; Maliva and Dickson 1992; Neugebauer 1973, 1974; Schatzinger et al. 1985; Scholle 1977; Van den Bark and Thomas 1980 Jurassic: Alvarez 1989; Bernoulli 1971; Bernoulli and Jenkyns 1970; Blendinger 1988; Castellarin 1970; Catalano et al. 1977; Bice and Stewart 1990; Böhm 1992; Böhm et al. 1999; Ebli 1997; Evans and Kendall 1977; Garrison 1967; Garrison and Fischer 1969; Gawlick et al. 1999; Germann 1971; Hollmann 1962, 1964; Jenkyns 1970, 1971; Kuhry et al. 1976; O’Dogherty et al. 2001; Sanantonio 1993; Seyfried 1980; Soussi and Ben Ismail 2000; Wierczbowski et al. 1999; Wilson 1969; Winterer and Bosellini 1981 Triassic: Bachmann and Jacobshagen 1974; Blendinger 1988; Dürkoop et al. 1986; Krystyn et al. 1971; Schlager 1974; Wendt 1969, 1973; Zankl 1971 Carboniferous: Bissell and Parker 1977; Yurewicz 1977 Devonian: Bandel 1974; Tucker 1973, 1974; Matti et al. 1975; Tucker and Kendall 1973; Vai 1980; Wendt and Aigner 1985; Wendt 1988 Ordovician: Cook and Taylor 1977; Ferretti and Kriz 1995; Jaanusson 1955, 1973; Lindström 1963, 1971, 1979; McBride 1970.
‘Pelagic carbonate platforms’ (Catalano et al. 1977) corresponding to type B of Fig. 15.30, are submarine intrabasinal highs representing drowned fragments of ancestor shallow-marine platforms and ramps (Santantonio 1993). They correspond to horsts or crests of rotated blocks and are common at rifted continental margins bordered by synsedimentary faults. These platforms are sites of condensed and discontinuous pelagic
Pelagic Carbonate Settings
sedimentation. In geological sections they are recognized by a pile of shallow-water limestones capped by a condensed pelagic sequence and bounded by paleofaults. Instructive examples were described by Wendt (1988) from the Devonian, and Winterer and Bosellini (1981) and Santantonio (1993) from the Jurassic. Other terms often used for pelagic platforms are seamount, swell, plateau, guyot or haut-fond pélagique. The terms seamount and guyot, however, should be restricted to submarine highs built on oceanic crust. Controls on the facies of pelagic carbonates are the sedimentation rate determined by the organic productivity, sea-floor currents and circulation patterns, and the water depth relative to the calcite compensation depth. These parameters control sea-floor cementation and dissolution. Pelagic carbonates can be attributed to five major lithofacies categories (Tucker 1990; see Fig. 15.31). Another spectrum of pelagic lithofacies is reflected by the increasing clay content ranging from pure limestones through limestones with shaley partings and pressure solution seams to nodular limestones and then to shales (marls) with dispersed calcareous nodules of early diagenetic origin. This limestone to shale spectrum may result from • topography: pelagic limestone preferentially deposited on topographic highs, shales with nodules in topographic lows; • increasing clastic input and decreasing carbonate productivity; • increasing depth with loss of CaCO3 as the CCD is approached. Biota: Common planktonic and nectonic constituents of Paleozoic pelagic limestones are radiolarians, goniatite cephalopods (Pl. 90/2), nautilid cephalopods, ostracods, tentaculitids (Pl. 91/3), and styliolinids (Pl.
Fig. 15.30. Range of depositional sites for pelagic carbonates. After Tucker (1990). Pelagic sediments are deposited in relatively shallow to very deep environments. A: Pelagic sediments occurring in neptunian dikes of shallow platforms. B: Pelagic sediments occurring on isolated platforms caused by isolated fault-blocks and on seamounts; generally little pelagic sediment is deposited. Hardgrounds and Fe-Mn crusts are common. C: Thick sequences of pelagic carbonates may develop upon subsiding platforms if there is little winnowing and dissolution. D: The pelagic facies on slopes is characterized by thin or thick resediments (Sect. 15.7.2). E: Thick pelagic carbonates will accumulate in basins. Marly intercalations and calciturbidites are common.
Paleozoic Pelagic Carbonates
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Fig. 15.31. Lithofacies categories of pelagic limestones resulting from variations in depth, current activity and clay content. After Tucker (1990). A: Thick sequences of parallel, thin-bedded limestones, sometimes with some lamination and bioturbation. Nearly continuous, out-of-suspension deposition. Common on deep-water slopes and subsided platforms. B: Bioturbated and nodular limestones. Slower sedimentation rates allow burrowing and destruction of lamination. Common on deep-water slopes and subsided platforms. C: Bedded limestones showing evidence of cementation in the form of intraclasts and hardgrounds with surfaces encrusted and bored. Common on topographic highs. D: Bedded limestones with evidence of sea-floor dissolution in the form of corrosional hardgrounds and molds of aragonitic bioclasts. Common on topographic highs. E: Pelagic skeletal sands originating in places where current winnowing removes lime mud, leaving larger pelagic skeletal grains (e.g. foraminifera, thin-shelled bivalves, styliolinids). Common on topographic highs. Fe-Mn encrustations and nodules are common in the categories C and D associated with discontinuity surfaces.
91/6). Common planktonic and nectonic constituents of Mesozoic pelagic limestones are coccolithophorids (Pl. 7/3, 5), planktonic foraminifera (Pl. 73), calcispheres (Pl. 73/2), radiolarians (Pl. 19/2; Pl. 76/1), calpionellids (Fig. 10.31), ammonites (Pl. 90/1), thinshelled bivalves (‘filaments’, Pl. 87/3), ostracods, and free-swimming crinoids (Fig. 10.56).
(Fig. 15.32) and contain abundant nautilid cephalopods and shells of trilobites, echinoderms, mollusks, ostracods, and bryozoans. Glauconite, hardgrounds and evidence of submarine lithification associated with nondeposition are abundant. Similar limestones but also dark, laminated pelagic lime mudstones are known from North America, associated with pelagic and hemipelagic black graptolite shales (Matti et al. 1975; Cook and Taylor 1977).
15.8.2 Examples and Case Studies of Pelagic Carbonates The following section presents examples displaying common microfacies types of Paleozoic and Mesozoic pelagic limestones formed in epeiric seas, at the margins of continents and in open ocean basins.
15.8.2.1 Microfacies of Paleozoic Basinal Carbonates Widely distributed pelagic Paleozoic microfacies types common in the Early Paleozoic are gray and red, often burrowed bioclastic wackestones and packstones with abundant hardgrounds. Paleozoic pelagic facies occur on epeiric ramps and at the margins of continents. Cambrian and Ordovician cephalopod limestones of the Baltic Shield Cephalopod limestones, informally named ‘Orthoceras limestones’ because of the occurrence of nautilid cephalopods, were formed in deep shelf environments of epeiric seas (Ferretti and Kriz 1996). These limestones are variously colored, often strongly burrowed
Fig. 15.32. Nodular bioclastic wackestone, an example of Early Paleozoic open-marine limestones deposited in deep shelf environments and regarded as pelagic sediments. The micritic matrix is very fine-grained. The source of the matrix is believed to be finely comminuted macrofossil debris (Lindström 1979). The sample is a burrowed skeletal wackestone. Most of the shells are trilobite fragments associated with some echinoderm plates. The picture is a detail of Pl. 19/1. Ordovician: Öland Island, Sweden.
796
Ammonitico Rosso Facies
carbonates of the peri-Mediterranean region are the Ammonitico rosso facies (Farinacci and Elmi 1981; Caracuel et al. 2000; Pl. 24/2) including the Adnet limestones in the Northern Calcareous Alps shown in Pl. 139. Many of these limestones were formed on the top of drowned Triassic carbonate platforms and represent examples of ‘pelagic carbonate platforms’ (Bernoulli 1972; Bernoulli and Jenkyns 1974). Depositional conditions were controlled by water depths (several hundreds of meters), oxygen levels near the sea bottom (producing gray or red rock colors), bottom current energy (low to moderate) and autochthonous or allochthonous sedimentation. Fig. 15.33. Styliolinid wackestone deposited on an intrabasinal swell. The origin of the fine-grained micrite matrix of the styliolinid limestones and the cephalopod limestones is unclear, but is commonly explained as resulting from a breakup of macroskeletons rather than from calcareous microfossils or nannofossils that are rare or missing in the Devonian. Middle Devonian (Givetian): Betic Cordillera, South Spain.
Devonian cephalopod limestones of western Europe Near the margin of the Old Red Continent pelagic limestones were formed adjacent to basins with silty shales on submarine highs (drowned reefs and volcanic seamounts), called ‘cephalopod limestone’ in Germany and ‘Griotte’ in France and Spain. The strongly condensed limestones consist of micrite. The biota are dominated by pelagic fauna of cephalopods (goniatites, Pl. 90/2; clymenids and orthocone nautilids), thinshelled bivalves, styliolinids (Pl. 91/6) and tentaculitids (Pl. 91/3; Sect. 10.2.4.5), ostracods, radiolarians, and conodonts. Styliolinid wackestones exhibit a distinctive and typical microfacies (see Fig. 15.33). The surfaces of hardgrounds are encrusted by sessile foraminifera and show evidence of submarine corrosion. Fe-Mn crusts and lithoclasts are common. Nodular fabrics are explained as the result of early lithification in the muddy calcitic sediments, and later pressure solution accentuating the nodular structure. The condensed cephalopod limestones were formed in different water depths on drowned reefs and platforms, volcanic seamounts, and basement highs (Bandel 1974; Tucker 1973, 1974; Wendt and Aigner 1985).
15.8.2.2 Microfacies of Mesozoic Basinal Carbonates Widely distributed Triassic deep-marine limestones are represented by the Late Triassic Hallstatt limestone (Pl. 23/3, 4; Pl. 43/2; Pl. 90/1). Common Jurassic pelagic
Microfacies of Jurassic limestones of the Ammonitico rosso and Adnet facies Microfacies criteria: Common microfacies criteria of Jurassic deep-marine limestones are • red, gray and white rock colors, • lime mudstones, wackestones and packstones, • stratigraphic condensation, • abundant hardgrounds and discontinuity surfaces, • Fe-Mn coatings and crusts; manganese nodules (see Sect. 4.2.4.1), • microbial (‘stromatolitic’) structures (Fig. 10.5), • cherts, • intercalations of allochthonous limestones (debris flows, turbidites), • neptunian dikes cutting Jurassic and underlying sediments, • ‘pelagic ooids’. Common microfacies types of Jurassic basinal limestones are (Wilson 1969; Kiessling 1996): • Dark pure lime mudstone, finely laminated on a millimeter scale. • Lithoclastic grainstone. Small rock fragments, allochthonous sediments. • Radiolarian wackestone and sponge-spicule wackestone. • Radiolarian packstone and radiolarian-spicule packstone. • Filament wackestone and packstone (Pl. 87/3). • Filament-Saccocoma-radiolarian wackestone and packstone. • Saccocoma wackestone and packstone (Fig. 10.56). • Very fine peloidal packstone or grainstone. Micropeloids derived from abraded skeletal or lithoclastic fragments upon the shelf and washed down into aprons at the foot of the slope. • Micro-biolithoclast packstone and grainstone (turbidites).
Pelagic Limestones
797
Adnet case study (Pl. 139): The Adnet limestones are variously bedded and nodular, red and gray limestones occurring in the central northern part of the Austrian and Bavarian Northern Calcareous Alps. The limestones exhibit all the microfacies criteria typical of pelagic carbonates deposited in different water depths below the photic zone (Fig. 15.36 on Pl. 139). Microfacies of the peri-Mediterranean Late JurassicEarly Cretaceous Aptychus and Biancone limestones The facies of the Late Jurassic Oberalm Formation in the Northern Calcareous Alps is an example of pelagic limestones widely distributed in the Alpine-Mediterranean region and better known under the names Aptychus limestones and Biancone. The facies is characterized by well-bedded, white micritic limestones, sometimes with marly partings and cherts. The carbonates comprise radiolarian limestones (Fig. 15.34; Pl. 76/1), calpionellid limestones (Fig. 10.30) and nannofossil limestones with coccolithophorids and nannoconids (see Fig. 4.4). The biotic composition and close association with radiolarites indicate deposition in very deep bathyal basins. Intercalations of turbidites (e.g. Barmstein limestones within the Oberalm limestones; Sect. 15.7.3.4) are common. The Early Cretaceous Biancone Formation consists of cherty, white to rose nannofossil micrite with calcite-replaced radiolarians and calpionellids.
Fig. 15.34. Radiolarian limestone. This example corresponds to the deposition site E of Fig. 15.30. The sample shows abundant radiolarians and sponge spicules. Note the different preservation of spumellarian and nasselarian radiolarians; some still exhibit characteristics of the test, whereas others are completely calcitized or infilled with micrite. Late Jurassic (Oberalm Formation, Tithonian): Hallein, Austria.
Microfacies of the Cretaceous Scaglia facies of the Venetian Alps, Italy The facies of the Cretaceous Scaglia Rossa Formation (Fig. 15.35), overlying the Biancone Formation, is an example of Late Cretaceous pelagic limestones that are widely distributed all over the Mediterranean region. The thin and evenly bedded micritic limestones of the Venetian Alps are typically developed as pink to red, slightly nodular skeletal wackestones rich in pelagic foraminifera and nannofossils. Trace fossils and
Fig. 15.35. Pelagic limestone with abundant planktonic foraminifera. Scaglia Rossa. This example corresponds to the deposition site B shown in Fig. 15.30. The sample is a bioclastic wackestone. All skeletal grains are shells of planktonic foraminifera or their debris. Late Cretaceous: Sirmione, Lake Garda, northern Italy.
798
Jurassic Adnet Limestone
Plate 139 Pelagic Deeper-Marine Limestones: Early Jurassic of the Northern Calcareous Alps In the Adnet area near Salzburg (Austria), Late Triassic reef and basinal carbonates are overlain by Jurassic sediments. The deposition of Liassic hemipelagic and pelagic sediments was controlled by the inherited Triassic relief. The basinal Triassic Kössen beds pass continuously into the Liassic Kendlbach Formation and further on into the gray cherty limestones of the Scheibelberg Formation. The former reef and upper slope were covered by red condensed limestones of the Adnet Formation characterized by poorly bedded limestones with ammonites (–> 9) and Fe-Mn crusts. Thin sections contain organic debris (–> 1, 4), ostracods, (–> 8), echinoderms (crinoids; –> 6), sponge spicules, dwarf gastropods, juvenile ammonites, foraminifera (nodosariids and involutinids, –> 7; textulariids), and Globochaete. Reworked hardground clasts (–> 8, 10) and intraclasts may be common. The source of the pelagic and hemipelagic micrite matrix was nannoplankton (Schizosphaerella), and probably also input from platforms. The microfacies was studied by Böhm (1992) and Böhm et al. (1999). The plate displays the microfacies of the limestones in the Sinemurian-Carixian Scheibelberg Formation (–> 1–3; Fig. 15.36) and the coeval red limestones in the Adnet Formation (–> 4–10) with well-known ammonitebearing limestones used since Roman times as decorative stone in southern and central Europe. 1 Gray burrowed microbioclastic packstone. Fossils are predominantly sponge spicules (arrows) and thin-shelled ostracods. The rock appears mottled because of burrowing. Interpretation: Lime muds consisting of calcareous plankton (Schizosphaerella) and deposited on the poorly aerated nutrient-rich deeper slope inhabited by abundant sponges. Water depths are believed to be several hundred meters. Scheibelberg limestone (Sinemurian): Osterhorn Mountains, Austria. 2 Cherts are common within the gray sponge-ostracod limestones. Take a look at the well-preserved sponge spicules (arrows). Nodular cherts in limestones can provide an important window on the original biota, sediment composition and diagenetic history of the host rocks, because of their low susceptibility to further diagenetic alteration. Same locality as –> 1. 3 Intercalations of crinoid-intraclast packstone within the Scheibelberg limestone, forming distinct, up to 30 cm thick beds. Arrows point to syntaxial rim cement around echinoderm fragments. Interpretation: Turbidites. The source area may be located on a swell with crinoid meadows. Plattenbruch, Adnet, southeast of Salzburg, Austria. 4 Transitional facies between the deeper Scheibelberg limestone and the red Adnet limestone, characterized by a red nodular microbioclastic wackestone. Interpretation: Deeper aerated slope environment. Breitenberg, Lake Wolfgang, Austria. 5 Red nodular Adnet limestone. Nodules (N) and matrix (M) are pelagic microbioclastic wackestones. Small calcite-filled structures are ostracods. Shells and echinoderms also occur. Interpretation: Mud deposition within a well-aerated slope environment below the photic zone and within the range of the aragonite lysocline. Nutrient input reduced by stronger bottom currents. Sporadical reduction in sedimentation rates resulted in complete oxidation of the sediment and patchy and local cementation of the mud. Kieferbruch, Adnet. 6 Red crinoid packstone/wackestone. Thin- to medium-bedded limestones. Interpretation: Accumulation of echinoderm fragments within depressions or fissures of swells caused by drowning of Upper Triassic reefs. Adnet, Austria. 7–9 Red lithoclastic crinoid-shell wackestone, characterized by diverse benthic foraminifera (–> 7: Involutina liassica: I and lagenids: L), echinoderms (E) and ostracods (–> 8), dissolved aragonite shells (–> 9, ammonites) and reworked hardground clasts, indicating the existence of strong bottom currents (–> 8). Interpretation: Well aerated deep-water environment, within the range of the aragonite lysocline. All samples from the Marmorbruch Adnet. 10 Red bioclastic wackestone/packstone with common hardground lithoclasts. Note the ferruginous coating (arrow). Prevailing fossils are ostracods and mollusk debris. Thick-bedded limestones. Interpretation: Upper slope environment with currents producing low sedimentation rates and erosion of hardground clasts. Adnet.
Fig. 15.36. Jurassic pelagic carbonates. Depositional model for Early Jurassic (Sinemurian-Carixian) deep-water limestones in the Adnet area near Salzburg (Austria). Slope setting and deeper basinal setting are an inherited Rhaetian topography consisting of a shallow reef zone and a marly basin. Jurassic sedimentation continued in parts of the Triassic basin but was interrupted during the lowermost Jurassic in areas over former reef swells. Reduced sedimentation is recorded by hardgrounds and ferruginous crusts in the Lower Adnet Formation that includes the famous red limestones and red nodular limestones, commonly summarized under the name Adnet limestone. Deeper basinal limestones are represented by gray cherty carbonates. The sedimentation was controlled by increasing plankton production triggered by enhanced upwelling. Bottom currents controlled the deposition of fine carbonate mud, the input of nutrients and the oxygenation of bottom waters. Water depths are inferred from the dip, geometry and microfacies. Width of the sketch is about 10 km. The arrows indicate characteristic microfacies types shown in Plate 139. Modified from Böhm (1992).
Plate 139: Early Jurassic Adnet Limestone
799
800
hardgrounds with ferruginous-phosphatic crusts are common. Bedding is disturbed by slides and slumps. Calcarenite and breccia intercalations as well as turbidites are common. The pelagic sequences are interpreted as being deposited at relatively shallow depths on the top of and adjacent to fault-bounded seamounts formed by drowned carbonate platforms (Colacicchi and Baldanza 1986). Similar pelagic deep-water facies and association with the shedding of shallow-water material from carbonate platforms into adjacent open-marine basins occur in the Cretaceous of the Middle East and Mexico (Wilson 1975; Enos 1977). Chalk: An exceptional pelagic sediment Chalk is a friable or hard pelagic lime mudstone containing planktonic and benthic foraminifera (see Sect. 10.2.2.1), calcispheres (see Sect. 10.2.1.9), bivalve fragments (chiefly Inoceramus, Pl. 87/8), echinoid plates, and locally bryozoans and other fossils. The fine fraction consists of up to 75% planktonic biota (coccolithophorids, calciodinoflagellates; Pl. 7/3-5), and calcite crystals representing disintegrated nannofossils and cement crystals (Fig. 15.37). Burrows and trace fossils suggest substrate stability and the availability of sufficient oxygen near the sediment-water interface. Extensive hardgrounds, glauconite and phosphatic mineralization are common. The lower parts are rich in clay and marls, and cherts. Resedimentation (indicated by common millimeter- to centimeter-sized intraclasts; Pl. 16/5) and intercalations of allochthonous carbonates are abundant in many chalk sections, e.g. in the North Sea chalk (Schatzinger et al. 1985). Most of the Cretaceous chalk of Europe, North America and the Middle East was deposited during global sea-level highstands in open shelf environments and in deep-sea settings. Microfacies data of chalks are important because of the poor visibility of sedimentological and biogenic structures in large parts of chalks. Maliva and Dickson (1992) studying the Cretaceous/Tertiary chalks of the Eldfisk Field in the Norwegian North Sea differentiated four microfacies defined by grain composition and fabric, abundance and distribution of authigenic minerals, and evidence of the mechanisms of porosity reduction: • Foraminifera mudstone characterized by volumetrically predominant planktonic foraminifera. • Radiolarian-foraminifera mudstone, differing from the foraminifera mudstone microfacies only by common molds of radiolarians that are empty or filled with calcite, pyrite or silica.
Cretaceous Chalk
Fig. 15.37. Chalk. The SEM photograph, a detail of Pl. 7/3 exhibiting a soft chalk, shows the accumulation of intact coccoliths and microcrystalline calcite particles, most representing disintegrated coccolith plates with varying degrees of calcite overgrowth. Interparticle microporosity is slightly reduced. Porosity reduction is largely due to dense packing of coccoliths and microcrystalline calcite particles resulting from consolidation (mechanical compaction) rather than from a large amount of cement crystals formed during weak compaction. Consolidation results from an increase in grain-packing density and concurrent expulsion of pore fluids. Consolidation is the main process by which pure calcite oozes, with an initial porosity of about 80%, are transformed into soft chalks with porosities between 40% and 50%. Late Cretaceous: Deep Sea Drilling Program Site 463, core 44, western equatorial Pacific.
• Extraclast-Bonetocardiella-foraminifera wackestone, characterized by planktonic foraminifera and calcispheres, and transported chalk clasts indicating debris flow deposits. • Extraclast-foraminifera packstone/wackestone characterized by foraminifera, extraclasts, sponge spicules, radiolarians, thoracosphaerid and other calcispheres, benthic foraminifera, ostracods, echinoderms, and detrital quartz grains. Many of these constituents also occur as minor components in the other microfacies. The extraclast-foraminifera packstone/ wackestone microfacies constitutes most of the turbidites within the chalks. Pelagic carbonates deposited on the ocean floor Pelagic sediments with an inferred oceanic basement have been described from various parts of the world, e.g. the Cretaceous Troodos Massif of Cyprus and the Franciscan Complex of California (see review in Jenkyns 1986), but the only reliable guide to the recognition of true oceanic sediments is an association with ophiolites. Most of the unequivocalocean-floor facies are Jurassic or Cretaceous in age. Bernoulli and Jenkyns (1974) distinguished five facies types: • Ophicalcites. Brecciated surface of the oceanic basement (serpentinite, gabbro) heavily brecciated. Fis-
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Contourites
Fig. 15.38. Strongly condensed pelagic cephalopod limestone. The red limestone exhibits most of the criteria characteristic of carbonates deposited in environments with strongly reduced or interrupted sedimentation: Stratigraphic condensation indicated by the occurrence of guide fossils revealing a tremendous time range recorded by only very thin sedimentary units, abundant hardgrounds and dissolution features, and internal brecciation and synsedimentary fissures. The sample comes from limestones occurring in close proximity with radiolarites, and volcanics. The limestone, a bioclastic wackestone with ammonites, shell debris and radiolarians is characterized by irregular layers of different thickness, separated by dissolution surfaces, and abundant Fe-Mn crusts. The primarily aragonitic shells of the ammonites (A) are extremely dissolved and filled with sediment or sparry calcite. Geopetal fabrics in the shells indicate reworking and redeposition. The lower part of the picture (1) shows a transition from dark thin-layered to light, domed parts. The latter are characterized by sparry calcite between dissolution clasts. Part 2 demonstrates the strong differences in the thickness of the irregular layers, caused by sedimentary differences in the content of non-carbonate constituents. Part 3 is an example of internal brecciation. Part 4 shows a thick Fe-Mn crust characterizing the top of a hardground. Millimeter- to centimeter-thick Mn-Fe crusts also occur on ammonite shells. Part 5 includes several relicts of dissolved ammonite shells. The nodular fabric reflects the strong dissolution of the pelmicritic sediment. The arrow points to the boundary of a vertical fissure (‘neptunian dike’, see Sect. 5.3.1) filled with dissolution relicts and ammonites. These fissures have been explained by the effect of initial rifting. Note the difference between this synsedimentary fissure and the tectonic fracture at the left of the picture. Early Middle Triassic: Asklepeeion in Epidaurus, Argolis, central Greece. After Dürkoop et al. (1986).
•
•
•
•
sures are filled by alternating white sparry calcite and red or green cryptocrystalline limestone. The matrix-rich breccias above serpentinites have been interpreted as magmatic carbonatites. Isotope data point to a sedimentary origin. Umbers: Brown iron-rich mudstones overlying mudstone (e.g. in Cyprus). Interpreted as products of precipitation from the interaction of magmatism and seawater. Radiolarites. Pillow lavas, as the uppermost part of ophiolite sections, are generally overlain by bedded radiolarites formed in a deep ocean below the calcite compensation depth. White nannofossil limestones typically overlie Jurassic radiolarites. Lithologically indistinguishable from Majolica, their continental counterpart. Black shales.
The example shown in Fig. 15.38 fits some of the criteria of pelagic carbonates potentially deposited at or near the oceanic basement (association with volcanics, radiolarites). On the other hand, strong stratigraphic condensation, the abundance of hardgrounds and of ammonites conforms to deep-marine limestones similar to those of the Hallstatt facies of the Alps formed in relatively shallow deep-water basins.
15.8.3 Contourites Contourites are deep-water drifts (more than 2000 m semipermanent bottom currents along the continents), mid-water drifts (300–2000 m on continental slopes), and outer shelf/upper slope drifts (50–300 m, formed under the influence of countercurrents and underflows).
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Recent research is focused on the question of how fossil contourites can be recognized. Many inferred fossil ‘contourites’ are now interpreted as bottom-reworked turbidites. Examples of ancient carbonate contourites have been described from the Ordovician Jiuxi Drift in Hunan, China (Duan et al. 1993), the Late Cretaceous Couches Rouges and Couches de Wang in the Swiss and French Alps (Villars 1991), the Cretaceous Talme Yafw Formation in Israel (Bein and Weiler 1976), the Paleogene Lefkara Formation in Cyprus (Kahler 1994), and the Neogene Mishaki Formation in Honshu, Japan (Stow and Faugéres 1990). These examples exhibit the following common criteria: • Location: along the slope or base-of-slope. • Prism shape of the sediment body, mound shape (several hundreds of meters in size). • Distinct contourite sequences. • Intensive bioturbation, continuous with deposition. Main contourite facies are bioturbated calcisiltite and burrow-mottled calcisiltites. • The sediment consists of calcilaminite, calcilutite, calcarenite, calcirudite and marls transported to the basin margin by downslope resedimentation of epicontinental platform carbonates. The coarse-grained facies is interpreted as being deposited by turbidity currents and debris flows; the fine-grained facies as having been dispersed along the slope by bottom currents and deposited as calcareous muddy contourites. • Contourites occur within a range of resedimented facies, pelagites, hemipelagites associated with cherts and micrites; or hemipelagites with pyroclastic sediments and concentrations of volcaniclastics. • The association with macrofossils and trace fossils is characteristic of deep-marine environments. • Calcareous contourites occur in regions of dominant pelagic biogenic input. Bedding indistinct, primary lamination may be preserved, mean grain size is most commonly silty clay or clayey. • The dominant biota are nannofossils and foraminifera. Many biogenic particles are stained and fragmented. Stow et al. (1998) suggested a three-stage approach to investigating of contourites, consisting of small scale (field, borehole or lab), medium-scale (formation and region) and large scale (system or continent) studies. Most of the questions summarized by the authors and related to the occurrence, structure, texture, fabric, structure and sequence of supposed contourites can be answered by combining microfacies data and outcrop observations.
Basics Deep-Marine Carbonates
Basics: Deep-marine pelagic carbonates Bandel, K. (1974): Deep-water limestones from the Devonian-Carboniferous of the Carnic Alps, Austria. – International Association of Sedimentologists, Special Publications, 1, 93-115 Bernoulli, D., Jenkyns, H.C. (1974): Alpine, Mediterranean and Central Atlantic Mesozoic facies in relation to the early evolution of the Tethys. – In: Dott, R.H., Shaver, R.M. (eds.): Modern and ancient geosynclinal sedimentation. – Society of Economic Paleontologists and Mineralogists, Special Publications, 19, 129-160 Böhm, F. (1992): Mikrofazies und Ablagerungsmilieu des Lias und Doggers der Nördlichen Kalkalpen. – Erlanger Geologische Abhandlungen, 121, 57-217 Cook, H.E., Enos, P. (eds., 1977): Deep-water carbonate environments. – Society of Economic Paleontologists and Mineralogists, Special Publications, 25, 336 pp. Farinacci, A. and Elmi, S. (eds., 1981): Rosso Ammonitico Symposium 1980. – 602 pp., Roma (Tecnoscienza) Garrison, R.E. (1972): Inter- and intrapillow limestones of the Olympic Peninsula, Washington. – Journal of Geology, 80, 310-322 Garrison, R.E., Fischer, A.G. (1969): Deep-water limestones and radiolarites of the Alpine Jurassic. – In: Friedman, G.M. (ed.): Depositional environments in carbonate rocks. A symposium. – Society of Economic Paleontologists and Mineralogists, Special Publications, 14, 20-56 Jenkyns, H.C. (1986): Pelagic environments. – In: Reading, H.G. (ed.): Sedimentary environments and facies. – 343390, Oxford (Blackwell) Leggett, J.K. (1985): Deep-sea pelagic sediments and palaeocenography: a review of recent processes. – In: Brenchley, P.J., Wulliams, B.O.J. (eds.): Sedimentology: Recent developments and developments and applied aspects. – Geological Society of London, Special Publication, 18, 95121 Santantonio, M. (1993); Facies associations and evolution of pelagic carbonate platform/basin systems: examples from the Italian Jurassic. – Sedimentology, 40, 1039-1068 Scholle, P.J., Arthur, M.A., Ekdale, A.A. (1983): Pelagic environments. – In: Scholle, P.A., Bebout, D.G., Moore, C.G. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists, Memoir, 33, 620-691 Stow, D.A.V., Faugères, J.-C., Viana, A., Gonthier, E. (1998): Fossil contourites: a critical review. – Sedimentary Geology, 115, 3-31 Stow, D.A.V. (ed., 1991): Deep-water turbidite systems. – IAS Reprint Series, 3, 488 pp. Tucker, M. (1974): Sedimentology of Palaeozoic pelagic limestones: the Devonian Griotte (Southern France) and Cephalopodenkalk (Germany). – International Association of Sedimentologists, Special Publications, 1, 7192 Wendt, J., Aigner, T. (1985): Facies patterns and depositional environments of Palaeozoic cephalopod limestones. – Sedimentary Geology, 44, 263-300 Wilson, J.L. (1969): Microfacies and sedimentary structures in ‘deeper water’ lime mudstones. – In: Friedman G.M. (ed.): Depositional environments in carbonate rocks. A symposium. – Society of Economic Paleontologists and Mineralogists, Special Publications, 14, 4-19 Further reading: K022, K181; see also Basics: Platform margins and carbonate slopes.
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16 Realizing Depositional Constraints and Processes
Having learned how microfacies can be integrated into facies models (Chap. 14) and how microfacies can be translated into depositional settings (Chap. 15), the following chapter summarizes the potential of microfacies for recognizing and interpreting major depositional constraints. These constraints are manifested by characteristic sedimentation patterns and require the application of specific facies analysis techniques. Sect. 16.1 deals with cyclic carbonates and underlines the benefit of microfacies for sequence stratigraphy. Sect. 16.2 discusses methods applied to reef carbonates and stresses the necessity of combining microfacies data and ecological characteristics. Sect. 16.3 explains how painstaking microfacies studies assist in reconstructing vanished carbonate platforms that are only recorded by pebbles and boulders. Sect. 16.4 includes case studies of ancient cold-water carbonates attesting to the significance of microfacies criteria. Sect. 16.5 treats ancient carbonates interpreted as products of marine seeps and vents. Sect. 16.6 gives a short summary of the facies criteria of mixed carbonate-siliciclastic paleoenvironments and discusses limestone-marl sequences. Many of the constraints on carbonate deposition as well as the biological contribution to the formation of carbonate rocks were subject to secular changes during the Earth’s history. Some of these changes, reflected by the occurrence, frequency and distribution of microfacies criteria, are the topic of Sect. 16.7.
16.1 Cyclic Carbonates, Microfacies and Sequence Stratigraphy The objective of this section is to document the value of microfacies in the study of cyclic carbonates within the framework of sequence analysis. Carbonate sequence stratigraphy plays a major role in reservoir char-
acterization and simulation. Intensive analysis of outcrop data, detailed facies studies of carbonate cycles, and computer-based forward models have led to the development of 3-D models describing depositional and diagenetic patterns of carbonate rocks and predicting carbonate reservoirs. Both parts of the section start with a brief overview of basic principles, continue with the discussion of relevant microfacies data, and close with some examples illustrating the advantage of thin-section studies for the understanding of cyclic carbonates and for sequencestratigraphy interpretations.
16.1.1 Cyclic Carbonates A cycle is a group of rock units that occur in a certain order, with one unit being frequently repeated throughout the succession. A stratification cycle is a group of beds which are regularly repeated (Schwarzacher 1987). Regularity may be recorded by bed composition, repeated sequences of bed thicknesses, or a constant number of beds composing the cycles. Cyclicity is a common feature of many limestone successions and occurs across different carbonate environments from platforms to reefs and down to basins. Understanding the origin of cyclicity in platform carbonates is crucial if cycles are to provide information on climate, eustacy and local tectonics, and if they are to be useful for high precision correlation. Many cycles are ascribed to corresponding, high-frequency and low amplitude sea-level changes. Cyclostratigraphy based on orbital cycles is a well-established approach in basin analysis (Fischer et al. 1990; House et al. 1995; Schwarzacher 2000). Pioneering studies of sedimentary cycles were undertaken in the Pennsylvanian and Permian ‘cyclothems’ of North America, followed by detailed and sophisticated investigations of Triassic platform carbonates in the Alps, and Cretaceous pelagic carbonates in different parts of the world. The term cyclothem, originally suggested for specific Late Paleozoic sedimen-
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_16, © Springer-Verlag Berlin Heidelberg 2010
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tary successions, is now commonly used for the sedimentary record of cyclic processes. Not included in the present section are rhythmic limestone-marl successions that are discussed in Sect. 16.6 dealing with sediments formed in mixed carbonate-siliciclastic settings.
16.1.1.1 Cyclic Carbonates: Some Basics Cyclic sedimentation produces vertical successions of sedimentary strata characterized by specific patterns. Based on the mechanisms that generate cyclic deposits, two types of cyclic successions can be distinguished: • Autocycles controlled by processes that take place within the basin itself, e.g. tides or storms. Autocyclic successions show limited stratigraphic continuity. • Allocycles caused mainly by variations external to the depositional basin, e.g. sea-level fluctuations, climatic changes or tectonics. Allocyclic successions may extend over long distances. Autocycles: Oscillations generated within the carbonate system are used in the Ginsburg model, that explains the common shoaling-upward cycles of carbonate platforms capped by fine-grained tidal flat deposits (Ginsburg 1971; Hardie and Shinn 1986). The model assumes continuous, steady basinal subsidence coupled with carbonate sediment supply rates that are self-regulated by the extent of the progradation. High sediment production in shallow lagoons leads to seaward progradation of tidal flat belts. This in turn reduces carbonate production and eventually ceases progradation. Extensive supratidal flats with nearly zero production lie adjacent to a narrow lagoon. The next cycle starts with an abrupt transgression and flooding of the tidal flats when subsidence has sufficiently lowered the beach ridges on the seaward side of the flats. Autocyclicity, related to changing sedimentation during tidal flat progradation, is favored by Hardie (1986), Pratt and James (1986), Bosellini and Hardie (1988), Cloyd et al. (1990), Anderson and Goodwin (1990), Strasser (1991) and Satterley (1996). Allocycles. Interpretations of cyclic carbonates relying on allocyclicity assume the existence of high-frequency, and low-amplitude, fourth- and fifth-order sealevel fluctuations that are often explained by changes in the Earth’s orbital parameters (Goldhammer et al. 1987; Strasser 1988; Koerschner and Read 1989; Goldhammer et al. 1990; Read et al. 1991). Orbital forcing and Milankovitch cycles. Insolation variations affecting the Earth’s climate, glacial volume
Cyclic Carbonates; Basics
and ocean volume have been advocated to explain corresponding cyclic patterns in climate, sea level, geochemical variables and sedimentary cycles. Many records of cyclic changes have frequencies that fit well with those of orbital parameters, and hence are likely causally related to them. The Earth’s orbit varies in a cyclic fashion in several ways. The axis about which the Earth rotates precesses with a fundamental period of about 26,000 years and modes between ca. 19,000 and 23,000 years (precession). It undergoes change in axis inclination (obliquity) in an oscillary manner, generating a cycle of about 41,000 years duration. The orbit of the Earth around the Sun changes from an almost circular to an almost elliptical (eccentricity) orbit in two cycles, one at a frequency of about 100,000 years, and the other 400,000 years. These variations in Earth’s orbital behavior produce periodic variations in insolation and climate, called Milankovitch cycles with periodicities ranging between about 20,000 and more than 400,000 years. The correlation of the Pleistocene deep-sea and coral-reef record with Milankovitch climatic rhythms (Hays et al. 1976; Berger 1984; Imbrie 1985) has resulted in a widespread acceptance of the Milankovitch theory for Pleistocene cycles and also for pre-Pleistocene cyclic sediments. Milankovitch climatic rhythms are recorded in lake sediments and evaporitic deposits, in shallow-marine carbonates as well as in hemipelagic and pelagic sediments. Orbital forcing and resulting sea-level fluctuations are regarded by many authors as the major controls on carbonate platform cycles. Individual shallowing-up (sometimes deepening-up) cycles are interpreted as the result of fifth-order and higher-order sea-level oscillations. Groups of fifth-order cycles form fourth-order sedimentary cycles. Third-order cycles in shallow-marine carbonates have been defined by systematic thickening and thinning of individual cycles over several hundred meters of profiles. Important studies on cyclic carbonates using the Milankovitch model include Fischer (1964), Goldhammer et al. (1990, 1991) and Satterley (1996a and 1996b). Problems with the application of the orbital forcing theory to carbonate cycles were discussed by Algeo and Wilkinson (1988), Drummond and Wilkinson (1993), Satterley (1996b) and Schlager (2002). Orders of stratigraphic cycles. Cycles are hierarchically classified by the duration of sea-level changes whereby rates of eustatic changes and amplitudes reflect the generating processes (glacio-eustasy or tectono-eustasy). Typical high-frequency, meter-scale cycles including fifth- and fourth-order cycles with durations of a few ten thousand years (for fifth-order)
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Recognizing Cyclicity
and a few hundred thousand years (for fourth-order) are grouped into third-order cycles formed on a scale of about 1 to 10 million years. Second-order cycles have durations ranging from 10 to 100 Ma and first-order cycles last several hundreds of million years. First- and second-order cycles attributed to longterm sea-level changes and tectonics (breakup of supercontinents, changes in spreading rates and in the flexural rigidity of the lithosphere) will not control the internal architecture of carbonate sequences, because the rates of the amplitudes of sea-level fluctuation (1– 2 cm/1000 years) are two orders of magnitude less than carbonate sedimentation rates (10–1000 cm/1000 years). The formation of sedimentary cycles in the third, fourth and fifth order bands have been explained by Milankovitch cycles, but also by autocyclic factors and tectonic controls. Causes of third-order cycles are the waxing and waning of polar ice caps (glacio-eustasy) and changes in the volume of ocean basins (tectonoeustasy). Shallow-marine carbonate platforms and ramps as well as pelagic carbonates record third-order, fourth-order and fifth-order eustatic cycles and depositional sequences (see Sect. 16.1.2.1).
Recognition of cyclicity: Cycles and sequences can be recognized by changes within the sedimentary succession expressed by: • Lithology (alternation of limestone and dolomite, or limestone and marls; insoluble residue; see Sect. 13.1.1.1). • Bedding (recurrent changes in bed thickness, bundling or composition). Sedimentary rhythms could be reflected by couplets of strata or groups of couplets. • Facies including depositional and diagenetic microfacies (see the following text). • Fossils (distribution, association, frequency). Ichnofossils and burrowing fabrics in pelagic carbonates are particularly useful in determining sea-level positions (Savrda and Bottjer 1994). • Geochemistry including fluctuations in stable isotopes (Sect. 13.2.2), major and minor elements (Sect. 13.2.1), and content of organic matter (Sect. 13.3). Carbon as well as oxygen isotopes are used to decipher the cyclicity of shallow-water micrites, but carbon values may offer more reliable data than oxygen values, which are more strongly affected by diagenesis. Distinct differences in the isotope data of the micritic matrix of tidal and subtidal limestones reflect cyclic changes in environmental conditions or different diagenetic controls (Joachimski 1991; Jimenez de Cisneros and Vera 1993).
• Statistical methods: Cyclicity in carbonates may not be obvious when looking at the changes in lithology, texture or biotic composition. In order to recognize hidden cyclicity patterns statistical methods are used. These methods include correlation analysis, principle component analysis and spectral analysis looking at time series in terms of their frequency composition (e.g. Pelosi and Raspini 1993; Gischler et al. 1994). Traditional time-series analysis assumes that the cyclic record is complete, so that each recorded stratigraphic rhythm corresponds to one pulse of the cycle-producing mechanism. The main assumptions are equal periodicity per cycle and no gaps in the data set. Power spectral analysis is applied to look for cyclicity and detect orbital-climatic forcing (see papers in De Boer and Smith 1994). These methods are applied to platform and ramp carbonates as well as pelagic carbonates. The basic data used comprise thickness, stacking and composition of strata as well as the distribution of paleontological or geochemical data. Cycle-thickness data can be analyzed for its directionality or randomness (Drummond and Wilkinson 1993; Grötsch 1996). Regarding the sedimentation rate as constant and periodicity as variable, Schwarzacher and Fischer (1982) have calculated periodic cyclicity from a statistical analysis of the thickness of the sequences. Spectral power spectra based on a fast Fourier transformation of cycle thickness to frequency (cycles per 1000 years) indicate whether the sedimentary cyclicity is within the Milankovitch spectrum or not (Schwarzacher 1991) and aid in the discussion of the recognized cyclicity in terms of eccentricity, obliquity and precession cycles (Goldhammer et al. 1987, 1990; Osleger 1991). The techniques of time series analysis are described in Einsele et al. (1991), Sprenger and Ten Kate (1993) and Schwarzacher (1993).
16.1.1.2 Microfacies and Cyclic Carbonates Variations in microfacies criteria within a sedimentary succession (e.g. a cycle or a parasequence) indicate changes in environmental controls and depositional rates. Changes in sea-level position (e.g. shallowingupward) are connected with shifts in depositional water depths, which in turn can be associated with changes in hydrodynamic conditions (e.g. low-energy to highenergy). Changes in depositional rates influenced by sea-level fluctuations may lead to the formation of boundaries between depositional units (sequences) pointing to significant alterations and/or breaks in sedimentation. These breaks are expressed by submarine discontinuities and unconformities, and by subaerial
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exposure surfaces. Discontinuity surfaces and exposures surfaces are recorded by characteristic microfacies criteria. Definitions of these criteria have been discussed in Chap. 4 and 5.
Shallowing-upward cycles of carbonate platforms Many Precambrian and Phanerozoic carbonate platforms are well stratified. The stratification is caused by stacking of meter-scale depositional sequences displaying a shallowing-upward trend. The shallowingupward sequences consist of repetitions of subtidal facies bounded by peritidal facies (peritidal cycles; e.g. Lofer cycle, Sect. 16.1.1.3), subaerial exposure surfaces (diagenetic cycles), and/or marine flooding surfaces (subtidal depositional cycles). More or less regular repetitions of sequences in platform and peritidal carbonates imply that cyclic processes (e.g. periodic changes in flooding through transgressions) controlled carbonate production and deposition. Diagenetic cycles originate in shallow and in deeper marine environments. In shallow environments, they consist of subtidal wackestones and grainstones, overlain by a thin cap of vadose cements, pisoids and various evidence of calichification (e.g. Latemar cycle, Sect. 16.1.1.3). Subtidal depositional cycles are characterized by a coarsening-upward trend in high-energy structures such as grain size, bed thickness and cross-bedding, and the lack of intertidal sediments or subaerial exposure. These cycles are a dominant cycle type on unrimmed shelves and ramps (Calvet and Tucker 1988; Osleger 1991). In deeper open shelf environments, cycles capped by hardgrounds can originate from a change in water depths and reduced sedimentation. Sediment accumulation takes place when the shelf was below the depth of wave abrasion, cementation at times when the shelf was above wave base and subject to water pumping. An example of this are Tertiary temperate limestones of southern Australia (Bone 1991). Meter-scale cycles dominated many shallow-water successions throughout the history of carbonate rocks. These high-frequency depositional cycles are fundamental sequence-stratigraphic units of carbonate platforms (parasequences; Sect. 16.1.2.1) and basic elements in the interpretation of transgressive-regressive processes. Peritidal shallowing-upward successions. Pratt et al. (1992) distinguished two types: • The low-energy type is characterized by a subtidal, burrowed, variably argillaceous lowest member. The
Shallowing-Upward Cycles
member consists of mudstone, wackestone or packstone, often with a basal bio- or intraclastic grainstone as a transgressive lag on top of the preexisting succession. The lower intertidal member exhibits thin-bedded, variably bioturbated mudstones, and bioclastic, peloidal and sometimes oolitic grainstones, locally with stromatolites. The upper intertidal and the supratidal member is usually a microbially laminated argillaceous mudstone/bindstone with fenestral fabrics and thin interbeds of intraclastic packstones and laminae consisting of peloidal or bioclastic grainstone. In arid climates, the inter- and supratidal members may be represented by nodular to wavy beds of anhydrite. • The high-energy type also exhibits subtidal bioturbated limestones, predominantly bio- and intraclastic grainstones, at the base, but the lower intertidal member consists of cross-bedded and cross-laminated bioclastic and/or locally oolitic grainstones, representing beach deposits. The upper intertidal and the supratidal members are characterized by desiccation-cracked microbial laminites. Both subtypes may exhibit a pedogenic capping horizon (paleosol, calcrete); the units may be separated by subaerial karst surfaces or erosion surfaces formed during environmental shifts. Common trends pointing to shallowing-upward in peritidal cycles starting with shallow subtidal sediments at the base are • upward decrease in grain-supported depositional textures, in grains formed in high-energy environments (e.g. ooids), and in normal marine biota (suggesting increased restriction of environmental conditions), • increase in the evidence of meteoric-vadose diagenesis, in the abundance of admixtures of terrigenous quartz and silt, and of extraclasts. Microfacies of shallowing-upward cycles. Ideal shallowing-upward cycles consist of distinct microfacies exhibiting criteria of shallow subtidal, intertidal and/ or supratidal deposition and evidence of subaerial exposure (e.g. desiccation cracks, pedogenic structures, pisoids, karst criteria). The diagnostic criteria of these depositional environments have been discussed in Sect. 15.5.1.2 and are summarized in Box 15.1. Microfacies criteria pointing to shallowing-upward conditions are mud peloids, oncoids, rhodoids and the types of fenestral fabrics (Sect. 5.1.5) as well as increasing meteoric-vadose cementation. Changes in the frequency of mud peloids within bedded micritric limestones indicate fluctuations in water energy and erosion of the sea bottom and can reflect cyclic shallowing patterns (Carss and Carozzi 1965; see Sect. 4.2.2).
Cyclicity in Reefs
Cyanoids occurring in different sedimentary environments are characterized by different criteria that can be used in recognizing changes in sea level. Shallowing is recorded by stromatolitic crusts overgrowing the oncoids, changes in the size, shape and composition of oncoids, and differences in depositional textures (Ratcliffe 1988). Rhodoids formed in shallow and deep waters and under different hydrodynamic conditions differ in composition, shape and size (Sect. 4.2.4.2). These differences can be used for estimating shallowing- and deepening-upward patterns. Burrow systems and taphofacies are sensitive indicators of shallowing and deepening both in shallow and deep environments (Brett 1995; Monaco and Giannetti 2002).
Cyclicity in reefs Sea-level fluctuations are major controls on the growth of shallow-water reefs. Reef-building organisms respond to shallowing and deepening of the reef top with changes in growth forms and guilds, succession patterns and the rate of carbonate production. Cyclic sea-level fluctuations can be recognized in ancient reefs using reef guilds and reef diagenesis, as demonstrated by Grötsch (1994) and Grötsch et al. (1994) studying an Early Cretaceous reef in Slovenia formed by corals and various encrusting organisms. Methods: The section studied in Slovenia comprises 133 m massive coarse-grained reefal limestones and fine-grained lagoonal limestones. Samples were taken with an equal-distance spacing of 0.5 m. A peel with an area of 16 cm2 was prepared from each sample. In order to set up a semiquantitative index to trace carbonate production per time unit, different environmentsensitive variables have been considered: • Relative abundance of reef guilds (constructor, binder, baffler, dweller, destroyer; see Sect 16.2.3.2 for discussion of the guild concept). • Primary content of biogenic aragonite comprising completely or partly aragonitic and now recrystallized organisms. • Textural types according to Dunham (1962). These qualitative facies variables were transformed into ‘facies numbers’ that increase with increasing sparry cement and decreasing micrite content, increasing grain size and increasing water energy. Results: From thin section analysis three facies types can be distinguished: (1) In-place and detrital framework facies (grainstones, rudstones, boundstones) characterized by abundant reef-building organisms pre-
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served as detritus and aragonitic skeletons; (2) detrital sand facies (grainstones) with abundant coated reef debris; and (3) baffler facies (wackestones, packstones) distinguished from the other facies types by the high amount of micrite and small amount of sediment-baffling algae. The multivariate data set was reduced by principal component analysis. Using component loadings as a relative index of framework production, vertical aggradation reef cycles controlled by regional subsidence rate, a third-order change in eustatic sea level, and superimposed high-frequency fluctuations in the Milankovitch band were recognized. Sea-level fluctuations were also documented by the changing abundance of marine and meteoric carbonate cement types. The boundaries of depositional cycles and diagenetic cementation cycles do not coincide because cementation lags behind sedimentation and early diagenetic meteoric overprint may extend several meters down into the sediment.
Shallow-marine and deep-marine cyclicity Relationships between cyclic sedimentation on carbonate platforms and cyclicity in basinal carbonates are reflected by • prograding bioclastic wedges of depositional sequences in outer platform settings correlating with packages of basinal limestone beds and bioclastic clinoforms (e.g. Quesne and Ferry 1994); • correlation of shallow-marine parasequence patterns with those of basinal carbonate cycles derived from compositional and geochemical data (Elder et al. 1994); • composition of calciturbidites (e.g. Reijmer et al. 1994; Sect. 15.7.5.2). Microfacies analysis indicates where the material has come from and whether a cyclic pattern exists. Answering the last questions requires statistical treatment of frequency data using spectral analysis and time analysis; • compositional analyses of allochthonous sediment intercalated within pelagic and hemipelagic beds, caused by highstand shedding (Sect. 16.1.2.2).
Pelagic cyclicity Recognizing and interpreting pelagic carbonate cycles with respect to climatic variations and sea-level fluctuations demands a combination of sedimentological, geochemical and paleontological data and the evaluation of these data by statistical analyses, e.g. spectral analysis. Commonly used basic criteria are strati-
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fication, number and thickness of beds, lithological variations (e.g. limestone-marl alternations), carbonate and non-carbonate contents, stable isotopes, main and minor elements, and rock color (light and dark beds) reflecting organic carbon content. Sampling within centimeter-scale and the study of planktonic microfossils both in washed samples and in microfacies thin sections allow biogenic carbonate productivity to be estimated (e.g. Noé 1993). Important papers dealing with pelagic cyclic carbonates are listed in Box 16.1. Late Paleozoic cyclothems The term cyclothem designates the rock record of cyclic accommodation processes characterized by coarsening-upward successions, reflecting a passage from marine sedimentation upwards into terrestrial sedimentation. The cyclothem concept was developed in the Pennsylvanian of the United States and generalized by Weller in Wanless and Weller (1932) and Moore (1931). At present, the most widely accepted cyclothem model is that of the Kansas-cycle cyclothems (Heckel 1977). Midcontinent cyclothems are dominated by very shallow and paralic facies with abundant subaerial exposure surfaces including well-developed paleosols. Meter-scale shallowing-upward cycles (parasequences) are common; they are usually capped by subaerial exposure surfaces. Interpretations of Pennsylvanian cycles include glacio-eustatic control as well as autocyclic and tectonic controls. Major global sea-level fluctuation documented by transgressive-regressive depositional patterns and having about 1–3 Ma lasting (3rd-order cycles) have been recorded from many Late Paleozoic depositional successions in North America and Europe (e.g. Russia) and were summarized in sea-level charts (Ross and Ross 1987, 1988, 1994; Heckel 1986, 1994). Permian sea-level fluctuation patterns are closely related to biostratigraphical changes in fossil zonal assemblages (e.g. fusulinid zonation). The fossil record is correlated with deposition during sea-level highstands. Limestone beds are key marker horizons in these cyclothems. Their interpretation as transgressive or regressive deposits is crucial in the analysis of larger cyclothems. Microfacies offers key data for evaluating the paleoenvironment and depositional depths of these carbonates (significant references are listed in Box 16.1, Carboniferous and Permian). Cyclothem boundaries may correspond to prominent marine flooding surfaces overlain by fossiliferous limestones. Centimeter-scaled
Lofer and Latemar Cycles
intraclastic beds with mudstone and limestone clasts represent lag deposits formed at transgressive surfaces truncating underlying paleosols (Miller and West 1998). The microfacies of the overlying limestones is characterized by fossiliferous wackestones and packstones. An example is shown in Pl. 56/8 and Pl. 118/1.
16.1.1.3 Case Studies: The Lofer Cycle and the Latemar Cycle (Triassic of the Alps) Shallowing-upward successions exhibit different microfacies patterns. Two patterns represented by Triassic carbonates have become key elements in the discussion of peritidal and diagenetic cycles – the Lofer or Dachstein cycle of the Austrian Northern Calcareous Alps and the Latemar cycle of the Dolomites in South Tyrol in Northern Italy. Microfacies and the controversial interpretations of these cycles are the topics of the following paragraph.
The Lofer (Dachstein) cycle The cyclic depositional pattern of Alpine Late Triassic carbonates was recognized by Sander (1936) in the Loferer Steinberge near Lofer, and the Steinernes Meer, Salzburg, and later investigated by Schwarzacher (1954) and, in great detail, by A.G. Fischer (1964), who published a fundamental study on the Lofer cyclothems. This study had a deep impact on the interpretation of cyclic carbonate sequences. The cycles of the Late Triassic Dachstein limestone are superbly exposed in huge mountain cliffs and can be studied over a thickness of several hundreds of meters (in the Lofer region about 600 m). Stratal patterns and controlling mechanisms of the Lofer cycles, however, are still controversial (Goldhammer et al. 1990; Satterley and Brandner 1995; Enos and Samankassou 1998). Comparable Late Triassic stacked high-frequency cycles were analyzed in the Southern Alps (Bosellini and Hardie 1988), Slovenia (Ogorelec and Rothe 1992), Hungary (Haas 1991, 1994; Balog et al. 1997), Sicily (Catalano et al. 1974) and Oman (Bernecker 1996; Weidlich and Bernecker 2003; Fig. 10.40). Lofer-type cycle patterns are not limited to the Late Triassic as shown by Liassic successions in Slovenia (Dozet 1993). Fig. 16.1 shows an idealized Lofer cyclothem illustrated by A.G. Fischer (1964) and appearing in many textbooks and overview articles. Fischer distinguished an ABC sequence consisting of three members:
Lofer Cycle
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present (Pl. 140/1). Member A is usually regarded as a supratidal deposit. • Member B is darker in color, predominantly dolomitic and characterized by distinct lamination structure (Pl. 140/3; Fig. 16.2C), fenestral fabrics including shrinkage pores (Pl. 140/2), and intraclasts. The laminites are very often crossed by vertical cracks (prism cracks, attributed to desiccation). This subtype corresponds to the ‘algal mat loferite’ described by Fischer (1964) and to subfacies B1 differentiated by Goldhammer et al. (1990). Member B also contains layers of burrowed, peloidal wackestones and packstones with restricted fauna consisting of thin-shelled gastropods and ostracods (‘pellet and homogeneous loferites’; subfacies B2). Intercalations of grainstones are common (Fig. 16.2A, B). Facies B is usually regarded as an intertidal deposit. • Member C makes up to 85% of the sequence. The facies comprises <1 m to more than 25 m thick, light gray massive limestone beds and includes wackestones and grainstones with submarine hardgrounds. The limestone always contains a large amount of organic debris. Oncoids formed by calcimicrobes are fairly common. Thick-shelled bivalves (Megalodon, Conchodon) are conspicuous (Pl. 140/4). The bivalves often form layers and occur in life position. Other common fossils are gastropods, dasyclad algae, and, approaching the margin of the platform, corals (Pl. 140/5). Member C is usually regarded as a subtidal deposit. Fig. 16.1. Diagrammatic representation of a Lofer cycle. The drawing is by A.G. Fischer (1964), the text is slightly modified. Following Fischer, an ideal cycle consists of three depositional units (members or facies) formed under different environmental conditions. Member A on the top of member C represents a basal disconformity, subaerial exposure and supratidal conditions. Member B records tidal conditions. Member C documents shallow-subtidal deposition. The typical ‘loferite’ facies characterized by a carbonate sediment riddled by shrinkage pores (sometimes synonymous with ‘birdseyes limestone’) occurs within Member B.
• Member A is an unconformity followed by a reddish and reddish-greenish or gray argillaceous, generally thin horizon of reworking with intraformational mud pebble conglomerates (Pl. 140/2; Fig. 16.2D) and diagenetic breccias. Associated with facies A are cracks and solution cavities in the underlying Member C. The latter is often filled with argillaceous limestone of Member A. Some of the solution cavities are aligned parallel to the bedding planes and could represent freshwater tables formed during the emersion of Member A. The red matrix is often interpreted as a soil related to subaerial exposure. Pedogenic structures may be
The repeated alternation of subtidal and intertidal/ supratidal deposits was explained by sea-level oscillations. Fischer (1964) favored a transgressive pattern of the ‘Lofer cycle’ in which basal intertidal/supratidal laminites (Member B) pass upward into shallow subtidal units (Member C) recording sea-level rise and a transgression of the shoreline. Soils capping the subtidal member mark a fall in sea-level. This deepeningupward interpretation has been opposed by Goldhammer et al. (1990) and Satterley (1996) who reinterpreted the classical Lofer cycles as shallowing-upward cycles, and Enos and Samankassou (1998), who stress the possible influence of autocyclic processes and interpret the pattern recorded by the bedded Triassic Dachstein limestone as alternating rhythms (...BCBCB...) rather than regular cycles. The crucial point in the different interpretations is the origin of the red ‘soils’ of Member A and the stratigraphic position of this member. The microfacies of the members forming Lofer cyclothems is highly variable. Common microfacies types include:
Lofer Cyclothem
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• Member A: Marl, silty marl, dolomitic marl; argillaceous mudstone, unfossiliferous lime mudstone (Pl. 140/1); lime mudstone with quartz grains; peloid wackestone; wackestone with sparse skeletal grains (e.g. ostracods); intraclast wackestone and floatstone with clasts consisting of reworked microbial mats; poorly sorted fine-grained breccia with black pebbles and other clasts; intraclast wackestone and floatstone with reworked material from Member C (Pl. 140/2; Fig. 16.2D). In addition, variegated-colored micrites within solution cavities and fine-bedded micrites are common features of member A. • Member B: Unfossiliferous lime mudstone; fenestral mudstone/bindstone (Pl. 17/3; Pl. 140/2) including micritic, peloidal (Fig. 16.2C) and intrapelmicritic loferites; laminated mudstone/bindstone, stromatolitic
mudstone/bindstone (Pl. 46/2; Pl. 140/3); peloid packstone; intraclast wackestone and floatstone with clasts consisting of reworked microbial mats; fine-grained polymict breccias containing material both of Member B and Member A; peloidal, intraclastic and bioclastic grainstone (shedding from subtidal areas; Fig. 16.2A, B). • Member C: Lime mudstone with solution cavities; mudstone and wackestone with sparse tiny shell debris; mudstone and wackestone with benthic foraminifera; peloid wackestone; bioclastic wackestone and packstone with low-diversity dasyclad algae and small gastropods; bioclastic wackestone and floatstone with bivalves (Pl. 140/4), in places shell coquinas; bioclastic wackestone and bafflestone with corals (Pl. 140/5); oncoid wackestone and floatstone with cyanoids formed
Plate 140 Cyclic Carbonates: The ‘Lofer Cyclothem’ of the Alpine Triassic The Late Triassic Dachstein Formation of the Northern Calcareous Alps in Austria comprises an up to 1,500 m thick sequence of bedded limestones, deposited within a several kilometer wide lagoonal belt behind a shelfmargin reef zone (see Fig. 15.29). The Dachstein facies is widely used as a reference example of meter-scale cycles (Lofer cycle named after the Loferer Steinberge) produced by high-frequency sea-level oscillations. The cycles form meter- to decameter-thick successions. The ideal Lofer cycle (Fig. 16.1) as defined by Fischer (1964) comprises a subtidal Member C with normal marine biota, an intertidal Member B, and a supratidal Member A. Both the intertidal and the supratidal members are characterized by fenestral laminites and stromatolites. The supratidal member A may or may not have a terrestrial horizon on top. Contacts between the members are gradational or sharp. Lithologies and microfacies reflect a broad spectrum of rapidly changing environmental conditions. 1 Argillaceous lime mudstone. Member A. The micritic and pelmicritic matrix contains spar-filled voids some of which are infilled with (red) calcisiltite (1). Large horizontal and vertical voids (2) and micritic and pelmicritic areas bounded by crinkled spar-filled cracks indicate pedogenic overprint. Late Triassic (Dachstein limestone, Norian): Breithorn, Steinernes Meer, Salzburg, Austria. 2 Limestone conglomerate (1) overlain by fenestral mudstone (2 to 4). The conglomerate consists of subangular and rounded micrite clasts embedded within a calcisiltite matrix. This floatstone texture corresponds to the ‘intraformational conglomerate’ attributed by Fischer (1964) to the Member A and regarded as an indication of supratidal reworking. The lower part of the fenestral mudstone (2) contains limestone clasts, too. This zone is overlain by characteristic loferites (3) with aligned spar-filled voids, assigned to Member B and considered to be intertidal in origin. Larger geopetally filled cavities are interpreted as shrinkage pores (4) within this zone and overlying it. Same locality as –> 1. 3 Laminated loferite consisting of flat and undulated micritic and sparry layers. The densely set micritic layers (1) consist of peloids forming a packstone fabric. The more widely set sparry layers (3) exhibit porostromate calcimicrobial fabrics in some areas. The center of the picture (2) shows stacked, spar-filled structures, probably tepees (arrows). The samples corresponds to the crinkled ‘algal mat loferite’ distinguished by Fischer (1964) and the ‘laminated loferite’ B1 described by Goldhammer et al. (1992). Member B. Same locality as –> 1. 4 Bivalve floatstone with recrystallized megalodontid shells (compare Fig. 10.40). The matrix is a fine-grained packstone with variously sized peloids, aggregate grains and a few benthic foraminifera. Small white sections are ostracod and gastropod shells. The fine-grained matrix indicates a protected platform interior environment. Member C. Subtidal part of the Lofer cycle. Same locality as –> 1. 5 Coral floatstone. Dendroid corals (Retiophyllia) embedded within a bioclastic packstone matrix containing abundant fine-grained shell and coral debris. Note the difference in the composition of the matrix as compared with –> 4, indicating the existence of small coral patches in the outer zone of the platform. Member C. Subtidal part of the Lofer cycle. Late Triassic (bedded Dachstein limestone, Norian): North of Riemannhaus, Steinernes Meer, Salzburg, Austria. 1–5: Thin sections courtesy of E. Samankassou (Fribourg).
Plate 140: Cyclic Carbonates: The ‘Lofer Cyclothem’
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The Latemar Cycle
Fig. 16.2. Dachstein (Lofer) cycle. A: Laminated mudstone of the intertidal Member B. The lower part of the picture shows a peloidal packstone layer (pellet loferite). Sand-sized particles forming a grainstone texture were deposited within pockets of the loferite layer. The grainstone is separated from an upper mudstone layer by spar-filled sheet cracks, peripherally filled with recrystallized radiaxial-fibrous cement. B: Detail of –> A showing the composition of the grainstone layer. Grains comprise peloids, micrite clasts, coated grains and fossils. The latter are represented by different types of benthic foraminifera, gastropods, and echinid spines. The grain association and the erosive lower boundary of the grainstone layer point to the deposition of subtidal sediment on tidal flats under the influence of sporadic high-energy events. C: Some of the micritic and pelmicritic layers forming the laminated mudstone of Member B are fragmented, possibly indicating desiccation of mud layers. The voids are infilled with crystal silt interpreted to be of vadose origin. D: Detail of a conglomerate floatstone shown in Pl. 140/2, developed at the top of the subtidal Member C and included within the supratidal Member A. The clasts are packstones with small shell debris similar to Pl. 140/4 possibly indicating reworking of subtidal carbonates of Member C. The matrix is microsparite. A, B, C: Late Triassic (Dachstein limestone, Norian): Pass Lueg, Salzburg, Austria. D: Late Triassic (Dachstein limestone, Norian): Steinernes Meer, Salzburg, Austria. Scale is 2 mm.
by porostromate and/or spongiostromate cyanobacteria and algae; wackestone and packstone with oncoids and small ooids; poorly sorted ooid wackestone and grainstone; ooid grainstone; intraclast wackestone and packstone; intraclast grainstone; bioclastic grainstone with diverse fauna (e.g. bivalve shells, brachiopods, echinoderms).
The Latemar cycle The Latemar cycles studied in the Southern Alps differ from the peritidal Dachstein cycles because they formed in subtidal shelf settings beyond the reach of tidal flats or shoreline sand. The cycles comprise depositional cycles with no depositional or erosional breaks,
The Latemar Cycle
and diagenetic cycles having a cap typified by subaerial or submarine diagenetic features. Controls discussed for the generation of repeated meter-scale subtidal Latemar cycles include autocyclic processes (vertical accretion, variations in sediment production and redistribution), allocyclic processes (episodic subsidence, short-term eustasy), or a combina-
Fig. 16.3. An ideal Middle Triassic ‘Latemar cycle’ of the Dolomites consists of an internally differentiated subtidal unit (1–3) and a thin subaerial exposure cap (4). The base of the submarine unit is defined by a sharp bedding plane, usually with a concentration of dolomite clasts derived from the cap of the underlying couplet. The submarine unit exhibits an upward coarsening trend in depositional texture, with increased grain-support toward the top. The sequence terminates with gravel-sized oncoids. 1 – Very fine to fine, sandsized peloidal, restricted skeletal wackestone to packstone. 2 – Medium to very coarse, sand-sized skeletal grain packstone to grainstone. 3 – Coarse, sand-sized to gravel-sized oncolitic, lithoclastic grainstone to rudstone. The limestone member is interpreted as a product of subtidal deposition within a shallow-marine, back-margin platform lagoon with normal salinity. The boundary between the subtidal limestone and the upper vadose dolostone cap (4) is irregular and gradational. The microfacies of the vadose cap is characterized by pisoid packstones and grainstone with caliche pisoids and other caliche fabrics. The vadose cap is interpreted as the product of an extended episode of subaerial exposure lasting several thousands of years. After Hinnov and Goldhammer (1991).
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tion of both. Both auto- and allocyclicity are also discussed as major controls on the origin of the peritidal cycles. Fig. 16.3 shows an ideal Latemar cycle as visualized by Hardie et al. (1986) and Goldhammer et al. (1990). The Middle Triassic Latemar Limestone in the Dolomites consists of hundreds of vertically stacked, meter-scale, platform carbonate cycles. Each cycle (couplets) is composed of a shallow-marine subtidal carbonate unit with a thin, centimeter-scale vadose diagenetic dolomitic cap. The couplets are interpreted as the products of repeated episodes of submergence/emergence with durations estimated at about 20,000 years. The dominant organization of successive couplet thicknesses into upward-thinning bundles of five suggested to Goldhammer et al. (1987) a control by 100,000 years sea-level oscillations. Spectral analysis gave evidence of Milankovitch-forced, small amplitude sea-level oscillations (Hinnov and Goldhammer 1991). This interpretation has been strongly criticized by Eggenhof et al. (1999) who found no evidence of a Milankovitch control on the Latemar cycle. In terms of Standard Microfacies Types the submarine unit of the Latemar model consists of SMF 16NON-LAMINATED at the base (common in platform interior environments), overlain by SMF 17 (aggregategrain grainstone, common in restricted platform interior environments), followed by SMF 12 (oncoid rudstone, known from different high-energy environments including back-margin lagoonal areas). The vadose cap corresponds to SMF 26 (pisoid rudstone, characterizing subaerial exposure). This sequence of SMF types reflects notably the trends in the changes of relative water depths postulated by Goldhammer et al. (1987). Basics: Cyclic carbonates Overviews De Boer, P.L., Smith, D.G. (eds., 1994): Orbital forcing and cyclic sequences. – International Association of Sedimentologists, Special Publications, 19, 576 pp., Oxford Drummond, C.N., Wilkinson, B.H. (1993): Carbonate cycle stacking patterns and hierarchies of orbitally forced eustatic sealevel change. – Journal of Sedimentary Petrology, 63, 369-377 Drummond, C.N., Wilkinson, B.H. (1993): Aperiodic accumulation of cyclic peritidal carbonate. – Geology, 21, 1023-1026 Einsele, G., Ricken, W., Seilacher, A. (eds., 1991): Cycles and events in stratigraphy. – 995 pp., Berlin (Springer) Fischer, A.G., Bottjer, D.J. (eds., 1991): Orbital forcing and sedimentary sequences. – Journal of Sedimentary Petrology, Special Issue, 61, 1063-1252 Ginsburg, R.N. (1971): Landward movement of carbonate mud; a new model for regressive cycles in carbonates. – American Association of Petroleum Geologists, Bulletin, 55, p. 340
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References on Cyclic Carbonates
Box 16.1. Selected references on cyclic carbonates. Most references relate to shallow-marine carbonates of platforms. Papers on cyclicity of pelagic carbonates (marked by an asterisk) usually relate to Cretaceous carbonates. Tertiary: Fischer and Roberts 1991; Herbert et al. 1995*; Neuser et al. 1982; Reijmer et al. 1988; Reining et al. 2002; Scholle et al.1998* Cretaceous: Arthur and Dean 1991*; Arthur et al. 1986*; Bachmann and Willems 1996; Barron et al. 1985; Bersezio 1993*; Boyd et al. 1994*; Claps and Masetti 1994*; Cotillon and Rio 1983*; D’Argenio et al. 1989, 1997; De Boer 1982*; Fiet and Gorin 2000*; Fischer et al. 1991*; Füchtbauer and Tisljar 1975; Grötsch 1994, 1996; Grötsch et al. 1994; Herbert et al. 1995*; Jimenez De Cisneros et al. 1993; Joachimski 1994; Longo et al. 1994; Mettraux et al. 1999*; Noé 1993*; Prokoph et al. 1995*; Ricken 1994*, 1996; Röhl and Strasser 1995; Russo et al.1991; Sabins 1964; Schindler and Conrad 1994; Schwarzacher 1994*; Stössel 1999; Strasser 1986, 1987, 1988, 1991, 1994 Jurassic: Bosence et al. 2000; Dozet 1993; Dragastan et al. 1987; Goldhammer et al. 1991; Hallam 1963, 1964; Oschmann 1990; Sun 1990; Talbot 1973; Weedon 1986; Zempolich and Erba 1999 Triassic: Balog et al. 1997; Bechstädt 1973, 1975; Bechstädt and Schweizer 1991; Bosellini 1967; Bosellini and Hardie 1988; Burchell et al. 1990; Dozet 1990; Egenhoff et al. 1999; Fischer 1964 1975; Froh 1976; Goldhammer and Harris 1989; Goldhammer et al. 1987, 1990, 1993; Haas 1982, 1991a, 1991b, 1994; Haas and Dobosi 1982; Hagemeister 1988; Hardie et al. 1986; Hinnov and Goldhammer 1991; Kalpakis and Lekka 1982; Klotz 1990; Masetti et al. 1988, 1991; May 1997; Olsen 1984; Reijmer et al. 1994; Röhl 1990; Sander 1936; Satterley 1996; Satterley and Brandner 1995; Schauer and Aigner 1997; Schwarzacher 1948, 1954; Schwarzacher and Haas 1986; Schweizer et al. 1988; Zeeh 1994
Heckel, P.H. (1984): Factors in Mid-Continent Pennsylvanian limestone deposition. – In: Hyne, N.J. (ed.): Limestones of the Mid-Continent. – Tulsa Geological Society, Special Publication, 2, 25-50 Heckel, P.H. (1994): Evaluation of the evidence for glacioeustatic control over marine Pennsylvanian cyclothems in North America and consideration of possible tectonic effects. – In: Dennison, J.M., Ettensohn, F.R. (eds.): Tectonic and eustatic controls on sedimentary cycles. – SEPM, Concepts in Sedimentology and Paleontology, 4, 65-87 House, M.R., Gale, A.S., Johanesse, E.P., Mathiu, C. (eds., 1995): Orbital forcing timescales and cyclostratigraphy. – Geological Association London, Special Publication, 85, 210 pp., London (Geological Society, Publishing House) Merriam, D.F. (ed., 1964): Symposium on cyclic sedimentation. – Kansas Geological Survey, Bulletin, vol. 1 and 2, 636 pp. Osleger, D. (1991): Subtidal carbonate cycles: implications for allocyclic vs. autocyclic controls. – Geology, 19, 917920 Pratt, B.R., James, N.P. (1986): The St.George Group (Lower Ordovician) of western Newfoundland; tidal flat island model for carbonate sedimentation in shallow epeiric seas. – Sedimentology, 33, 313-343
Permian: Buggisch et al. 1994; Grant et al. 1994; Ludwig 1994; Mack and James 1986; Mazzullo 1999; McCorne 1963; Miller and West 1993; Mutti and Simo 1993; Osleger 1998; Osleger and Tinker 1999; Saller et al. 1994; Silver and Todd 1969; Vai and Venturini 1997; Yang et al. 1994 Carboniferous: Adlis et al. 1988; Archer and Feldman 1994; Brown et al. 1990; Connolly and Stanton 1992; Dott 1958; Driese and Dott 1984; Eldrick and Read 1991; Gibling and Bird 1994; Goldhammer et al. 1991a, 1991b, 1994; Harbaugh 1964; Heckel 1983, 1984, 1986; Homann 1969; Horbury and Adams 1989; Klein 1990, 1993; Klein and Kupperman 1992; Knapp and French 1991; Langenheim and Nelson 1992; Moore 1931; Olszewski 1996; Saller et al. 1994, 1999; Soreghan 1994; Vai and Venturini 1997; Wanless and Weller 1932; Weller 1931, 1964; West et al. 1997; Wiberg and Smith 1994; Yancey 1991 Devonian: Brownlaw et al. 1996, 1998; Chen et al. 2001; Dietrich and Wilkinson 1999; Eldrick 1995; Fejer and Narbonne 1992; Gischler 1995; Holmes and Christie-Blick 1993; Preat and Racki 1993; Racki 1992; Schudack 1993; Yang et al. 1995; Yang and Yang 1984 Silurian: Jeppson 1990; Johnson and Lescinsky 1986 Ordovician: Goldhammer et al. 1992, 1993; Kerans and Lucia 1989; Kim and Lee 1998; Lee et al. 2001; Read and Goldhammer 1988; Steinhoff and Walker 1995; Webber 2002; Young et al. 1972 Cambrian: Alvaro et al. 2000; Chow and James 1987; Elrick and Snider 2002; Glumac and Walker 2002; Hardie et al. 1991; Koerschner and Read 1989; Osleger 1991; Osleger and Read 1991; Selg 1988; Waters et al. 1989 Precambrian: Chan et al. 1994; Grotzinger 1986, 1989; Jackson 1985; Sami and James 1994
Read, J.F., Kerans, C., Webber, L.J., Sarg, J.F., Wright, F.M. (1995): Milankovitch sea-level changes, cycles, and reservoirs on carbonate platforms in greenhouse and icehouse worlds. – SEPM, Short Course, 35, 212 pp. Schwarzacher, W. (1994): Cyclostratigraphy and the Milankovitch theory. – Developments in Sedimentology, 52, 238 pp., Amsterdam (Elsevier) Schwarzacher, W. (2000): Repetitions and cyclic stratigraphy. – Earth-Science Reviews, 50, 51-75 Smith, D.G. (1994): Cyclicity or chaos? Orbital forcing versus non-linear dynamics. – In: DeBoer, P.L., Smith, D.G. (eds.): Orbital forcings and cyclic sequences. – International Association of Sedimentologists, Special Publication, 19, 531-544 Sprenger, A., Ten Kate, W.A. (1993): Cross-spectral analyses of two late Berriasian rhythmic limestone-marl successions in SE Spain and SE France favours orbital control. – Geologie en Mijnbouw, 72, 69-83 Strasser, A. (1991): Lagoonal-peritidal sequences in carbonate environments: autocyclic and allocyclic processes. – In: Einsele, G., Ricken, W., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 709-721, Berlin (Springer) Microfacies and cyclic carbonates Enos, P., Samankassou, E. (1998): Lofer cyclothems revis-
Cyclic Carbonates: References
ited (Late Triassic, Northern Alps, Austria). – Facies, 38, 207-228 Fischer, A.G. (1964): The Lofer cyclothems of the Alpine Triassic. – In: Merriam, D.F. (ed.): Symposium on cyclic sedimentation. – Kansas Geological Survey, Bulletin, 169, 107-149 Goldhammer, R.K., Dunn, P.A., Hardie, L.A. (1990): Depositional cycles, composite sealevel changes, cycle stacking patterns, and the hierarchy of stratigraphic forcing: examples from Alpine Triassic platform carbonates. – Geological Society of America, Bulletin, 102, 535-562 Goldhammer, R.K., Harris, M.T. (1989): Eustatic controls on the stratigraphy and geometry of the Latemar buildup (Middle Triassic), the Dolomites of northern Italy. – In: Crevello, P.D., Wilson, J.L., Sarg, J.F., Read, J.F. (eds.): Controls on carbonate platforms and basin development. – SEPM, Special Publication, 44, 323-338 Goldhammer, R.K., Lehmann, P.J., Dunn, P.A. (1993): The origin of high-frequency platform carbonate cycles and third-order sequences (Lower Ordovician El Paso Group, West Texas): Constraints from outcrop data and stratigraphic modeling. – Journal of Sedimentary Petrology, 63, 318-359 Grötsch, J. (1994): Guilds, cycles and episodic vertical aggradation of a reef (late Barremian to early Aptian, Dinaric carbonate platform, Slovenia). – In: De Boer, P.L., Smith, D.G. (eds.): Orbital forcing and cyclic sequences. – International Association of Sedimentologists, Special Publications, 19, 227-242 Grötsch, J. (1996): Cycle stacking and long-term sea-level history in the Lower Cretaceous (Gavrovo Platform, NW Greece). – Journal of Sedimentary Research, 66, 723-736 Grötsch, J., Koch, R., Buser, S. (1994): Fazies, Gildenstruktur und Diagenese des nördlichen Randes der Dinarischen Karbonatplattform (Barreme-Apt, W-Slowenien). – Abhandlungen der Geologischen Bundesanstalt Wien, 50, 125-153 Haas, J. (1994): Lofer cycles of the Upper Triassic Dachstein platform in the Transdanubian Mid-Mountains, Hungary. – In: De Boer, P.L. and Smith, D.G. (eds.): Orbital forcing and cyclic sequences. – International Association of Sedimentologists, Special Publication, 19, 303-322 Hinnov, L.A., Goldhammer, R.K. (1991): Spectral analysis of the Middle Triassic Latemar Limestone. – Journal of Sedimentary Petrology, 61, 1173-1193 Osleger, D., Read, J.F. (1991): Relation of eustasy to stacking patterns of meter-scale carbonate cycles, Late Cambrian, U.S.A. – Journal of Sedimentary Petrology, 61, 1225-1252 Satterley, A.K. (1996): The interpretation of cyclic successions of the Middle and Upper Triassic of the Northern and Southern Alps. – Earth-Science Review., 40, 181-207 Strasser, A. (1988): Shallowing-upward sequences in Purbeckian peritidal carbonates (lowermost Cretaceous, Swiss and French Jura Mountains). – Sedimentology, 35, 369383 Strasser, A. (1994): Milankovitch cyclicity and high-resolution sequence stratigraphy in lagoonal-peritidal carbonates (Upper Tithonian-Lower Berriasian, French Jura Mountains). – In: De Boer, P.L., Smith, D.G. (eds.): Orbital forcing and cyclic sequences. – International Association of Sedimentologists, Special Publications, 19, 285301 Further reading: K183, K195, K196
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16.1.2 Carbonate Sequence Stratigraphy The facies models described in Chap. 14 are limited with regard to their predictive capacity by their static view of time and relative sea-level changes. These difficulties can be overcome by models based on sequence stratigraphy showing how depositional sequences and systems tracts respond to lowstand, transgressive and highstand conditions of relative sea level under humid and arid conditions (Handford and Loucks 1993). Definition. Sequence stratigraphy is defined as the study of rock relationships within a chronostratigraphic framework, where the succession of rocks is cyclic and composed of genetically related stratal units (sequences and systems tracts; Posamentier et al. 1988). Seismic and sequence stratigraphy. Sequence stratigraphy was born in the 1960’s. It has its roots in seismic stratigraphy that extracts stratigraphic information from seismic data and inferred stratal geometries from seismic-reflection patterns. A relationship between the geologic rock record and its seismic response can be recognized by comparing the seismic record to cored drill holes and combining this with seismic modeling (Fagin 1991). The link between geology and seismic images are the physical properties of the rocks, in particular the acoustic impedance. Sonic velocity can be related to carbonate rock lithology and facies types (Rafavich et al. 1984; Wand et al. 1991; Anselmetti and Eberli 1993; Kenter and Ivanov 1995). Variations in sonic velocity are a result of the combined effect of variations in depositional lithology and diagenetic alterations (Anselmetti and Eberli 1993) also revealing distinct differences between carbonates deposited on shallow-water platforms and carbonates formed on slopes and in basins (Anselmetti et al. 1997). For important papers see ‘Basics’. Sequence stratigraphy of carbonates Sarg (1988) applied the sequence-stratigraphic model, originally proposed by Vail et al. (1977) for siliciclastic sediments, to carbonate platforms and bank margins. New sequence-stratigraphic models for carbonate shelves concern ramps, different types of platforms (e.g. intracratonic or isolated platforms) and carbonate-evaporite basins (Calvet et al. 1990; Hunt and Tucker 1993; Tucker et al. 1993). Recent overviews on the state of carbonate sequence stratigraphy have been given by Harris et al. (1999), Moore (2001) and Schlager (2002). Depositional bias: Sequence patterns are strongly influenced by depositional bias (differences among
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depositional systems; Schlager 1991). The interpretation of sequences in carbonate platforms requires more basic controls (facies and core analysis) than sequence stratigraphy analysis of siliciclastic rocks (Schlager 1992). Siliciclastic and carbonate sequences differ with regard to the additional control of sediment supply by changes in carbonate productivity (Schlager 1991; Driscoll et al. 1991). Siliciclastics shed most of their sediment during sea-level lowstands, shallow-water carbonates during sea-level highstands. Input of siliciclastics is strongly controlled by tectonics (influencing the relief and erosion of the hinterland). Carbonate systems are strongly controlled by environmental factors governing biogenic carbonate production. Carbonate and siliciclastic systems generate their own specific relief on the sea floor. Tropical carbonate platforms and reefs are often terminated by drowning unconformities. Drowning may be caused by environmental stress (caused, for example, by shifts in currents or hinterland drainage) reducing the growth potential of carbonate systems rather than by sea level. In contrast, sequence-stratigraphic patterns of temperate carbonates exhibit similarities with those of siliciclastic sediments (Read 1995).
Sequence Analysis: Basics
cal climates favor shallow-marine carbonate deposition. Arid climate favors high rates of evaporation and the deposition of evaporites. Humid climate promotes karstification of exposed carbonate platforms during sea-level lowstands by rainfall.
16.1.2.1 Sequence Analysis: Some Basics Sequence stratigraphic analysis is based on the subdivision of sedimentary sequences into units (sequences) that are bounded by surfaces of stratal discontinuities or their correlative conformities. These units are vertically arranged into depositional sequences (genetically related packages). The development of a sequence-stratigraphic framework for carbonate rocks requires integrating observations from core, wireline logs and outcrop analogs. The sequences (systems tracts; parasequences) are bounded updip by unconformities and downdip by correlative conformities. Systems tracts consist of all facies deposited during either lowstand, transgression or highstand sea levels. Lowstands and highstands are separated by the transgressive surface (flooding surface). The transgressive stand is separated from the highstand by the maximum flooding surface.
Controls The basic assumption for sequence stratigraphy is that sequences and their systems are essentially controlled by sea-level change and tectonics (Vail et al. 1977; Van Wagoner et al. 1988; Posamentier et al. 1988; Emery et al. 1996). However, the original and classical interpretation in terms of sea-level cycle is not the unique solution. The volume and composition of the sediment will have effects similar to sea-level fluctuations, especially in carbonate depositional systems (Schlager 1992). Environmental factors and not only sea level may represent a primary control of the internal architecture of sequences. Therefore, dominant controlling factors may change over time (Posamentier and James 1993). Carbonate stratigraphic sequences may be regarded as a composite record of • sea-level fluctuations (causing changes in accommodation space), • changes in sediment supply (product of carbonate production and erosion) controlled by • environmental factors (climate; growth potential of organisms) independent of sea level. Changes in the water temperature and nutrients, e.g. control the formation of reef carbonates. Warm tropical and subtropi-
Systems tracts An ideal sequence can be divided into three major components that deposit at different phases of the relative sea level. These systems tracts are defined mainly on the basis of their stratal geometries, using terminations of seismic reflectors of onlap, downlap, and offlap, but must be also defined in terms of facies and microfacies (see Sect. 16.1.2.2). • Lowstand systems tract (LST). Includes all deposits accumulated after the onset of relative sea-level fall. The shelf is exposed to erosion and/or karstification producing secondary porosity. Clastics will be brought closer to the basin of accumulation. The carbonate factory is moved basinward and is productive over a narrower area, while deeper basins may be starved of carbonates and exhibit condensed horizons. Characteristic features of LST are common to abundant siliciclastic input, exposure surfaces on marine carbonates, karst, caliche horizons and increased influx of siliciclastics. • Transgressive systems tract (TST). Comprises the deposits accumulated from the onset of coastal transgression until the time of the maximum transgression of the coast (maximum flooding surface), just prior to
Parasequences
a renewed regression. If the rise is rapid, carbonate platforms will drown, producing a drowning unconformity surface (characterized by a rapid lithological change from shallow-marine carbonates to deep shelf, slope or basin deposits). Characteristic features of TST are deepening-upward and increasingly open marine conditions indicated by the composition of the biota, the increase in bioclastic and oolitic carbonates as well as the relatively small thickness of thin-bedded deposits. • Highstand systems tract (HST): Formed by regressive deposits that accumulate when the sedimentation rates exceed the rate of relative sea-level rise. The shallow shelf is widened, producing and shedding the greatest quantity of carbonate sands having the best potential for reservoir rocks. Carbonate sand production at shelf margins outpaces the relative sea-level rise or standstill and the basinward migration of carbonates result in aggrading prograding geometries. As accommodation space decreases, cycles become progressively thinner. Characteristic features of HST are thick-bedded or massive deposits, often light rock color, marine biota, shoaling-upward patterns and a change from open-marine to restricted conditions. Framework reefs are generally regarded as characteristic features of HST but may also form during the development of LST and TST. Because TST or HST may be the result of a relative sea-level change or a change in sediment supply, the relations of shoaling/deepening trends to sea-level fluctuations must be critically evaluated.
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very rapid flooding or a sudden increase in water depth. A thin transgressive unit may be developed between the top of the underlying parasequence and the upwardshoaling part of the overlying sequence. Parasequences are often organized into parasequence sets or bundles grouped into larger, third-order depositional sequences. These sequences have a well-defined internal structure in which subunits are arranged into packages or parasequences with contrasting stacking patterns. The packages of parasequence sets are bounded by disconformities. Parasequence sets display trends of shallowing, deepening or no change in water depth. Lehrmann and Goldhammer (1999) have argued for significant secular variations in the development of parasequence and facies stacking patterns during the Earth’s history.
Depositional sequences and systems tracts record the interplay of accommodation (the space available for potential sediment accumulation) and sediment supply. The separation of effects of accommodation, sediment supply and sea-level fluctuations needs diagnostic facies criteria, which are more distinctive in siliciclastic sediments (Posamentier et al. 1988; Posamentier et al. 1992) than in carbonates, but can nevertheless be effectively derived from outcrop, well data and microfacies.
Parasequences correspond to successions developed in response to periodic or episodic environmental changes occurring in cycles of fourth-order (10 3 years) to fifth-order (104 years). Many parasequences relate in part to high frequency, Milankovitch-driven climate changes and their associated sea-level fluctuations. This is indicated by (1) estimates of cycle durations, (2) bundling of carbonate cycles into sets of five, (3) spectral analysis of time series constructed from cyclic stratigraphic sections, and (4) evidence of high-frequency sea-level drops in the platforms, generating disconformities, breccias, soils, or vadose fabrics on the tops of parasequences (Koerschner and Read 1989; Goldhammer et al. 1990). On the inner platform, parasequences consist of shallow lime sands or muds, capped by tidal flat facies, and exposure surfaces. On the shelf edge and shoal complexes of ramps, parasequences may be poorly cyclic to non-cyclic and grainstone-dominated. In slightly deeper water, parasequences show coarsening upward from lime mud deposits and interbedded muds and sands up into grainy caps.
Parasequences Most sequences consist of small-scale, upward-shallowing successions without significant breaks. These parasequences forming the smallest fundamental composite building blocks of systems tracts have a thickness of <1 to >10 m and can be delineated by the stratal architecture of their component lithofacies, their bounding surfaces or the contacts between each system. Parasequences are bounded by minor marine flooding surfaces, across which deeper water facies abruptly overlie shallower water facies (Van Wagoner et al. 1988). In shallow-marine environments, parasequences are characterized by an upper boundary that indicates
Sequence boundaries Following the classical sequence stratigraphy approach, the most important criteria for recognizing sequence boundaries are incised valley and related unconformities. These criteria are difficult to trace on carbonate platforms. Sequence boundaries of platforms are often characterized by hardgrounds and meteoric caps, in basins by increasing terrigenous input, abundant allochthonous sediment and coarse microfacies. Main sequence boundaries are • Emersion or subaerial surfaces. Sequence boundaries directly underlying deeper-water facies.
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Microfacies and Sequence Boundaries
• Transgressive surfaces: Surfaces directly underlying deeper-water facies. • Maximum flooding surfaces: Characterizing the relatively deepest depositional environment.
are reflected by statistically-based patterns described in Sect. 14.4 as ‘dynamic microfacies types’. This method assists in definiting sequence tracts and of hidden sequence boundaries.
Van Wagoner et al. (1988) and Schlager (1998) distinguished three major types of sequence boundaries: Type-1 sequence boundary forms when the relative sea level falls below the shelf break of the preceding sequence. Type-1 sequence boundaries are characterized by the emergence of the platform out beyond the shelf edge. This type is more common on rimmed shelves than ramps. This type is characterized by distinct signatures that are recorded by microfacies criteria (e.g. exposure; paleokarst, Pl. 129; pedogenic overprint on marine deposits; Pl. 128/1). Type-2 sequence boundary forms when relative the sea level falls to somewhere between the old shoreline and the shelf break. Only the inner shelf becomes exposed. Type-2 sequence boundaries may have erosional disconformities updip, but the sea level does not significantly expose the margin. The sequence boundary, therefore, is conformable. This type, associated with small sea-level falls, is typical of ramps. Type-3 sequence boundary forms when the sea level rises faster than platforms can aggrade and a transgressive systems tract with a significant marine hiatus overlies the preceding highstand tract (Schlager 1998, 1999). Marine erosion is particularly intensive on the top of drowned platforms (see Sect. 15.6.2).
Sequence-related diagenesis: Diagenetic features are key criteria to sequence stratigraphy (Moore 2001; Booler and Tucker 2002) and the recognition of subaerial exposure phases (Fouke et al. 1995). Significant porosity may develop beneath sequence-bounding unconformities and parasequence disconformities on carbonate platforms (Read and Horbury 1993). Normalmarine parasequences exhibit first-generation isopachous cement in grainstones. The rarity of marine cement may indicate rapid carbonate accumulation during a relative sea-level rise (keep-up carbonate system; Sarg 1988). Arid climate would prevent meteoric vadose/phreatic cement and produce grain-to-grain contacts and common pressure solution. Late highstand conditions in peritidal sequences are recorded by meteoric-vadose cement and crystal silt.
16.1.2.2 Microfacies Data Applied to Sequence Stratigraphy Relative sea-level fluctuations will affect the composition and distribution of microfacies types both of platform and ramp carbonates as well as on slopes and in basins. Sea-level fluctuations causing shallowing-upward cycles and parasequences are clearly documented by microfacies data and biota, allowing changes in depositional water depths and accommodation space to be estimated. Microfacies data are of particular value in evaluating sequence boundaries (discontinuity surfaces; Sect. 5.2.5; Fig. 16.4), changes in paleo-water depths (compare Sects. 9.3.4, 12.3 and 14.2) and shallowing/deepening trends of the sea level, reflected by changes in hydrodynamic energy (Sect. 12.1.1). Recent studies (e.g. Della Porta 2003) use microfacies as an integrated part of the interpretation of depositional models in terms of sequence stratigraphy. Shifts in the relative importance of grains due to environmental changes caused by sea-level fluctuations
Recognizing sequence boundaries Discontinuities and unconformities are key elements in recognizing sequence boundaries. Small-scale and short-lived discontinuities as seen in thin sections and outcrops can reflect large-scale variations in sea level. Excellent examples of the potential of microfacies for differentiating various orders of sea-level changes have been provided by Evarts et al. (1995) and Hillgaertner (1998), see Sect. 5.2.5. Emersion (exposure) and subaerial surfaces pointing to sea-level fall are characterized by marine carbonates capped or penetrated by sediments exhibiting terrestrial facies features (Sect. 15.1.2). These features are documented by: • Surfaces with desiccation marks occurring in supratidal and upper intertidal zones (Sect. 15.5.1.2). • Pedogenic structures (Sect. 15.1.1) including paleosols capping or penetrating marine carbonate strata; roots and alveolar structures; caliche pisoids and caliche crusts, sometimes with tepee structures truncated by erosion. • Paleokarst recorded by microkarst structures, karstic solution cavities and joints in the underlying marine limestone; speleothems. Surfaces with a network of cavities penetrating tens of cm into the underlying limestones and filled with sediment of different texture: • Terrestrial sediments (red beds, green shales, detrital sands or conglomerates) overlying marine carbon-
Sequence Boundaries
ates; infiltration of sand or shale into fissures and karstic cavities. • Brecciated carbonates adjacent to unconformity and filling caves. • Updip evaporite solution breccia. • Diagenetic overprint. Change from marine diagenesis to freshwater diagenesis. Meteoric-vadose diagenesis at the top of shallowing-upward sequences. Dedolomitization. Multiple early diagenetic overprint indicating by crosscutting generations of vugs with geopetal infilling. Crosscutting can be used to establish diagenetic logs showing the succession of diagenetic events (Loreau and Cros 1988). • Pisoids (Sect. 4.2.6). Pisoids are common constituents of supratidal vadose caps of peritidal cycles. Pisolitic horizons intercalated within marine carbonates in-
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dicate breaks in sedimentation caused by sea-level drops and can record long-term non-marine phases (e.g. karstification). • Leaching and preservation of fossil shells (Vollbrecht and Meischner 1993). On rapidly subsiding platforms the sequence boundary may be cryptic and located within very thin cyclic carbonates showing evidence of maximum emergence (quartz sand/silts on tops or reworked into bases of cycles, clayey paleosols capping the cycles, and tepees, calichified or brecciated caps of cycles). Transgressive and flooding surfaces • Surfaces directly underlying deeper facies (transgressive surfaces).
Fig. 16.4. Sequence boundary and Transgressive Systems Tract. The thin section shows a limestone rich in corallinacean red algae, bryozoans and pelagic foraminifera. These limestones, resting on Cretaceous platform carbonates, are widely distributed in the Southern Apennines (Carannante 1982). Bryozoans (B) and melobesiid red algae (M) grew in an open-shelf environment. The association corresponds to rhodalgal grain association (see Sect. 12.2) indicating warm-temperate water. Algal thalli were bored (white arrows) and infilled with pelagic lime mud containing globigerinid foraminifera (GF; Orbulina and Globigerina). Multiple boring is indicated by boring in fractured algal fragments (black arrow). Boring also continued during further deposition of pelagic foraminiferal mud that developed into thin-bedded pelagic wackestones. The microfacies of the sample characterizes a sequence boundary at the top of shelf carbonates formed during a lowstand phase. The boundary corresponds to a minor flooding surface developed on a hardground. Hardground conditions are recorded by intensive and multiple boring. Deepening upward is indicated by the significant microfacies change (SMF 8 to SMF 3). Tertiary (Miocene): Cività di Pietraroia, eastern Matese Mts., Southern Apennines, Italy.
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Differentiating Systems Tracts
• Facies change surfaces characterized by marked microfacies contrasts, differences in grain sizes and significant differences in diagenetic patterns. • Marked landward shift of distal facies over proximal facies. • Hardgrounds (Sect. 5.2.4.1). Surfaces coated by mm-thick crusts of iron oxide. • Condensation surfaces (Sect. 5.2.4.2). • Borings (Sect. 9.3). Surface with extensive borings may indicate interruption in sedimentation or non-deposition and characterize a flooding surface. • Phreatic cements. • Breccias. Sedimentary breccias can mark underlying sequence boundaries (Biddle 1984; Garcia-Mondjar 1990; Sect. 5.3.3.5). • Hardground intraclasts indicate breaks in sedimentation, often connected with maximum flooding surfaces. • Oncoids are common at the base of transgressive sequences of low-energy environments and can mark maximum flooding surfaces (Wright 1983). • Palynomorphs. Because palynofacies parameters are linked more or less directly to the distance of the coast line and/or water depths, a rapid marine flooding will cause a significant signal in the composition of the palynofacies assemblage (Rameil et al. 2000). • Organic matter and bioturbation. Intensive burrowing suggesting low sedimentation rates, and concentration of organic matter may indicate the relatively deepest environment and the maximum flooding surface.
cessions. Siliciclastic intercalations and terrestrial deposits are common as well as mixed siliciclastic-carbonate deposits. Grains are predominantly bioclasts. Normal marine biota but skeletal grains reflecting adaption to brackish-water or restricted conditions are common (e.g. ostracods; miliolid foraminifera). Ooids scarce. Extraclasts (Sect. 4.2.8.2) may indicate an increased drainage of river systems caused by sea-level lowstands. Black pebbles are good markers of sea-level lowstand phases connected with subaerial exposure (Box 4.22).
Differentiating LST, TST and HST of platforms and ramps
Box 16.2. Preliminary guide to identifying systems tracts in a single section or borehole. After Schlager (2002).
Generally carbonate systems tracts are identified from geometric relations derived from shore-to-basin cross sections recorded in seismic profiles, large outcrops or well-correlated boreholes. However, there is an increasing demand to recognize sequences and systems tracts in outcrop sections or single boreholes as well. Some microfacies criteria are listed below. Schlager (2002) suggested a preliminary guide for identifying systems tracts shown in Box 16.2. This approach is speculative, but may become more reliable by using selected microfacies criteria as shown by the case studies in Sect.16.1.2.3. Lowstand Systems Tracts Shoalwater belts are much narrower than transgressive and highstand tracts. They are well flushed, rich in reefs, hardgrounds and rocky shores. LSTs are characterized by shallowing and deepening-upward successions forming meter-sized facies suc-
Transgressive Systems Tracts Platform interior well flushed and normal marine. Reefs and sand shoals narrow, with wide passages. Patch reefs occur far landward. TSTs develop deepening-upward and increasingly open-marine trends. Typically thin-bedded carbonates. Non-skeletal grains, e.g. ooids and peloids, tend to be more common in TSTs and HSTs, because their growth is triggered by tidal currents and extensive winnowing by waves on the flooded platform top. Intraclasts common. Skeletal grains diverse, comprising benthic foraminifera, mollusks, echinoderms, and algae. Lime mud rare. Highstand Systems Tracts Reefs and sand shoals at the platform margin common and wide and more continuous. Shoaling-upward facies successions. Restriction increases in platform interior environments. HSTs are characterized by an
The identification must start with the knowledge of the position of the section with respect to the long-term platform edge. This knowledge may be derived from regional geological information, clues from the use of Standard Facies Models or the evaluation of fossil assemblages. Section location seaward of the platform edge: • LST: All shoalwater intercalations in slope deposits. • TST: Slope deposits with platform debris low, decreasing upward in the section. • HST: Slope deposits with platform debris high, increasing upward in the section. Ooids, peloids, platform-derived mud and specific platform interior skeletal grains abundant. Section location landward of the platform edge: • LST: Represented by exposure surfaces and terrestrial deposits. • TST: Deepening upward and increasingly open-marine trend. • HST: Shallowing upward and increasingly restricted trend.
Sea-Level Changes
increase in the number of facies belts (particularly on ramps) and shallowing-upward successions forming meter-sized cycles. Often thick-bedded and light-colored. Carbonate mud in lagoons, restricted environments and tidal flats. Significant increase in the diversity of microfacies types and biota as compared with TSTs. Fully marine biota prevail, exhibiting different associations in different subenvironments. Ooids abundant. Aggregate grains (Sect. 4.2.7) are common grains of sea-level highstands carbonates. Recognizing high-frequency sea-level changes High-resolution sequence analysis aims to understand short-term depositional processes that took place on a time scale of some ten thousand to hundred thousand years. These analyses require a bed-per-bed investigation with respect to facies and depositional environment, the determination of depositional sequences (supported by photomosaics; Weidlich and Bernecker 2003), and the identification of sequence boundaries and maximum flooding surfaces. Hierarchical stacking of depositional sequences common in HSTs, and facies changes allow us to define small-scale, medium-scale and large-scale sequences (Strasser et al. 1999). These sequences of fourth, fifth, and even higher order express the smallest changes in the environment through facies changes (e.g. Goodwin and Anderson 1985; Grotzinger 1986; Strasser 1987). The changes produce relatively thin shallowing-upward sedimentary successions (parasequences) bounded by marine flooding surfaces. These parasequences exhibit different vertical successions, depending on their setting in platform, ramp and offshore environments. An instructive example from the Middle Jurassic of Switzerland was described by Gonzales (1996). Synchronous changes in sea-level are recorded • in off-shoal environments by a parasequence with marl layers at the base of thin upward-shoaling units, covered by increasingly proximal tempestites; • in ramp environments by oblique-stratified oolitic and bioclastic grainstones bounded by marine flooding surfaces; and • on the platform by marine flooding surfaces overlain by coral beds, oncoid beds and platform tempestites deposited under moderate energy conditions, which are in turn covered by high-energy, obliquestratified oolitic grainstones. Such parasequences deposited in response to smallscale, high-frequency sea-level changes can be corre-
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lated over distances of more than 100 km and may provide a chronological framework permitting detailed paleogeographical reconstructions (Strasser 1987). Differentiating HST and LST in slope and basinal environments Carbonate sediment is exported from the platform to platform flanks. Changes in carbonate production of the platform may be detected by a quantified analysis of grain composition of platform flanks. These changes may or may not coincide with stratal geometry and bed surfaces. Highstand and lowstand shedding. Sea-level fluctuations affecting the carbonate production in platform environments are recorded by shallow-marine aragonite-rich carbonate material exported to the slope and into basins. Most sediment is shed during sea-level highstands (Schlager et al. 1994). Highstand shedding is triggered by the flat tops of carbonate platforms providing large production areas that are many times larger than during lowstands, and by the presence of rimmed platform margins. Highstand shedding is pronounced in low-latitude platforms. Evidence of highstand shedding comes from interglacial periods of the Quaternary of the Bahamas, the Nicaragua Rise in the Caribbean and the Queensland Shelf. Examples of ancient highstand shedding have been described from the Pleistocene back to the Devonian (compare the case study documented in Sect. 15.7.5.2). Note that the highstand model is restricted to flat-topped platforms. It does not apply to carbonate ramps (Betzler et al. 2000; Westphal et al. 2004). Shedding is strongly reduced during lowstands by the very strong reduction or demise of carbonate production and the rapid meteoric and marine lithification preventing erosion of carbonate material (Droxler and Schlager 1985). Sediment shed from an exposed platform top (lowstand) differs from platforms with flooded tops (highstand) in the composition of coarse skeletal and nonskeletal grains. Quantitative compositional analysis of turbidites and other mass-flow intercalations in slope and basinal sediments reflect lowstand and highstand conditions: • Highstand shedding is indicated by the dominance of non-skeletal grains including abundant ooids, peloids and aggregate grains (grains that require flooded platform tops for their formation), in association with specific, shallow-marine platform-interior skeletal grains. Primarily aragonitic grains dominate.
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• Lowstand shedding is indicated by skeletal grains in association with reef-derived (or slope-derived) lithoclasts. Primarily calcitic grains are common. Deposition of carbonate breccias and blocks transported by debris flows will punctuate the pelagic sedimentation in the basin.
Microfacies in Sequence Stratigraphy
16.1.2.3 Case Studies: Sea-Level Fluctuations and Systems Tracts Documented by Microfacies The case studies in this section dealing with carbonate ramps and platforms demonstrate the potential of microfacies criteria and microfacies types in sequence
Plate 141 Microfacies, Sea-Level Fluctuations and Sequence-Stratigraphic Analysis: Case Study Part 1 The microfacies types shown in Pl. 141 and Pl. 142 characterize an Early Jurassic succession from the central Peloponnesus in Greece. The section consists of 33 limestone beds, each one with a thickness of 25 to 40 cm. The succession exhibits alternations of shallow-marine and deep-marine limestones. Pl. 141/1 shows the microfacies of a sample from the bottom of the section, Pl. 142/20 the microfacies of the topmost sample. Microfacies criteria and SMF types can be used to recognize changes in paleobathymetry, sequence boundaries and systems tracts. The section starts with a sample characterized by reworked and redeposited, possibly reef-derived material (–> 1). The samples shown in –> 2 and 3 correspond to SMF 10 typically developed in the Standard Facies Zones FZ 7 or FZ 2. Sedimentation in FZ 7 (open-marine platform interior) is reliable because of the association of ooids, benthic foraminifera and diverse biota. Figs. 4–6 correspond to SMF 9. Sedimentation in a deep shelf environment (FZ 2) is indicated by the association of echinoderms, bivalves, ostracods and foraminifera, and by strong burrowing. The samples of –> 7 and 8 display pelagic criteria (filaments, ammonites) supporting an attribution to SMF 3 and FZ 3 (deep shelf margin). Regarding the sample shown in –> 1 as allochthonous intercalation in the open-marine platform facies recorded by –> 2 and 3, the microfacies of –> 1 to 8 reflects a deepening-upward succession of a HST connected with sea-level rise. A distinct sequence boundary and subaerial exposure at the top of basinal sediments and LST conditions are recorded by the paleocaliche shown in –> 9 and 10. This boundary would demand a significant sea-level fall corresponding to a Type1 sequence boundary. 1 Poorly sorted packstone. Larger grains are recrystallized fragments of chaetetid coralline sponges (C). Smaller grains are foraminifera, shell debris and peloids. The microfacies indicates allochthonous redeposition. Sample B 1. 2 Bioclastic packstone. Grains: Small micritized radial ooids, coated echinoderm clasts (E) surrounded by rim cement, foraminifera (F; Involutina), ostracods (O). SMF 10 (indicating textural inversion and redeposition). Sample B 3. 3 Oolitic-bioclastic packstone. Grains: Radial ooids, echinoderms (E), bivalve and brachiopod shells, foraminifera (F). The matrix is inhomogeneous micrite. SMF 10. Sample B 4. 4 Strongly burrowed, fine-bioclastic wackestone. Larger grains are micritized echinoderm fragments, smaller grains are shell debris, ostracods and rare foraminifera (arrow). SMF type 9 (defined by a bioclastic wackestone in which diverse organisms have been jumbled through burrowing). Sample B 8. 5 Burrowed bioclastic wackestone. Grains: Ostracods, echinoderm debris (E), foraminifera (F), small gastropods (G). Grains with black Fe/Mn coatings and fossil molds filled with dark micrite indicate reworking. SMF 9. Sample B 10. 6 Burrowed fine-bioclastic packstone. Larger grains are foraminifera (F) and ostracods (O), smaller grains are densely packed shell debris, foraminifera and ostracods, some echinoderm debris (E). SMF 9. Sample B 12. 7 Pelagic bioclastic wackestone. Grains: Juvenile ammonites (A), filaments (arrows), some echinoderm fragments (E). SMF 3. Sample B 14. 8 Pelagic ‘filament’ wackestone. Grains: Filaments are very thin shells of pelagic bivalves (see Sect. 10.2.4.2), echinoderm (E) fragments. SMF 3-FIL. Sample B 15. 9 Paleocaliche. Diagnostic criteria are the nonfossiliferous clotted and microfractured, dark matrix (M) and the abundance of quartz grains (Q). Sequence boundary. Sample B 18. 10 Paleocaliche. Diagnostic criteria are the laminated texture (L), in-place brecciation (black arrows), fissures filled with quartz grains (white arrow), clasts exhibiting alveolar texture (A), and dark-colored reworked grains (‘black pebbles’, BP). Sequence boundary. Sample B19.
Plate 141: Microfacies, Sea-Level Fluctuation, Sequence Stratigraphy; Part 1
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analysis. Relative sea-level changes will affect the composition and distribution of microfacies types both of platform and ramp carbonates as well as on slopes and in basins. Box 16.3 lists papers dealing with the study of carbonate rocks using both the sequence-stratigraphic and the microfacies approach. Detailed bed-per-bed microfacies analysis allows the complex reactions taking place during the formation of parasequences to be demonstrated. The present section includes examples exhibiting the use of microfacies types in recognizing sequence boundaries and systems tracts. The example from Early Jurassic carbonates from Greece shown on Pl. 141 and Pl. 142 illustrates the recognition of sequence boundaries and systems tracts on a thin-section scale. The case studies dealing with
Shelf Carbonates from Greece
Cretaceous carbonate ramps and platforms of the Sinai (Egypt) demonstrate the differences in the composition and diversity of microfacies types in different systems tracts.
Early Jurassic shelf carbonates of central Greece Pl. 141 and Pl. 142 display microfacies types of a section composed of bedded limestones. The microfacies types are defined using grain composition and texture. Most microfacies types can be attributed to Standard Microfacies Types (SMF) that characterize major depositional facies belts (Standard Facies Zones, FZ). Microfacies criteria reflect changes in depositional water depths and the existence of subaerial and submarine sequence boundaries.
Plate 142 Microfacies, Sea-Level Fluctuations and Sequence-Stratigraphic Analysis: Case Study Part 2 (continued from Pl. 141) Microfacies types of an Early Jurassic succession of the central Peloponnesus in Greece. The topmost sample of this section is –> 20. The oncoid limestone of –> 11 (as well as the sample taken from the overlying bed, not shown in the plate) corresponding to SMF 13 was deposited in a shallow platform interior environment (FZ 2). The sample manifests sea-level rise and TST conditions. A sequence boundary is recorded by the lag deposit (SMF 14) shown in –> 12. The sample indicates reduced or no sedimentation. LST conditions are indicated by the fenestral limestones of tidal lagoons shown in –> 13 to 15. The coral/microbial boundstone of –> 16 and a wackestone with megalodontid bivalves in the overlying sample (not shown in the plate) point to changes from intertidal/supratidal to subtidal environments. A sequence boundary is recorded by the microbreccia of –> 17. This erosional boundary may record a transgressive surface on the top of a drowned platform. The samples displayed in –> 18–20 indicate sea-level rise and TST/HST conditions. The samples correspond to SMF 9 and the Standard Facies Zone FZ 2 (deep shelf), the same depositional environment as inferred for the samples from the lower part of the section (–> 4–6 on Pl. 141). 11 Oncoid rudstone. The oncoids consist of tubular porostromate cyanobacteria. Nuclei are shells. Oncoids are bound together by a network consisting of peloids containing a few foraminifera. Interparticle pores are filled with calcite cement. SMF 13 (defined by large cyanoids and porostromate oncoids). Sample B 20. 12 Lithoclastic floatstone. Large grains are rounded calcitic lithoclasts and pebbles (reworked fossils), smaller grains are lithoclasts. The arrows point to a blackened and iron-stained crust with quartz grains. SMF 14. Lag deposit. Sequence boundary. Sample B 22. 13 Laminar fenestral bindstone made of densely packed peloids. Note the downward extensions of the roofs of the fenestrae (indicating microbial contribution; black arrow) and the infilling with gray silt (indicating vadose diagenesis; white arrows). SMF 21-FEN (defined as fenestral peloid bindstone with birdseyes and stromactoid voids). Sample B 23. 14 Fenestral bindstone with coprolites. Note decapod coprolite grain (Favreina; arrows) and thin isopachous cement rims. SMF 21. Sample B 25. 15 Fenestral rudstone/grainstone. The arrow points to a reworked vadoid grain. SMF 21-FEN. Sample B 26E. 16 Coral/microbial boundstone/floatstone. Reworked corals (C) connected by microbial crusts (M). SMF 5. Sample B 27. 17 Microbreccia consisting of mudstone and packstone clasts exhibiting small voids. Interspaces between the clasts are partly infilled with fine-grained peloidal sediment (arrows). Sequence boundary. Sample B 29. 18 Bioclastic wackestone with gastropods, algal debris (A) and a few foraminifera. SMF 9. Sample B 31. 19 Burrowed fine-bioclastic wackestone with shell-echinoderm debris. SMF 9. Sample B 32. 20 Burrowed fine-bioclastic wackestone. Grains: Echinoderm debris (E), ostracods (O), small gastropods (G), foraminifera (F) and peloids. SMF 9. Sample B 33.
Plate 142: Microfacies, Sea-Level Fluctuation, Sequence Stratigraphy; Part 2
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The Middle Cretaceous carbonate ramp of the northern Sinai, Egypt Bachmann and Kuss (1998) combined detailed field observations of sedimentary structures, depositional geometries and paleontological criteria with microfacies studies, aiming at a reconstruction of the changing depositional environments and an interpretation in terms of sequence stratigraphy. The investigations were concentrated on Late Albian to Cenomanian carbonate-dominated ramp deposits in the northern Sinai. The lateral and vertical facies variability in the stratigraphic sections and the paleoenvironmental data derived from the microfacies data of 500 thin sections allow estimates of sea level-controlled environmental changes during deposition and the interpretation of facies variations in terms of sequence boundaries, transgressive surfaces, and maximum flooding surfaces. The resulting facies patterns characterize systems tracts and parasequences whose formation is reflected by microfacies composition and distribution. Microfacies of limestones and marls were defined according to texture, matrix and grain types. The microfacies spectrum comprises oolitic, bioclastic and intraclastic grainstones representing carbonate shoals, bioclastic packstones and floatstones recording rudist biostromes, bioclastic packstones with foraminifera and other biota formed on the open-marine shallow ramp, peloidal and bioclastic wackestones deposited in protected inner ramp settings, and peritidal wackestones and lime mudstones. The distribution of 24 microfacies types were analyzed with respect to their preferred occurrence within lowstand, transgressive and highstand systems tracts. Some microfacies types in the studied area are characteristic of specific systems tracts: Lowstand Systems Tract: This systems tract was characterized by a proximal emerged facies belt where non-deposition of continental sediments prevailed, a facies belt with restricted environments characterized by peritidal, lagoonal or brackish conditions, and an open-marine low-energy facies belt separating the inner ramp from the mid-ramp area. Common microfacies are: • grainstones with terrigenous quartz associated with sandstones and claystones, • fenestral mudstones and quartz-bearing dolomites of peritidal carbonates, • mudstones to wackestones with abundant miliolid foraminifera of protected inner ramp environments, • wackestones to packstones with brackish-water ostracods. Transgressive Systems Tract: The transition from a lowstand systems tract to transgressive systems tract is
Cretaceous Carbonate Ramp from Sinai
Box 16.3. Selected references dealing with sequencestratigraphic studies of carbonate rocks Tertiary: Baum and Vail 1988; Betzler 1989; Coffey and Read 2002; Cucci and Clark 1993; Donovan et al. 1988; Franseen et al. 1993; Friebe 1993; Geel 2000; Grötsch and Mercadier 1999; Kulbrock 1995; Kusumastutu et al. 2002; Luterbacher et al. 1991; Mutti et al. 1999; Pujalte et al. 1993; Rusciadelli 1999; Saller et al. 1993; Vecsei and Sanders 1997 Cretaceous: Arnaud-Vanneau and Arnaud 1990; Bauer et al. 2002; Bernoulli et al. 1992; Booler and Tucker 2002; Buchbinder et al. 2000; Christie-Blick 1990; Di Stefano and Ruberti 2000; Donovan et al. 1988; Dromart et al. 1993; Elder et al. 1994; Everts 1994; Everts et al. 1995; Gräfe and Wiedmann 1993; Hunt and Tucker 1993; Kulbrock 1995; Lehmann et al. 2000; Mettraux et al. 1999; Philip 1994; Pinet and Papenoe 1985; Quesne and Ferry 1994; Röhl and Strasser 1995; Rusciadelli 1999; Simo 1986; Strasser 1987, 1994; Wilmsen 2000 Jurassic: Aurell and Melendiz 1993; Brachert 1992; De Matos and Hulstrand 1994; De Raffaelis et al. 2001; Dromart 1992; Dromart et al. 1993; Durlet and Thierry 2000; Fürsich et al. 2003; Helm et al. 2003; Koenick et al. 1994; Lasemi 2002; Leinfelder 1993; Leinfelder and Brachert 1991; Leinfelder and Wilson 1998; Martire 1992; Merzeraud et al. 2000; Miall 1986; Monaco and Giannetti 2002; Pittett et al. 2000; Strasser 1994; Vail et al. 1987 Triassic: Aigner and Bachmann 1992; Goldhammer et al. 1990, 1991, 1993; Götz 1996; Lenoble and Canerot 1993; Lopez-Gomez and Arche 1993; Milroy and Wright 2000; Rameil et al. 2000; Reijmer 1998; Tucker et al. 1993 Permian: Brown and Loucks 1993; Ehrenberg et al. 2000; Gerard and Buhrig 1990; Hovorka et al. 1993; Hunt and Fitchen 1999; Mazzullo 1999; Mutti and Simo 1993; Osleger and Tinker 1999; Ross and Ross 1995; Saller et al. 1999; Sarg 1988; Sonnenfeld and Cros 1993 Carboniferous: Grammer et al. 1996; Khetani and Read 2002; Olszewski 1996; Rankey et al. 1999; Reid and Dorobek 1993; Smith 1999; Waite 1993; Weber et al. 1995; Wendte et al. 1992 Devonian: Middleton 1987; Southgate et al. 1993; Wendte et al. 1992 Ordovician: Choi et al. 1999; Long 1993; Meng et al. 1997; Pope and Read 1997 Cambrian: Chafetz et al. 1986; Meng et al. 1997; Montanez and Osleger 1993
marked by a total change in the previously common microfacies types, because the shallow ramp is now subdivided into two facies belts that are controlled by different hydraulic regimes. The microfacies types formed under high-energy conditions (carbonate shoal) include: • various bioclastic, oolitic and some intraclastic grainstones, prevailing together with • bioclastic and intraclastic packstones,
Cretaceous Platform from Sinai
• bioclastic floatstones with rudists and organic crusts, and • oolitic and oncolitic packstones. The microfacies types formed under low-energy conditions (protected inner ramp) include: • bioclastic wackestones and packstones with dasyclad algae, and • biopelpackstones and wackestones with benthic foraminifera. Highstand Systems Tract: Again, the transition from a transgressive systems tract to highstand systems tract is characterized by major changes in the distribution of microfacies types. The differentiation of the ramp into more facies belts is reflected by a significant increase in microfacies diversity. Widely distributed microfacies types are: • grainstones and packstones with rudist debris, • bioclastic floatstones with rudists, • peloidal and bioclastic wackestones and packstones, • packstones with benthic foraminifera, • fenestral mudstones and dolomites. The high amount of oolitic microfacies types compared with bioclastic grainstones is interpreted as result of an increased development of flat, high-energy areas during the TST, favoring enough accommodation space for the formation of ooids (see Jenkyns and Strasser 1995; Pittett et al. 1995). Increasing accommodation space supports the formation of ooids in comparison with skeletal grains (Schlager et al. 1994).
Late Cretaceous platform interior carbonates of the Sinai Fig. 16.5 shows microfacies and systems tracts of Turonian shallow-marine platform sediments studied by Bauer et al. (2002). The Sinai platform extended over a maximum length of about 200 km and was differentiated in the Turonian by a shallow intrashelf basin. The Sinai platform was drowned during the Early Turonian. Platform environments recovered in the Middle Turonian and carbonate platform interior sediments prevailed until the Late Turonian. From the Coniacian onward, siliciclastic input from the hinterland increased significantly and mixed siliciclastic-carbonate deposits accumulated in basins and swells. The study was focused on the vertical and lateral microfacies distribution and the relationships between grain composition and depositional environments based on the semiquantitative investigation of 450 thin sections from 20 sections. The latter support shorefaces, restricted lagoons, shallow subtidal environments and
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inner platform shoals, and deep open-marine basinal environments. Facies belts and the distribution of microfacies types were integrated within a sequence-stratigraphic framework. Sequence boundaries are documented by emersion (paleosols, caliche, vadose cements) or erosional surfaces connected with changes in lithology and hardgrounds, and by stratigraphic gaps. Typical sediments of LSTs are evaporites, claystones with plants, ostracods and oysters, and silt- and sandstones. Relative sealevel rise is recognizable from the reduction of siliciclastic influx and the development of inner-platform marls and coarse bioclastic and oolitic carbonates. Deep-water TST sediments are represented by marls and chalks with planktonic foraminifera and ammonites. HSTs are characterized by thick, shallow-subtidal deposits and muddy subtidal and lagoonal carbonates. Note that LST sediments are underrepresented in this example. Most microfacies types occur both in sediments formed under TST and HST conditions, but the diversity of microfacies types increases from TST to HST. Basics: Sequence stratigraphy of carbonates (overviews) Berg, O.R., Woolverton, D.G. (eds., 1985): Seismic stratigraphy II. An integrated approach. – American Association of Petroleum Geologists, Memoir, 39, 276 pp. Bracco Gartner, G.L., Morsilli, M., Schlager, W., Bosellini, A. (2002): Toe-of-slope of a Cretaceous carbonate platform in outcrop, seismic model and offshore seismic data (Apulia, Italy). – International Journal of Earth Sciences, 91, 315-330 Cross, T.A., Lessenger, M.A. (1988): Seismic stratigraphy. – Annual Reviews of Earth Planetary Science, 16, 319-354 Emery, D., Myers, K., Bertram, G., Griffith, C., Reynolds, M.N.T., Richards, M. (1996): Sequence stratigraphy. – International Association of Sedimentologists, 297 pp. Frostick, L., Steel, R. (eds., 1994): Tectonic controls and signatures in sedimentary successions. – International Association of Sedimentologists, 528 pp. Handford, C.R., Loucks, R. (1993): Carbonate depositional sequences and system tracts. – In: Loucks, R.G., Sarg, J.F. (eds.): Carbonate sequence stratigraphy: recent developments and applications. – American Association of Petroleum Geologists, Memoir, 57, 3-41 Harris, P.M., Saller, A.H., Simo, J.A.T. (eds., 1999): Advances in carbonate sequence stratigraphy: application to reservoirs, outcrops and models. – SEPM, Special Publication, 63, 421 pp. Lehrmann, D.J., Goldhammer, R.K. (1999): Secular variation in parasequence and facies stacking patterns of platform carbonates: a guide to the adaption of stacking-pattern analysis in strata of diverse ages and settings. – In: Harris, P.M., Saller, A.H., Simo, J.A.T. (eds., 1999): Advances in carbonate sequence stratigraphy: application to reservoirs, outcrops and models. – SEPM, Special Publication, 63, 187-225
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Systems Tracts and Microfacies
Fig. 16.5. Systems tracts and microfacies. The example from the Late Cretaceous of the Sinai (Egypt) demonstrates the relationships between systems tracts, environmentally diagnostic microfacies criteria and characteristic microfacies types. From left to right: Sequence boundaries and their features; systems tracts defined by sea level (LST, TST, HST; 1–14; mfs: maximum flooding surface); facies-diagnostic fossils and microfacies criteria; stratigraphic distribution of prevalent microfacies types. Modified from Bauer et al. (2002). Facies of systems tracts: 1 - HST, marls and limestones with mollusks and echinoderms. 2 - LST, siltstones and shales with plants and ferruginous ooids. 3/4 - TST/HST, transgressive surface (ts), decrease of quartzose sediments; maximum flooding surface (mfs), dolomites, limestones and marls with abundant rudists, oysters and corals. 5 - TST, ts, condensed deposits with abundant ammonites overlain by marls with planktonic foraminifera. 6 - HST, mfs, dolomites and oolitic limestones. 7 - LST, gypsum, siltstones and shales with plant remains and brackish ostracods. 8 - TST, ts, onset of shallowsubtidal deposits. 9 - HST, mfs, thick-bedded limestones and dolomites. 10 - LST - siliciclastics. 11 - TST/HST, bioclastic dolomites and limestones. 12 - LST, marls, shales and siltstones with plants, gypsum, Fe-ooids and caliche. 13 - TST, ts, onset of carbonates with ooids and reworked grains. 14 - HST, condensation at mfs, dolomites alternating with limestones. Microfacies types: Siliciclastic shoreface – S 2. Quartzose bioclastic packstones and grainstones. Lagoon – L 2, bioclastic packstones and wackestones with mollusks, echinoderms, ostracods, and benthic foraminifera. Shallow subtidal – P 1, peloidal wackestones and packstones with abundant mollusks. P 2, bivalve floatstones with oysters and rudist fragments. P 3, bioclastic packstone with green algae. P 4, foraminiferal wackestones with miliolids and packstones with bivalves. P 5, winnowed packstones and grainstones rich in peloids. High-energy subtidal – W 1, bioclastic packstones rich in coated grains. W 2, winnowed packstones rich in bioclasts. W 3, algal packstones with dasyclads and rhodophyceans, and with coated grains. W 4, oolitic packstones and grainstones. Deep open-marine environment – B 1, dolomitic wackestones and packstones with planktonic foraminifera. Loucks, R.G., Sarg, J.F. (eds., 1993): Carbonate sequence stratigraphy - recent developments and applications. – American Association of Petroleum Geologists, Memoir, 57, 545 pp. Mcdonald, D. (ed., 1991): Sedimentation; tectonics and eustasy. Sea-level changes at active margins. – International Association of Sedimentologists, Special Publication, 12, 528 pp.
Miall, A.D. (1997): The geology of stratigraphic sequences. – 433 pp., Berlin (Springer) Moore, C.H. (2001): Carbonate reservoirs. Porosity evolution and diagenesis in a sequence stratigraphic framework. – Developments in Sedimentology, 55, 460 pp., Amsterdam (Elsevier) Muto, T., Steel, R.J. (2000): The accomodation concept in sequence stratigraphy: some dimensional problems and
References: Carbonate Sequence Stratigraphy
possible redefinition. – Sedimentary Geology, 130, 1-10 Palaz, I., Marfurt, K.J. (eds., 1997): Carbonate seismology. – Society of Exploration Geophysics, Geophysical Development Series, 6, 443 pp. Posamentier, H.W., Allen, G.P. (2000): Siliciclastic sequence stratigraphy - concepts and applications. – Concepts in Sedimentology and Paleontology, 7, 1-209 Posamentier, H.W., James, N.P. (1993): An overview on sequence stratigraphic concepts: uses and abuses. – International Association of Sedimentologists, Special Publication, 18, 3-18 Posamentier, H.W., Summerhayes, C.P., Haq, B.U., Allen, G.P. (eds., 1993): Sequence stratigraphy and facies associations. – International Association of Sedimentologists, Special Publication, 18, 656 pp. Sarg, J.F. (1988): Carbonate sequence stratigraphy. – In: Wilgus, C.K., Kendall, C.G.St. C., Posamentier, H.W., Ross, C.A., Van Wagoner, J.C. (eds.): Sea level changes: an integrated approach. – Society of Economic Paleontologists and Mineralogists, Special Publications, 42, 155-181 Schlager, W. (1992): Sedimentology and sequence stratigraphy of reefs and carbonate platforms. – SEPM, Continuing Education Course Notes, 34, 71 pp. Schlager, W. (1993): Accomodation and supply - a dual control on stratigraphic sequences. – Sedimentary Geology, 86, 11-136 Schlager, W. (2002): Sedimentology and sequence stratigraphy of carbonate rocks. – 146 pp., Amsterdam (Vrije Universiteit/Earth and Life Sciences) Schlager, W., Reijmer, J.W.G., Droxler, A. (1994); Highstand shedding of carbonate platforms. – Journal of Sedimentary Research, B64, 270-281 Vail, P.R., Mitchum, R.M., Todd, R.G., Widmier, J.M., Thompson, S., Sangree, J.B., Bubb, J.N., Hatledid, W.G. (1977): Seismic stratigraphy and global changes in sea level. – In: Payton, C.E. (ed.): Seismic stratigraphy - applications to hydrocarbon exploration. – American Association of Petroleum Geologists Memoir, 26, 49-212 Van Wagoner, J.C., Posamentier, H.W., Mitchum, R.M., Vail, P.R., Sarg, J.F., Loitit, T.S., Hardenbol, J. (1988): An overview of the fundamental of sequence stratigraphy and key definitions. – In: Wilgus, C.K., Hastings, B.S., Kendall, C.G.St.C., Posamentier, H.C., Ross, C.A., Van Wagoner, J.C. (eds.): Sea-level changes - an integrated approach. – Society of Economic Paleontologists and Mineralogists, Special Publication, 42, 39-45 Weimer, P., Posamentier, H.W- (eds., 1993): Siliciclastic sequence stratigraphy - recent developments and applications. – American Association of Petroleum Geologists, Memoir, 58, 492 pp. Wilgus, V.K., Hastings, B.S., Kendall, S.G.St.C., Posamentier, H.W., Ross,, C.A., Van Wagoner, J.C. (eds., 1988): Sealevel changes. an integrated approach. – Society of Economic Paleontologists and Mineralogists, Special Publication, 42, 407 pp. Williams, G.D., Dobb, A. (eds., 1993): Tectonics and seismic sequence stratigraphy. – Geological Society London, Special Publication, 71, 226 pp. References on seismic stratigraphy and seismic modeling of carbonate rocks (K183) Andrews 1988; Anselmetti and Eberli 1993; Anselmetti et al. 1997; Berg and Woolverton 1985; Biddle et al. 1991; Boyd
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et al. 1993; Bracco Gartner et al. 2002; Campbell and Stafleu 1992; Chiocci et al. 1991; Cross and Lessenger 1988; Cross et al. 1993; Erlich et al. 1990; Fagin 1991; Floden 1980; Fontaine et al. 1987; Gerard and Buhrig 1990; Grötsch and Mercadier 1999; Jacquin et al. 1991; Jervey 1988; Kenter and Ivanov 1995; Mettraux et al. 1999; Mitchum and Uliana 1985; Palaz and Marfurt 1997; Payton 1977; Pinet and Papaneoe 1985; Rafavich et al. 1984; Rudolph et al. 1989; Stafleu and Schlager 1991; Stafleu and Sonnenfeld 1994; Stafleu et al. 1994; Underhill 1991; Vail 1987; Vail et al. 1977a, 1977b, 1977c; Zinke et al. 2001 Examples of the application of microfacies in sequence stratigraphy Bachmann, M. (1994): Die Karbonatrampe von Organya im oberen Oberapt und unteren Unterapt (NE-Spanien, Prov. Lerida): Fazies, Zyklo- and Sequenzstratigraphie). – Bericht aus dem Fachbereich Geowissenschaften der Universität Bremen, 54, 202 pp. Bachmann, M., Kuss, J. (1998): The Middle Cretaceous carbonate ramp of the northern Sinai: sequence stratigraphy and facies distribution. – In: Wright, V.P., Burchette, T.P. (eds.): Carbonate ramps. – Geological Society, London, Special Publications, 149, 253-280 Buchbinder, B., Benjamini, C., Lipson-Benitah, S. (2000): Sequence development of Late Cenomian-Turonian carbonate ramps, platforms and basins in Israel. – Cretaceous Research, 21, 813-843 Geel T. (2000): Recognition of stratigraphic sequences in carbonate platforms and slope deposits: empirical models based on microfacies analysis of Palaeogene deposits in northeastern Spain. – Palaeogeography, Palaeoclimatology, Palaeoecology, 155, 211-238 Hunt, D., Tucker, M.E. (1993): Sequence stratigraphy and carbonate shelves with an example from the mid-Cretaceous (Urgonian) of southeast France. – In: Posamentier, H.W., Summerhayes, C.P., Haq, B.U., Allen, G.P. (eds.): Sequence stratigraphy and facies associations. – International Association of Sedimentologists, Special Publication, 18, 307-341 Krawinkel, H., Seyfried, H. (1996): Sedimentologic, paleoecologic, taphonomic and ichnologic criteria for high-resolution sequence stratigraphy: a practical guide for the identification and interpretation of discontinuities in shelf deposits. – Sedimentary Geology, 101, 79-101 Maurer, F., Reijmer, J.J.G., Schlager, W. (2003): Quantification and input and compositional variations of calciturbidites in a Middle Triassic basinal succession (Seceda, Dolomites, Southern Alps). – International Journal of Earth Sciences, 92, 593-509 Miller, K.B., West, R.W. (1998): Identification of sequence boundaries within cyclic strata of the Lower Permian of Kansas, USA: problems and alternatives. – Journal of Geology, 106, 119-132 Preat, A., Carliez, D. (1994): Microfacies et cyclicité dans le Givetien supérieur de Fromelennes (synclinorium de Dinant, France). – Annales de la Societé géologique de Belgique, 117, 227-283 Reijers, T.J.A., Peters, S.W. (1997): Sequence stratigraphy based on microfacies analysis: Mfamosing limestone, Calabar flank, Nigeria. – Geologie en Mijnbouw, 76, 197215 Further reading: K183, K196
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16.2 Understanding Reef Carbonates Reef limestones are different. They differ from other carbonate rocks by the strong biological control on their formation. This control in turn is responsible for specific sedimentation patterns and cementation types. The evolution of reefs and reef organisms, and the development of reef attributes during time has been discussed in several texts (Fagerstrom 1987; Webb 1996; Wood 1999; Stanley 2001; Kiessling et al. 2002). Secular changes and variations, e.g. in main reef-building groups, the composition of reef communities, reef abundance, paleolatitudinal distribution, or reef setting, are evident. Modern reefs have been discussed in Sect. 2.4.3.4 and Sect. 2.4.4.3. Reefs are formed in tropical and nontropical, warm-water and cold-water environments (see Tab. 2.3). The following text is focused on ancient reefs and the question how reef limestones should be described with regard to microfacies and paleontological criteria. 16.2.1 What is a Reef? Definition. Riding (2002) presented a lucid and precise overview of what the term ‘reef’ means to geologists and which of the many names used in describing ancient structures interpreted as reefs are meaningful. He defined a reef as ‘a calcareous deposit created by essentially in place sessile organisms’. This broad definition aims to encompass reefs of all types, ignoring their scale, and unites structures that fundamentally are products of substrate colonization and controls by sessile organisms. Riding’s definition agrees widely with the definition proposed by Flügel and Flügel-Kahler (1992) and the slightly modified use in the book on ‘Phanerozoic Reef Patterns’ (Kiessling et al. 2002): ‘Laterally confined biogenic structures, developed by the growth or activity of sessile benthic organisms and topographic relief and (inferred) rigidity’. Only these broad definitions allow a comparative analysis of reef development and patterns in space and time.
16.2.2 Reef Types Beginning in the 1950’s, a differentiation of two major categories became fashionable: ‘organic reefs’ and ‘mounds’ • Organic reefs (also called ecologic reefs, biological
Reef Types
reefs, true reefs) are aquatic biosedimentary deposits formed essentially in place by procaryote and eucaryote skeletal or non-skeletal organisms. This category also includes structures often called framework reefs. Frameworks may be built by large in-place skeletons being in contact with small-scaled in-situ skeletons (micro-framework or detrital framework; Weidlich and Flügel 1995) or with non-skeletal microbes contributing to the formation of automicrites (Webb 1996). • Carbonate mud mounds are carbonate mud-dominated (micrite and fine-silt) deposits with topographic relief. Carbonate mounds differ from ecologic reefs in the scarcity of in-place skeletons, stromatolites and thrombolites, but probably represent a continuum of shared ecologies and sedimentary characteristics (Wood 2001). They are the least well understood subjects for reef studies owing to their often genetic obscurity and potential heterogeneity, as shown by Bridges et al. (1995) and Kopaska-Merkel and Haywick (2001). In practice, the terms mud mound and reef mound are widely used. Both types are characterized by relief. Mud mounds consist of a high amount of finegrained carbonate (some authors suggest it is more than 50% of the rock volume). Reef mounds (or skeletal mounds) are defined by bioclastic lime mud with minor amounts of organic bindings. All these terms refer to autochthonous, biologically controlled deposits. Geometrically similar moundshaped structures may, however, also develop by allochthonous accumulation of skeletal grains, e.g. crinoid particles or fine-grained reef debris. These calcareous bioclastic deposits have been referred to as hydrodynamic reefs. A careful inspection of sedimentary structures must clarify whether the deposit is an allochthonous, current-swept and transported deposit or an in-place accumulation. Crinoid limestones are used here as an example to reveal their varying origin. In-place degradation of crinoids or other skeletons (e.g. segmented algae) together with the baffling of sediment and stabilization of the substrate by microbes can contribute to the formation of mounds. Crinoid particles tend to be transported, and can also form dunelike structures without a genetic relationship to reefs or carbonate mounds. Hydrodynamic controls on the deposition of sand-sized reef debris of randomly scattered skeletal grains in reefs are possible (Fagerstrom and Weidlich 1999). Reefs and mounds are encompassed within the term buildup, originally defined as carbonate masses representing in-place accumulations of largely skeleton-de-
Reef Biota
rived carbonate sediment having some topographic expression above the sea floor during growth (Stanton 1967). In practice, buildup has become an often used term suggesting the existence of bodies of massive or only coarsely bedded limestones genetically related or not at all related to reefs, occurring within bedded carbonates or other lithologies. Seismic reefs recognized by geophysics are defined by geometry and reflection properties. They are often not congruent in size and geometry with ecologic reefs and mounds because they may also include massive forereef debris as well as the backreef apron (Schlager 1992).
16.2.3 Reef Fossils Reef fossils comprise sessile, skeletonized organisms contributing to the formation of reef structures, and organisms using the diverse ecologic niches of reef dwelling. Preservation of reef fossils is relatively good in Paleozoic reef limestones because many skeletons consist of calcite. Mesozoic and Cenozoic reefbuilders with a prevailing aragonite skeleton mineralogy are often poorly preserved and recorded by leached and recrystallized fossils, or by open or closed molds. These molds or recrystallized material also provide some clues to the guild assignment and areal distribution of reefbuilders.
16.2.3.1 Reef Biota: Compositional Changes during Time The dominant reef builders changed considerably through time. Kiessling et al. (1999) and Kiessling (2001) defined seven major Phanerozoic reef phases based on prevailing high-ranked taxonomic composition and compositional reef types: • Early Cambrian to Early Ordovician: Volumetric predominance of microbes and microbial mounds, in the Cambrian in association with archaeocyath sponges (Pl. 82). • Middle Ordovician to Frasnian: Dominance of stromatoporoid sponges (Pl. 80) and corals (Pl. 83). Reefs and reef mounds prevailed. Bryozoans played a significant role until the Silurian. • Famennian to Early Permian: Metazoans except bryozoans (Pl. 144/7, 8) were subordinate to microbes and algae in reef construction. Reefs and reef mounds. Microbes were more abundant in the Famennian-Early
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Carboniferous interval (Pl. 8/6). Algae prevailed with the onset of the Pennsylvanian (Pl. 56, Pl. 58/1). • Middle Permian to early Late Triassic. Reef mounds, mounds or reefs dominated by coralline calcareous sponges (Pl. 42, Pl. 79, Pl. 145/1-3), corals (Pl. 116/2, 3) and microbes (Pl. 98/8). • Latest Triassic to earliest Cretaceous: Shallow-water reefs commonly dominated by scleractinian corals (Pl. 84). Siliceous sponges important in Middle and Late Jurassic mounds (Pl. 78/1, 4), bivalves in Jurassic banks. • Cretaceous: Reef mounds, reefs and biostromes dominated either by rudist bivalves (Pl. 88) or corals. • Cenozoic: Dominance of corals (Pl. 147/1, 4) and coralline red algae (Pl. 54). In fact, the temporal distribution of reefbuilders through time is more complex in regard to number of different reef-building associations and the predominance of a particular community. Statistical analysis of PaleoReef data revealed a great variety in the diversity of compositional reef types during the Ordovician, the Viséan to Early Carnian, the Callovian to Early Aptian, and the latest Cretaceous to Danian (Kiessling 2002).
16.2.3.2 Reef Guilds: Ecologic Units The reef guild concept was introduced, amplified and exemplified by Fagerstrom (1987, 1988, 1991), Weidlich and Fagerstrom (1998), and Fagerstrom and Weidlich 1999a, 1999b). A guild is defined as a group of species competing for the same class of environmental resources without regard to taxonomic position. The communities of organic reefs can be subdivided into five functional units corresponding to guilds. Three guilds contribute to the formation of reef frameworks – the Constructor Guild, the Baffler Guild and the Binder Guild (see Fig. 16.6). The Destroyer Guild includes those organisms that are active in the biological destruction of the reef framework by boring, rasping and biting. The Dweller Guild consists of the organisms living in the reef. Criteria used in distinguishing guilds include growth direction, growth form, skeletonization, transportability and taphonomic features, and skeletal packing density and areal cover. Typical constructors and bafflers have upward growth directions producing erect growth skeletons. Binders have predominating lateral growth producing relative thin, tabular-lenticular skeletons. Constructors differ from bafflers in having much larger and stronger skeletons, making the largest contribu-
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The Guild Concept
Criteria
Constructor
Baffler
Binder
Importance of criteria
1
habit, growth habit, growth direction
upwards, erect
upwards, erect
lateral, reptant
very high
2
form, life-form, growth form
massive, domes, branches cups, columns
cylinders, cones, blades
sheets, lenses, runners, webs, plates, umbrellas
high
3
skeletonization, skeletal strength
well-skeletonized strong, rigid
poorly skeletonized, mostly as skeletal fragments
well skeletonized
very high
4
skeletal volume, cloniality
large, clonal or gregarious
smaller, solitary or clonal
medium size, clonal or gregarious
moderate
5
biostratonomy, taphonomy, transportability
in growth position or in situ (toppled, broken)
in situ (toppled, broken), commonly transported
in growth position, encrust, roof-over or trap sediment
high
6
skeletal packing density, areal cover
as in life or less dense
highly variable from dispersed to concentrated
as in life
moderate
Fig. 16.6. Reef guilds. Criteria used in recognizing and assigning reef-building guilds. The habit, form and packing/density of reef-building organisms should be evaluated from outcrop observations combined with thin-section data. Skeletonization and skeletal volume can be derived from digital analysis of thin sections and grid studies of acetate sheets taken in the field (Bernecker et al. 1999; Fagerstrom 2002). After Fagerstrom and Weidlich (1999).
tion to the skeletal volume of the reef. Bafflers have small, more fragile skeletons. Bafflers are commonly preserved as toppled fragments. Constructors can be in life position, toppled on edge, toppled or overturned. Bafflers often reduce current velocities. Slower and more turbulent currents enhance the deposition of sediment within framework pores, larval recruitment, the feeding-filtration efficiency of suspension feeders. Interpreting the guild structure of fossil reef communities and the roles of individuals and taxa in framebuilding is a two-step process: • Assign the reef-building fossils to their membership in one or more constructor, baffler or binder guilds considering the nature and degree of intraspecific morphological plasticity (a part of a colony may belong to one guild, another part to a different guild) and the possibility of variable guild assignment (a species may be a member of several guilds). Important criteria for a correct determination of reef-building guilds are summarized in Fig. 16.6. • Determine the relative importance of each guild in the overall frame-building process, based on skeletal size, strength, the abundance of its assigned members and the areal cover by analyzing large acetate sheet tracings, close-up photos and inspecting in detail defined quadrat-surfaces associated with the investigation of thin-section samples.
Large thin sections often allow guild assignment, but field observations on a larger scale are pivotal. Strength and limitations of the reef guild concept: Guild structure analysis has been successfully used in interpreting the formation of various Phanerozoic reefs and recognizing changes in the overall contribution of constructors, bafflers and binders to reef formation during the Earth’s history. Secular variations are obvious for the Constructor Guild and the Binder Guild, whereas the importance of the Baffler Guild seems not to have varied strongly in the Phanerozoic (Kiessling 2002). Limitations of the concept relate to the • role of the Baffler Guild in reducing current velocity and enhancing sediment deposition within the reef framework. Baffler must not necessarily baffle; • guild assignment of microbes that may act as binders and constructors as well; • guild assignment of cryptic reef communities, e.g. pendant biota within reef cavities; • possibility that some reef organisms may be members both of the reef-building guild and of the Dweller Guild; • good community preservation during rising sea levels, limited preservation during falling sea levels.
Reef Carbonates
16.2.4 How to Classify Reef Carbonates? The unique position of reef carbonates is demonstrated by their specific grouping within textural limestone classifications that subdivide autochthonous reef carbonates with regard to the inferred role of sessile ‘reefbuilding’ organisms for the formation of the reef rock (Sect. 8.2; Fig. 8.1A). The differentiation of framestone, bafflestone and bindstone (Embry and Klovan 1971) is useful, but is not without problems. Problems relate to linking the rock names with specific skeletal shapes (e.g. bafflestone with stalk-shaped organisms and framestone with massive growth forms) and the incorporation of process-related and necessarily subjective terms in the classification. Using the terms bafflestone, bindstone or framestone (Pl. 41 and Pl. 42), or one of the other names listed in Box 8.1 requires concise information on the paleoecological and taphonomic criteria of supposed reef-builders. Otherwise, not only will the selected name be wrong, but also the environmental interpretation derived from the rock name, if the observed fabric is believed to reflect ecologic constraints within ancient reefs (see Fig. 8.1B and Fig. 8.1C).
Fig. 16.7. Coral reef limestone with the rugose coral Phillipsastraea. Description and classification of this sample must take into account growth forms of corals (dendroid), narrow or wide distance between corallites (close), mono- or polymict composition of reefbuilders (monomict), proportions of reefbuilders/matrix/cement (skeletons of reefbuilders and micrite matrix dominate; carbonate cement is restricted to shelter voids between corallites), composition of the matrix (peloidal micrite, probably microbial) and cement types (thin radiaxial-fibrous cements and late blocky cement in a few voids). The dendroid growth form and the micrite between the corallites, would make the samples correspond to a bafflestone, but genetically it is a framestone because corals and microbes formed a framework. The sample comes from mounds (High Relief Carbonate Mud Mounds after the classification of Riding 2002), famous because of the variegated colors of the lime mud (récif rouge) and a distinct depth-dependent vertical succession of facies (Boulvain 1990). The mounds were formed on a ramp. They exhibit marked differences in dominant reefbuilders and microfacies: Facies A formed below storm wave base is represented by red, stromatactis and sponge-rich mounds. Facies B developed on the top of facies A, characterized by pink coral-stromatoporoid mud mounds formed below storm wave base. Facies C starting development around the lower base of the photic zone is represented by gray coralthrombolite-calcimicrobe mounds. The sample has been taken from this facies. Late Devonian (Early Frasnian): Beauchâteau Quarry, Philippeville, southern Belgium. Width of the picture 5 cm.
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–> Caution: Note that the classification of rock samples into bafflestones, bindstones and framestones is not necessarily parallel to the assignment of the fossils within the samples to reef-building guilds! Guilds interpret the ecologic function of reefbuilders based on defined parameters. Reef rock categories are highly subjective. A bindstone sample may yield fossils belonging to the Binder Guild and to the Constructor Guild. A framestone can be formed by assemblages of Constructor-, Baffler- and Binder Guilds.
16.2.5 Microfacies Approach to Reef Studies Microfacies studies of reef limestones require large samples that provide data on the higher reef-building taxa present, and the major growth forms and orientations. A combination of quantitative studies using grids and core sampling is recommended (Webb 1999). In the field, reef limestones often appear as massive or only coarse bedded lens-shaped or domed carbonate bodies, varying in size from a few centimeters to several hundreds of meters. Stacked reefs can achieve thicknesses of tens to hundreds of meters. Reef lime-
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Reef and Mound Carbonates
Fig. 16.8. Rudist breccia. Reef debris? The thin section exhibits a fine-grained limestone breccia consisting mostly of poorly sorted rudist shells. Rudist breccias have often been interpreted as ‘forereef breccias’, but many breccias are more likely high-energy current deposits, storm layers, carbonate bodies of prograding carbonate platforms, or gravity-flow deposits. The sample is a bioclastic packstone of a storm deposit exhibiting larger rudist shells embedded within skeletal debris consisting of worn and rounded rudist fragments and some echinoderms. Reworked rudist reef material might be present, but this must be checked by a paleontological survey looking for characteristic reef taxa. Late Cretaceous: Melenisce, SSlovenia.
stones cover a broad textural spectrum, including boundstones, lime mudstones with or without sparfilled cavities, wackestones and packstones as well as grainstones. Reef talus and reef debris occur as lithoclastic and biolithoclastic floatstones and rudstones, and as carbonate breccia. Massive reef limestone. Fig. 16.7 shows a thin section of a massive reef limestone characterized by abundant corals and a fine-grained matrix. In the field the occurrence of corals is restricted to areas of only few square meters. Most of the limestone adjacent to the coral sample is a microbial lime mudstone forming the highest part of a carbonate mud mound. Reef debris. Many reefs have foreslopes consisting of eroded reef-margin material (see Pl. 27/2). Reef material can be transported far into the basin by turbidity currents or debris flows. Storm-related transport can lead to the accumulation of reef blocks and reef debris both on the slope and in backreef and lagoonal areas (see Sect. 12.1.2.2). The interpretation of the source area is based on the occurrence of reef-diagnostic fossils and characteristic microfacies types. However, a careful analysis of the composition of the carbonates is necessary as shown by the sample displayed in Fig. 16.8. The sample would be regarded in the field as an indication of the existence of reefs because of the abundance of rudist bivalve shells. Does the thin section give a distinct clue to a deposition of reef debris? The life habitats of rudist bivalves ranged from protected restricted to open platforms, and to platform margins, and also included deeper parts of ramps. Rudists were able to build topographically relevant organic mounds, but rarely gave origin to true reefs. ‘Rudist
breccias’ are a common feature of Cretaceous carbonate platforms and ramps. These breccias represent just one of three common limestone types with rudists that can be recognized in outcrops and the subsurface (Cestari and Sartorio 1995): • Limestones with rudists in original growth position (contributing to the formation of mounds); • limestone with rudists not in original growth position (contributing to tabular and mound-like geometries of rudist limestones); and • limestones composed of rudist fragments. This type is the most common one. It is characterized by grainor mud-support and various levels of fragmentation, abrasion and rounding of rudist shells. The texture of the grain-supported microbreccia shown in Fig. 16.8 indicates accumulation of strongly reworked rudist shells. Whether the clasts were derived from rudist reefs or not can be answered only by looking for characteristic reef taxa within the clasts of the breccia. 16.2.5.1 Basic Constituents of Reef and Mound Carbonates Rates of organic and skeletal growth, particulate sedimentation and early cementation determine accretion rates of reefs and mounds. In-place organic deposits can be described compositionally in terms of the relative proportions of three major components - matrix (M), skeletons (S) and cavity or cement (C). The MSC triangular diagram proposed by Riding (2002) assists in differentiating reef and mound types and separating reef rock samples with respect to differences in origin and environment. Fig. 16.9 and Fig. 16.10 can be used for evaluating the major constituents of reef rocks.
Example: Triassic Reef
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Fig. 16.9. Ancient reefs are built by the upward accretion of (1) sessile, ‘reef-building’ skeletal organisms, filling of voids between and within reef organisms with (2) sediment and/or with (3) marine cement. A: In this sample, reef-building organisms are represented by diverse coralline sponges (see Sect. 10.2.3.1); DS - large disjectoporids; IS - inozoan sponges including SS sphinctozoan sponges (Colospongia encrusting inozoans sponges) and a few scleractinian corals (C - Stylophyllum); CC - spar-filled cavities. Corals and sponges are biogenically encrusted (E). The white arrow points to Alpinophragmium, an encrusting foraminifer characteristic of the ‘central reef’ of many Late Triassic reefs typical of protected areas with active reef growth. Geopetal structures within sponges, the good preservation of the sponges, and the corresponding orientation of shelter pores suggest that most reefbuilders are in life position. The sediment between and within these organisms is a black micrite with bioclasts forming a packstone texture (see details in pictures B and C). Carbonate cement occurs in shelter pores beneath sponges and within some inozoan sponges. Cementation (CC) of some shelter pores started with cloudy radiaxal-fibrous marine cement (black arrow) followed by white granular cement. Other shelter pores are basically filled with sediment. The photographs B and C show enlarged details of the sediment between the reefbuilders. B: Consistent infilling of sediment both between as well as within sponges is recorded by fine-graded layers, possibly stabilized by microbial binding (black tubular structures). Enlarged details of the lower part of A. C: Skeletal debris within the homogeneous micrite consists of variously sized, broken mollusk shells and sponge debris. Considering the proportions of reef-building skeletal organisms, sediment forming the matrix of the reef limestone, and carbonate cement, the sample would correspond to a ‘close cluster reef’. This category, proposed by Riding (2002), is characterized by closely spaced skeletons that inhibit hydrodynamic removal and sorting of trapped sediment. The rock name of the sample is bafflestone although the sponges must be assigned to the Constructor Guild. The sample comes from a platform margin reef. Late Triassic (Dachstein limestone, Norian): Gosaukamm, Austria. After Wurm (1981).
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Reef Carbonates
Fig. 16.10. ‘Cement reef’. This reef limestone is characterized by the dominance of only two main constituents – cement and encrusting reefbuilders. Reef limestones formed by aggregates of serpulid worms (S), green (?) algae (A) and volumetrically dominant synsedimentary carbonate cements correspond to a specific reef type and denote specific environmental conditions. Laminated cement crusts (C), probably caused by microbial precipitation, provide substrate for encrusting organisms, e.g. microbes, sponges or serpulids. Rapid synsedimentary cementation imparts extra strength and stability to reefs. The microfacies of the sample is characteristic of meter- to decameter-sized mounds formed in particular paleogeographic locations in the Late Triassic of the western Mediterranean Tethys. Disaerobic, mesohaline and eutrophic conditions have been considered as important controls for this reef type. Cement reefs were common in Permian and Triassic, but are also known from other time intervals (e.g. Cambrian and Precambrian). Late Triassic (Norian): Cahorros de Monachil, Sierra Nevada near Granada, Betic Cordillera, Southern Spain. After Flügel et al. (1984).
16.2.5.2 Describing Reef Carbonates: A Practical Guide Box 16.4 summarizes the criteria of reef limestones that should be looked for carefully in thin sections and larger rock samples when evaluating the formation of ancient reefs and their diagenetic history.
16.2.6 Case Studies of Some Ancient Reefs The case studies demonstrate diagnostic microfacies criteria of well-known ancient mud mounds and reefs and discuss the problems involved in interpreting the formation of reefs and reef environments. Different genetic interpretations are possible for the spectacular Devonian Kess Kess mounds in Morocco, depending on the weighting given to microfacies, paleontological features or geochemical data. The Early Carboniferous Waulsortian Muleshoe Mound in the Sacramento
Mountains of New Mexico was a key for the creation of mud mound models, but as Lloyd Pray said, ‘there is an enormous gulf between seeing and perceiving’. The Muleshoe model has been remodeled again and again. The same happened to the world-famous Capitan Reef and the Permian Reef Complex in Western Texas and New Mexico. Mounds: The term ‘mud mound’ was originally used for biodetrital carbonate bodies in Florida Bay, whose origin is related to baffling of sediment by marine grasses and a surplus of carbonate production by epibiontic organisms. Other modern carbonate mud mounds represent hydrodynamic accumulations of mud produced by various algae and the breakdown of skeletal material. The formation of these mounds is not controlled by framework-building or accretionary organisms. Phanerozoic mud mounds are found in various settings, including ramp, foreslope, deeper shelf and seamount positions. They generally consist of more than
Guide for Reef Limestones
837
Box 16.4. Guide for the description of reef limestones. Describe biota, matrix (sediment) and cement with regard to their specific types and abundance. Take care that even a large thin section may not provide full information on the frequency of the three main constituents. Growth forms of reefbuilders observed on a thin-section scale should be checked by field data. Biota • Reef-building organisms. Describe type (on higher or lower taxonomic level). Try to estimate ranking and the volumetric importance of potential reefbuilders.
• Relationship between matrix and reef organisms. Can you suggest a causal relationship between the type or abundance of matrix and the existence, growth form, size or distribution of reef organisms?
• Growth forms and guild structure. Try to define the growth forms of sessile organisms using Fig. 12.2. Growth forms provide clues to hydrodynamic conditions and relative sedimentation rates. Assign reefbuilders to guild groups (Fig. 16.6).
• Potential microbial contribution. Can you find criteria supporting a microbially triggered origin of the reef rock matrix (see Box 9.1)?
• Encrusting organisms. Micro-encruster associations are environment-sensitive (Sect. 9.2.3; Fig. 9.6 and 9.10). Are encrusters completely absent? Are encrusters common or abundant (pointing to low sedimentation rates)? Are encrusters restricted to reef-building organisms? Or do they occur on particles of reef debris? • Micro- and macroborings in reefbuilders (indicating biological destruction and potential sediment production). Boring patterns can provide clues to paleo-water depth (Sect. 9.3.4). • Microfossils in the matrix between reefbuilders. Specific environments and subenvironments of reefs are often characterized by facies-diagnostic benthic foraminifera, algae and microproblematica (Sect. 14.2.2; Fig. 14.13). • Associated fossils. Which fossils occur in association with reefbuilders? Can you see recurrent patterns in the frequency and taxonomic composition of these fossils? Are these organisms constant reef dwellers using the niches within the reef or were they just visiting the reef environment? Matrix and sediment A close look at the matrix is necessary in assigning your sample to guild types and classifying the sample as bafflestone, bindstone or framestone. The abundance of matrix assists in differentiating reef types, e.g. mud mounds and reef mounds. • Abundance of matrix. Estimate or measure the abundance of matrix in comparison to reef biota and syngenetic carbonate cement. • Texture of the matrix. Describe matrix and reef sediment in terms of the Dunham classification (e.g. lime mudstone, wackestone, packstone, floatstone; Sect. 8.3.2). • Matrix composition. Characterize type, distribution and preservation of grains within the matrix (see Fig. 16.9). • Can you allocate the matrix to a specific genetic category? The matrix may be allomicrite, automicrite formed on free surfaces or as internal micrite within cryptic cavities, or calcisiltite (Sect. 4.1.4). See Fig. 4.1 and Box 4.4.
Carbonate cement Cements are essential constituents of ancient reefs. Reef cementation is controlled by water pumping and biological activity (e.g. microbes, organic matter), and depends on the size and connection of pores. Syngenetic Mg-calcite and aragonite cements contribute to the rigidity of reefs and assist in maintaining steep reef slopes. Cementation is particularly abundant at seaward margins of high-energy reefs, e.g. in forereef areas (Fig. 7.7), characterized by a high rate of water movement and low sedimentation rates. • Distribution of cement and pores. Do reef cements occur on free surfaces or in cavities? Are they found only in the interspace between autochthonous or parautochthonous reefbuilders (e.g. growth-framework pores), in intraskeletal pores (Pl. 29/2), or do they occur within reef cavities (e.g. stromatactis)? Are these cavities primary pores or solution cavities (e.g. biomolds)? • Cement types and fabrics. Describe cement types, cement sequences and cement fabrics (Fig. 7.12). Common types are fibrous, bladed, radiaxial and botryoidal cements. Radiaxial and botryoidal cements may dominate volumetrically over reefbuilders and sediment (Sect. 7.4.4.1, 7.4.4.2; Fig. 7.14; Box 7.7). Be aware that ‘micrite’ may not represent sedimentary matrix, but early diagenetic microcrystalline or peloidal microcrystalline cement! • Cement and fossils. Look for the interrelationship between cement and fossils, e.g. layered cement sequences (Fig. 7.14; Fig. 16.10) that indicate (often microbially mediated) changes in water flow and chemical composition of fluids within the pore system of reef structures. Intergrowth of micritic (potentially microbial) crusts or organisms depending on hard substrates and cement crusts indicates rapid synsedimentary cementation (Pl. 145/1). • Diagenetic environments. Evaluate cement types and cement distribution in terms of marine, meteoric and burial diagenesis (Sect. 7.2.1; Box 7.6). Ecologic shallow-water reefs tend to be affected by meteoric-vadose and marinevadose diagenesis, reflected by dripstone and meniscus cements, and deposition of crystal silt within solution cavities.
838
50% fine-grained carbonate mud (micrite and peloidal micrite, silt-sized carbonate). In contrast to modern mud banks, many of the Paleozoic and Mesozoic mud mounds are characterized by organically formed textures, originally called ‘cryptalgal’ and today ‘microbial’, stressing the predominant role of procaryotic organisms (e.g. bacteria) and eucaryonts. Siliceous sponges, various ‘crusts’ as well as mm- to cm-sized sediment- and spar-filled cavities (stromatactis) are commonly associated with these textures (see Reitner and Neuweiler 1995). Frame-building and baffling organisms seem to be of minor importance in the formation of most ancient mud mounds. Pivotal to the understanding of the origin of ancient mud mounds is the genetic interpretation of the microcrystalline carbonate (automicrite and/or allomicrite; see Sect. 4.1.2) and of the processes controlling the formation of these micrites. Mud mounds are formed by very different biological processes triggering the accumulation of fine-grained sediment (Pratt 1995). Ma-
Devonian Mud Mounds
jor controls on the formation of mud mounds are the shelf configuration (position in mid ramp to outer ramp part of homoclinal settings; imported mud can be distributed toward the mud mounds), reduced allochthonous sedimentation rates during rising sea level (transgression and early highstand), increased alkalinity of seawater facilitating calcification and allowing low to moderate grazing (facilitating growth of biofilms and mats).
16.2.6.1 Mud Mounds: The Early Devonian ‘Kess Kess’ Mounds in the Anti-Atlas, Southern Morocco The Kess Kess mounds are unique in comparison to other Paleozoic mud mounds. Many studies indicate that Paleozoic mound accretion is biologically controlled, either by sponges or microbes that may induce or trigger CaCO3 precipitation, or by delicate mud-trap-
Plate 143 Mud Mounds: The Early Devonian ‘Kess Kess’ Mounds in the Anti-Atlas, Southern Morocco Spectacular conical mounds crop out in the Hamar Lakhdad Ridge, 15 km southeast of Erfoud in the Tafilalt province. Mound and intermound carbonates are crudely bedded. They differ only slightly with regard to biotic composition and microfacies. Dominating texture types are lime mudstone and skeletal wackestone and packstone with crinoids, trilobites, tentaculitids and brachiopods and a scarce tabulate coral fauna consisting of auloporids, a few thamnoporids and favositids. Stromatactis-like spar-filled cavities and sheet cracks occur in the mound and in the intermound facies. Neptunian dikes crosscut the mound and the intermound facies. The mounds show no differentiation into specific core and flank facies. Intercalations of skeletal rudstones and grainstones are interpreted as storm-controlled eventstones. Slumping criteria are absent. 1 Mound facies. Densely packed crinoid wackestone. The role of benthic organisms for the origin of the mounds was most probably that of a parautochthonous accumulation of bioclastic sediment produced by the degradation of crinoids, trilobites and thamnoporid corals. 2 Tentaculitid wackestone with trilobite fragments. Intermound facies. Tentaculitids increase in frequency distally from the mounds toward the more pelagic open-marine basin. 3 Bioerosion is weak, but some boring occurs in larger crinoid columnals. Skeletal grains other than crinoids are trilobite fragments and a few auloporids. 4 Trilobite lumachelle and calcite cements. Bioclastic rudstones intercalated both in the intermound and the mound facies are interpreted as high-energy current deposition caused by storm events. Evidence of storm deposits are significant breaks in the size of skeletal grains from bed to bed. Fine-grained sediment is winnowed and interskeletal pores are filled with marine radiaxial cement. 5 Open-space fabric. Spar-filled, laterally elongated cavities occur both in the intermound and the mound facies, but are more common in the mound facies. Many of these stromatactoid cavities have been enlarged by dissolution between skeletal grains. The structures range from millimeters to tens of centimeters laterally. 6 Open-space fabric. Typical stromatactis cavity filled with basal micritic sediment and recrystallized sparry calcite. These millimeter- to centimeter-sized cavities are interpreted very differently (see Box 5.2). Following the hydrothermal venting model, the cavities might have acted as conduits. Mound facies. 7 Tabulate corals. Auloporid wackestone, a common microfacies of the mound carbonates. 8 Encrusting corals. Auloporids encrusting echinoderm fragment. Auloporids preferred firm and hard substrates; their occurrence provides a clue to rapid synsedimentary lithification of the fine-grained sediment of the mounds. 9 Auloporid corals and carbonate cement. Auloporids seem to have acted as initial points for early marine cementation. This is indicated by the frequent patches consisting of densely spaced auloporids surrounded by radiaxial and fibrous carbonate cements.
Plate 143: Mud Mounds (Kess Kess) of Morocco
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Mud Mounds
Fig. 16.11. South view of the Hamar Lakhdad Ridge southeast of Erfoud, southern Morocco, showing the Early Devonian Kess-Kess mounds exhumed from the overlying Emsian to Famennian shales and limestones. The height of the mounds ranges between 35 and 60 m. The flanks are asymmetrical, exhibiting a steep slope (about 50°) and a shallower slope. Most mounds have similar shapes. Courtesy of W. Buggisch (Erlangen).
Plate 144
Mud Mounds: Early Carboniferous (Lower Mississippian) Muleshoe Mound, Sacramento Mountains, New Mexico, U.S.A.
The dome-shaped Muleshoe buildup, 12 km SSE of Alamagordo in New Mexico, formed on the outer and deeper part of a ramp (Ahr 1989), comprises two, superimposed knoll-form ‘cores’ with a total thickness of more than 100 m and a width of several hundreds of meters across. The mound consists of a syndepositional wackestone core facies (100 m thick) flanked by steeply dipping crinoidal beds (Bolton et al. 1982). The lower slopes of the core facies and the proximal flanks were probably the major growth sites of crinoids at the mound. Much crinoidal sediment may be essentially in place. The lower part of the central core facies is mud-rich, lacking porosity and permeability and consisting of skeletal wackestones and lime mudstones with bryozoans and crinoids, micrite and void-filling cements (–> 8, 9). The upper part of the mound differs in the abundance of fenestrate bryozoans and sparry calcite (–> 3). The flanking beds (–> 1, 4, 5) are characterized by encrinites (crinoid packstones) exhibiting abundant solutional grain contacts. The intermound facies consists of skeletal wackestones (–> 2). 1 2
Proximal flank facies. Thin-bedded crinoid-bryozoan limestone. Eastern slope, lower part. Mound base. Sediment below the mud mound. Interbedding of thin-bedded crinoid-bryozoan limestone beds and marls. Grains are bryozoan fragments and crinoids. 3 Upper mound facies. Clast of a megabreccia deposited as debris flow along the flanks of the mound. The clast was derived from the upper core facies, as indicated by bryozoans embedded in micrite and the void-filling cement. 4 Flank facies. Thick-bedded crinoid limestone. Flank facies, interfingering with massive micritic core facies. Note compaction and stylolitic boundaries (arrows) between crinoid fragments. The biofabric of the sample is characterized by dense packing of abundant crinoidal elements. Proximal facies of the eastern slope, higher part. 5 Flank Facies. Bedded crinoid limestone. The flank facies interfingers with the massive upper core facies. Moderately sorted crinoid packstone. The arrows point to syntaxial echinoderm rim cement. Western slope. 6 Upper mound facies. Massive limestone of the central part of the upper mound facies. Poorly sorted crinoid packstone. In contrast to –> 5, syntaxial cement is absent, indicating differences in primary intergranular porosities. 7 Upper mound facies. Massive limestone. Central part of the upper mound core facies. Bryozoan-crinoid wackestone. Fistuliporid bryozoans (FB) are parautochthonous. 8 Lower mound facies. Hard, massive limestone. Mound facies. Some of the core facies are mudstones and wackestones poor in fossils. Common fossils are bryozoans and in places brachiopods. 9 Lower mound facies. Hard massive limestone. Mound facies. Burrowed bioclastic wackestone. Skeletal grains are echinoderms (E), bryozoans (B), ostracods (O) and sponge spicules. The micritic matrix is inhomogeneous and contains a high proportion of fine-bioclastic and peloidal particles. 10 Mound facies. Dark massive limestones with open-space structures (not shown). Interfingering of the upper mound facies with crinoid-rich grainstones and wackestones of the flank facies. Burrowed wackestone with crinoids and cryptostome bryozoans (Rhombopora, R).
Plate 144: Muleshoe Mound, New Mexico
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ping and sediment-stabilizing organisms, e.g. bryozoans. The Early Devonian Kess Kess mounds in southern Morocco do not fit into this model. The cone-shaped, steep-sided mounds occur on the top of a volcanic massif. The mounds are confined to a 0.5 x 3.5 km area above a basaltic high enplaced during the earliest Devonian above a shale-dominated Ordovician-Silurian succession. The upper surface of the volcanic high was colonized during the Pragian to Early Emsian time by crinoids that formed an up to 180 m thick cover of bioclastic sands and gravels. Mounds and intermound facies developed on top of the crinoid beds during the Late Emsian, and were buried by Late Emsian siliciclastic-rich mud. Key questions for understanding the formation of the Kess Kess mounds are: Why were the mounds formed where they are? What stimulated and controlled the carbonate mud production? Which factors controlled the shape of the mounds? And if there was a control on mound development by hydrothermal venting, how did venting affect or even kill mound organisms? The beautifully exposed mounds (Fig. 16.11) have been studied with respect to field data (mapping, paleocurrent measurements using orthocone nautilids), distribution and structure of the mounds, biota and microfacies (Brachert et al. 1992; Hüssner 1994) and with regard to isotope geochemistry (Belka 1998; Mounji et al. 1998). The paleoenvironment of the mounds is interpreted as deep-water, below fair-weather wave base but within the range of major storms. The absence of algae and stromatoporoids indicate non-photic conditions. Different genetic interpretations have been discussed, depending on the weighting of these criteria: • The model derived from the microfacies approach favors mound formation by a self-sustaining interplay of physically induced changes in the environment (bottom currents, storms) and intrinsic biological factors (preferred accumulation of skeletal material in the mound area because of the increased growth of crinoids due to better nutrient access). The mounds developed on shoals originating from the transport of sediment by storms. Carbonate mud production is explained by the degradation of skeletons. Stabilization of mound flanks is explained by rapid synsedimentary marine cementation starting with the early cementation of auloporid corals. The asymmetry of the mounds is interpreted as a result of shaping by bottom currents. • Geochemically-based models explain the formation of microcrystalline carbonate in the mounds by hydro-
Waulsortian Mud Mounds
thermal venting related to thermal flux above the volcanic massif (Mounji et al. 1998) or by carbonate precipitation from brines comprising a mixture of hydrothermal fluids and seawater, associated with the aerobic oxidation of thermogenic methane (Belka 1998). Inorganic carbonate mud production might be possible. Most mounds developed over crosspoints of radial and tangential faults resembling a network of thermally induced joints. These faults could have served as conduits for the migration of hydrothermal fluids, which might have had a positive effect on providing a nutritional source.
16.2.6.2 Waulsortian Mud Mounds: Early Carboniferous (Lower Mississippian) Muleshoe Mound, Sacramento Mountains, New Mexico, U.S.A. The Muleshoe Mound near Alamogordo in New Mexico is a well-studied example of a ‘Waulsortian’ mound (Pl. 144). Waulsortian banks and mounds (named after Waulsort, a locality in Belgium) with tabular, knolland sheet-like geometries developed during midDinantian (Late Tornaisian to Early Viséan) times in many parts of the world, mainly in central Europe and in North America. Mounds and banks are lime mud-dominated. Carbonate muds are ‘ polymuds’, e.g. they include both automicrites and allomicrites. Most authors assume a strong microbial contribution in mud formation. Diagnostic microfacies criteria are skeletal grains, nonskeletal grains (predominantly peloids), a wackestone and packstone matrix, and coarse sparry fabrics within masses of various size and forms (some of which correspond to stromatactis). The facies is characterized by three end-members: Fenestrate bryozoan fronds buried in wackestone or coated with spar, wackestones with or without ‘sparry masses’ (some open-space structures) and crinoid-rich packstones. Major biota, in order of decreasing volumetric importance, are crinoids, fenestrate and other bryozoans, brachiopods and mollusks. Note that the dominance of specific groups (e.g. crinoids and bryozoans in the Waulsortian mounds) is not necessarily a proof that intensified growth of these organisms significantly influenced mound formation (Ahr and Stanton 1994). Calcareous algae are almost absent, but microbial filaments may be present. Simple foraminifera are ubiquitous, but plurilocular forms are absent or rare and restricted to the shallow upper aphotic and the photic parts of the
Permian Reef Complex
mounds. The Waulsortian mounds grew in tropical, relatively deep subphotic to shallow photic environments on ramps and within basins. Lees and Miller (1985, 1995) distinguished four depth-related phases on the basis of grain associations, whereby the deepest phase A is characterized by an association consisting of crinoids and fenestrate bryozoans, and the shallowest phase D by micritized skeletal debris and cryptalgal coating. Organism diversity increases markedly with decreasing water depth. Peloids and intraclasts tend to be concentrated in more shallow phases. Correspondence analysis demonstrates that the phases correspond to relays and are part of a continuum (see Sect. 14.4). Mud mound formation is explained as starting with biofilms triggering CaCO3 precipitation.
16.2.6.3 Reefs: The Capitan Reef, Permian Reef Complex, Guadalupe Mountains, Texas and New Mexico, U.S.A. The Guadalupe and Delaware Mountains contain some of the finest outcrops of reef and reef-related rocks in the world. Original facies relations are preserved from extensive erosional modifications due to rapid deposition of evaporites at the end of Guadalupian time. The Permian Basin region of west Texas and New Mexico provides an opportunity to study the interrelationships of depositional facies, sequence stratigraphy, diagenetic patterns, and the development of carbonate hydrocarbon reservoir rocks in detail. Depositional facies belts include a shelf-interior facies comprising red beds and evaporites deposited on sabkha-like plains, pisolitic deposits (see Pl. 14/1, Pl. 126/1), a shelf facies consist-
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ing of skeletal grain- and packstones (Pl. 145/ 4) and marginal reef limestones (Pl. 145/1-3; Fig. 16.12), and a slope and basin facies. The Capitan Reef is an excellent example of an ancient prograding reef located at the edge of a carbonate shelf. Basics: Understanding reef carbonates Belka, Z. (1998): Early Devonian Kess-kess carbonate mounds of the eastern Anti-Atlas, Morocco) and their relation to submarine hydrothermal venting. – Journal of Sedimentary Research, 68, 368-377 Bolton, K., Lane, H.R., LeMone, D.V. (1982): Symposium in paleoenvironmental setting and distribution of Waulsortian facies. – 202 pp., El Paso (El Paso Geological Society and The University of Texas) Brachert, T.C., Buggisch, W., Flügel, E., Hüssner, H.M., Joachimski, M.M., Tourneur, F., Walliser, O.H. (1992): Controls of mud mound formation: The Early Devonian Kess-Kess carbonates of the Hamar Laghdad, Antiatlas, Morocco. – Geologische Rundschau, 81, 15-44 Camoin, G.F., Davies, P.J. (eds., 1998): Reefs and carbonate platforms in the Pacific and Indian Oceans. – International Association of Sedimentologists, Special Publication, 25, 440 pp. Fagerstrom, J.A. (1987): The evolution of reef communities. – 600 pp., New York (Wiley Interscience) Fagerstrom, J.A. (1991): Reef-building guilds and a checklist for determining guild membership. – Coral Reefs, 10, 47-52 Fagerstrom, J.A., Bradshaw, M.A. (2002): Early Devonian reefs at Reefton, New Zealand: guilds, origin and paleogeographic significance. – Lethaia, 35, 35-50 Fagerstrom, J.A., Weidlich, O. (1999): Strength and weakness of the reef guild concept and quantitative data: application to the Upper Capitan-Massive community (Permian), Guadalupe Mountains, New Mexico-Texas. – Facies, 40, 131-156 Flügel, E. (1989): ‘Algen/Zement’-Riffe. – Archiv für Lagerstättenforschung, Geologische Bundesanstalt Wien, 10, 125-131 Flügel, E., Flügel-Kahler, E. (1992): Phanerozoic reef evolution: Basic questions and a data base. – Facies, 26, 167-278
Fig. 16.12. Permian Reef Complex. A: Reefbuilders. Strongly encrusted inozoan calcareous sponges. Interskeletal voids are filled with acicular cement. B: Reef cement. Former aragonite fans forming botryoids. Walnut Canyon and Slaughter Canyon, Guadalupe Mountains.
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Flügel, E., Kiessling, W. (2002): A new look at ancient reefs. – In: Kiessling, W., Flügel, E., Golonka, J. (eds., 2002): Phanerozoic reef patterns. – SEPM, Special Publications, 72, 3-10 Haywick, D.W., Kopaska-Merkel, D.C. (eds., 2001): Carbonate mounds: sedimentation, organismal response, and diagenesis. – Sedimentary Geology, 145, Issue 3-4, 157414 Hubbard, D.K., Burke, R.B., Gill, I.P. (1998): Where is the reef: the role of framework in the Holocene. – Carbonates and Evaporites, 13, 3-9 Hüssner, H.M. (1994): Reefs, an elementary principle with
Ecologic Reef: Capitan Reef
many complex realizations. – Beringeria, 11, 3-99 Insalaco, E. (1998): The descriptive nomenclature and classification of growth fabrics in fossil scleractinian reefs. – Sedimentary Geology, 118, 159-186 James, N.P., Bourque, P.-A. (1992): Reefs and mounds. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sea level change. – 323-348, Ottawa (Geological Association of Canada) Kiessling, W. (2002): Secular variations in the Phanerozoic reef systems. – In: Kiessling, W., Flügel, E., Golonka, J. (eds.): Phanerozoic reef patterns. – SEPM, Special Publication, 75, 625-690
Plate 145 Ecologic Reef or Something Else? The Capitan Reef, (Guadalupe Mountains), Texas and New Mexico The origin of the reef rocks and the ecologic interpretation of the organisms of the famous Capitan reef have produced a long ongoing debate. The Late Permian Capitan Limestone consisting of a massive/unbedded reef limestone (Capitan-Massive Member) and a basinward dipping, bedded biodetrital member (Capitan Breccia Member) is well exposed in a narrow belt in the southern Guadalupe Mountains and continues in the subsurface around the northern margin of the Delaware Basin (Harris and Groover 1989). The formation of the Massive Member of the Capitan Formation has been interpreted in various ways, including a barrier reef, a deep-water skeletal wackestone mound and a cement boundstone mound. Principal constituents of the reef framework were framebuilding heterotroph organisms (sponges, Tubiphytes, bryozoans), encrusting organisms (Archaeolithoporella and microbialite), infilling internal sediment and an amazingly abundant mass of synsedimentary marine cements (–> 1; Fig. 16.12B). The most important constructional element of the reefs is a micro-framework (e.g. Tubiphytes, Archaeolithoporella, sponges), a consortium of low-growing reefbuilders (coralline sponges, Fig. 16.12A) and marine-phreatic cement (Weidlich and Fagerstrom 1999). Open surface communities were probably less common than cryptic communities (Wood 1996). The reef was constructed in part by microbially bound sediments enclosing growth cavities formed by fenestellid bryozoan fronds and laminar sponges and strengthened by extensive synsedimentary carbonate cements and biogenic encrustations. Postmortem biogenic encrustations and extensive early cementation are regarded as major controls on the growth and preservation of the reef structure. Some of the sponges were ‘upside down’ hanging in pendant fashion within crypts. Most reef growth is believed to have taken place not in shallow but in deeper water below the wave base (Saller et al. 1999). 1 ‘Cementstone’. Marine cement is volumetrically the most significant component through much of the Massive-Member of the Capitan Formation. The rock consists predominantly of submarine, originally aragonitic radial-fibrous cement (C) formed synchronously with biogenic crusts (Archaeolithoporella, arrows) growing on sponges (SP) as well as on cements. Note the botryoids composed of cement fans. Botryoids may comprise more than 90% of reef rock samples. The cementation took place at or near the shelf margin and has greatly reduced primary porosity. Late Permian (Upper Capitan Limestone, Capitan-Massive Member): Mouth of Walnut Canyon. 2 Carbonate cement and encrusters. Synsedimentary intergrowth of Tubiphytes Maslov (see Sect. 10.2.6.2) and originally aragonitic botryoidal cement within a growth cavity. Reef biota exhibit distinct internal zonation patterns, reflecting changing environmental conditions passing up the slope and onto the shelf. Same locality as –> 1. 3 Cement. Botryoidal carbonate cement forms a great part of the reef rock. The thin section shows primary fibers of calcite paramorphic after marine aragonite. Note the plane straight sides of the crystals (see Fig. 16.12B). On the basis of analogies with Holocene reef cements, this calcite cement is interpreted as the diagenetic equivalent of a submarine, aragonite precursor (see Sect. 7.1.5.2). Massive marine cementation is largely confined to the reef facies and the immediate nearback-reef grainstones. Same locality as –> 1. 4 Backreef facies. Recrystallized green algal packstone. Most grains are dasyclads (Mizzia; arrow points to a longitudinal section of Connexia, C, an alga known up to now only from the western Tethyan Permian) and peloids, associated with a few gastropods (bottom left). Mizzia grainstones, packstones and wackestones occur in a backreef position (outer shelf lagoon) immediately shelfward of the Capitan reef. The stands of the dasyclads probably formed in hypersaline lagoons. Mizzia limestones represent a worldwide Middle and Late Permian shelf facies. Standard Microfacies SMF 18-DASY. Late Permian (Tansill Formation, coeval with the Uppermost Capitan Formation): Guadalupe Mountains, Texas, U.S.A.
Plate 145: Capitan Reef, Guadalupe Mts.
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Permian Reef Complex
Fig. 16.13. Permian Reef Complex. A: West side of the southern Guadalupe Mountains. West Side Scarp. El Capitan at the southern end of the Guadalupe Mountains and most of the Guadalupe Peak represent the upper foreslope facies consisting of very thickly bedded ‘forereef’ talus extending far into the basin. The Capitan-Massive reef facies is exposed in the upper part of the Guadalupe Peak and in the adjacent cliff walls toward the Shumard Peak. ‘Back-reef’ facies occurs at the top of the Guadalupe Peak. The brown sediments in the foreground below the El Capitan are the Cherry Canyon Formation and the Brushy Canyon Formation. These units are basinal, predominantly siliciclastic sediments with tongues of reef-derived calcarenites, indicating the progradation of the Capitan platform margin over the sediments of the Delaware Basin. Width of the picture about 5 km. Photograph courtesy of W. Buggisch (Erlangen). B: Drawing by Norman Newell from Newell et al. (1953). Kiessling, W., Flügel, E. (2002): Paleoreefs - a database on Phanerozoic reefs. – In: Kiessling, W., Flügel, E., Golonka, J. (eds., 2002): Phanerozoic reef patterns. – SEPM, Special Publication, 75, 77-92 Kiessling, W., Flügel, E., Golonka, J. (2000): Fluctuations in the carbonate production of Phanerozoic reefs. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems: components and interactions. – Geological Society of London, Special Publication, 178, 191-215
Kiessling, W., Flügel, E., Golonka, J. (eds., 2002): Phanerozoic reef patterns. – SEPM, Special Publication, 75, 775 pp. Kirkland, P.L., Dickson, J.A.D., Wood, R.A., Land, L.S. (1998): Microbialite and microstratigraphy: the origin of encrustations in the middle and upper Capitan Formation, Guadalupe Mountains, Texas and New Mexico, U.S.A. – Journal of Sedimentary Research, A68, 956-969 Kiessling, W., Flügel, E., Golonka, J. (eds., 2002): Phanero-
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Lost Platforms
zoic reef patterns. – SEPM, Special Publications, 72, 774 pp. Kopaska-Merkel, D.C., Haywick, D.W. (2001): Carbonate mounds: sedimentation, organismal response, and diagenesis. – Sedimentary Geology, 145, 157-159 Lees, A., Miller, J. (1995): Waulsortian banks. – In: Monty, C.L.V., Bosence D.W.J., Bridges, P.H., Pratt, B.R. (eds.): Carbonate mud mounds. Their origin and evolution. – International Association of Sedimentologists, Special Publication, 23, 191-271 Monty, C.L.V., Bosence, D.W.J., Bridges, P.H., Pratt, B.R. (eds.): Carbonate mud mounds. Their origin and evolution. – International Association of Sedimentologists, Special Publication, 23, 537 pp. Mounji, D., Bourque, P.A., Savard, M.M. (1998): Hydrothermal origin of Devonian conical mounds (Kess-Kess) of Hamar Lakhdad ridge, Anti-Atlas, Morocco. – Geology, 26, 1123-1126 Noé, S. (2003): Spätstadium einer sterbenden Karbonatplattform: Schelfrand- und Außenschelf-Entwicklung der Tansill-Formation (Permian Reef Complex, New Mexico, USA). – Kölner Forum für Geologie und Paläontologie, 11, 1-254 Piller, W., Riegl, B. (2003): Vertical versus horizontal growth strategies of coral frameworks (Tulamben, Bali, Indonesia). – International Journal of Earth Sciences, 92, 511519 Riding, R. (2002): Structure and composition of organic reefs and carbonate mud mounds: concepts and categories. – Earth-Science Reviews, 58, 163-231 Saller, A.H., Harris, P.M., Kirkland, B.L., Mazzullo, S.J. (eds., 1999): Geologic framework of the Capitan reefs. – SEPM, Special Publication, 65, 224 pp. Stanley, G.D. (ed., 2001): The history and sedimentology of ancient reef systems. – 458 pp., New York (Kluwer Academic/Plenum Publishers) Stanton, R.J., Flügel, E. (1989): Problems with reef models: the late Triassic Steinplatte ‘reef’ (Northern Alps, Salzburg/Tyrol, Austria). – Facies, 20, 1-138 Webb, G.E. (1996): Was Phanerozoic reef history controlled by the distribution of non-encymatically secreted carbonate / microbial carbonate and biologically induced cement)? – Sedimentology, 43, 947-971 Weidlich, O., Bernecker, M., Flügel, E. (1998): Combined quantitative analysis and microfacies studies of ancient reefs: an integrated approach to Upper Permian and Upper Triassic reef carbonates (Sultanate of Oman). – Facies, 28, 115-144 Wendt, J., Belka, Z., Kaufmann, B., Kostrewa, R., Hayer, J. (1997): The world’s most spectacular carbonate mud mounds (Middle Devonian, Algerian Sahara). – Journal of Sedimentary Research, 67, 424-436 Wood, R. (1995): The changing biology of reef-building. – Palaios, 10, 517-529 Wood, R. (1999): Reef evolution. – 414 pp., Oxford (Oxford University Press) Wood, R. (2001): Are reefs and mud mounds really so different? – Sedimentary Geology, 145, 161-171 An excellent web site on the Permian Reef Complex describing a virtual field trip and showing the present state of the art is offered by Peter Scholle: geoinfo.nmt.edu/staff/scholle/guadalupe.hmtl. Further reading: K191, K213
16.3 Fingerprinting Lost Platforms One of the most exciting topics of microfacies studies is the reconstruction of depositional environments that are only recorded by eroded relicts of former carbonate platforms and reefs. This approach is different than the microfacies analysis of mass flow deposits occurring adjacent to preserved platforms (Sect. 15.7.2) because the former platform is already lost. Most of the relicts of these lost platforms are preserved within debris flow and other mass flow deposits intercalated in deep shelf, deep ramp or basinal sediments and forming tongues and wedges of limestone conglomerates and breccias. Significance: The reconstruction of lost platforms based on the provenance analysis of erosional products of platform material offers possibilities for • recognizing sequence-stratigraphic breaks connected with major sea-level changes leading to the destruction of platforms, • reconstructing the paleogeography of tectonically dislocated regions.
16.3.1 Methods A study of these allochthonous deposits for reconstructing lost platforms requires the following steps: Field study • Stratigraphic sections exhibiting the occurrence and geometry of the allochthonous intercalations. • Measurements of the distribution, frequency, size and shape of the limestone pebbles and boulders. The clasts should be differentiated with regard to their lithology, rock color and abundance of individual types (see Sect. 5.3.3.2). • Using grids and point counting results in data that can be used in estimating the proximal versus the distal position of the depositional area with respect to the source area. The abundance of the same clast types exhibiting variable shapes may indicate a more proximal position, rare and small clast types a more distal position. • Sampling should consider lithologic variability and statistical distribution of clasts (see Sect. 3.1.2.1). Laboratory study • Age determination is one of the main targets in the reconstruction of lost depositional environments. Thin sections should be evaluated with regard to agediagnostic fossils, e.g. foraminifera or algae. Paleozoic deposits have been successfully dated by mi-
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•
•
crofossils found in acid residues, e.g. conodonts or silicified foraminifera. Microfacies criteria and microfacies types of boulders should be typified using large thin sections in order to get an idea of microfacies variation and biotic diversity. Unlike common microfacies studies, grains and fossils should be described in great detail, both qualitatively and quantitatively. These details facilitate a comparison of boulders and the differentiation of groups characterized by similar microfacies. Assignment of the microfacies seen in thin sections of boulders to Standard Microfacies Types might assist in recognizing rough facies zones, but could be dangerous because of circular thinking. Use of the zones distinguished in the Wilson model will lead to platform environments and ooid sand bars will be placed near platform margins. See Sect. 14.3.4 for the differences when using ramp models. The original depositional areas on former shelves, ramps or platforms may be derived from the estimation of environmental conditions. Water energy and circulation, and substrate are reflected by matrix, grains and biota (see Chap. 12). These criteria should be used in evaluating water depths, wave base and storm wave base.
Reconstructing Lost Platforms
16.3.2 Case Studies The following case studies demonstrate the methods of using microfacies for reconstructing depositional environments that are only recorded by clasts and pebbles.
16.3.2.1 Case Study: Platform Facies Patterns Derived from the Microfacies of Early Carboniferous Conglomerates (Southern Spain) Herbig (1984) studied limestone boulders occurring in the post-Bashkirian Marbella Formation outcropping in the Betic Cordillera of southern Spain. The Marbella Formation represents deep-water sediments, predominantly conglomerate lenses incised in shales. The lower part of the formation is characterized by oligomict limestone conglomerates whose clasts can be dated with conodonts and benthic foraminifera as Late Viséan in age. A thorough analysis of the boulders, based on thin sections and acid residue data indicates a Late Viséan age (zone 16 of the foraminiferal zonation) of the limestones recorded by the boulders in the conglomerates. This allows the use of the microfacies types observed in the boulders to reconstruct a relatively short phase of Early Carboniferous shelf sedimentation. Pl. 146
>>> Fig. 16.14. Reconstruction of the Late Viséan shelf derived from the microfacies analysis of limestone boulders of the Marbella Formation near Malaga and Granada, southern Spain. Modified from Herbig (1984). The model shows a highly differentiated, shallow carbonate shelf embracing the Standard Facies Zones FZ 8 (restricted platform interior), FZ 7 (open-marine platform interior), FZ 5/6 (platform-margin sand shoals and reefs) and FZ 4 (slope). The spatial arrangement of the microfacies types is based on joint grain associations and joint biota, e.g. the close proximity of microfacies b1 and b2 is indicated by the joint occurrence of ooids. Rare ooids in the oncoid rudstone of microfacies c1 point to a proximity of this microfacies type to microfacies b1 and b2. The sketch displays common matrix and grain types as well as the distribution of facies-diagnostic fossils. Pseudodonezella is a dasyclad green alga (see Pl. 60/1). Fasciella is interpreted as green or red alga (see Pl. 56/3, 4). The name Saccaminopsis refers to a rock-building, generally single-chambered foraminifer. Corals are represented by single and colonial rugose corals and heterocorals (see Pl. 83/2). Microfacies types: Slope. a2 – crinoid grainstone. a3 – limestone breccia consisting of cm-sized clasts representing various microfacies. Ooid and lime sand bars and adjacent areas. b1 – ooid grainstone (Pl. 146/1). b2 – oolitic dasyclad grainstone (Pl. 146/2). b3 – coated crinoid packstone (Pl. 146/3). b4 – skeletal grainstone with diverse biota. Protected and/ or deeper parts of the lime sand areas. c1 – oncoid rudstone (Pl. 146/4). c2 – fine-grained foraminiferal grainstone (Pl. 146/ 5). c3 – skeletal packstone. Luvward and central part of the open-marine shelf lagoon. e1 – grainstone with abundant aggregate grains. e4 – skeletal packstone and grainstone with crinoids and bryozoans. e5 – crinoid packstone. e6 – bioclastic wackestone (Pl. 146/6). Leeward part of the platform interior. f1 – wackestone with micrite clasts. f2 – grainstone with aggregate grains and amalgamated peloids (Pl. 146/8). f3 – grainstone with Saccaminopsis. Platform interior with restricted water circulation. g1 – wackestone with Saccaminopsis. g2 – fine-grained peloidal grainstone. g2 – peloidal grainstone with densely packed peloids (Pl. 146/8). g3 – laminated fenestral mudstone/bindstone (Pl. 146/7). g4 – quartz-bearing mudstone. g5 – floatstone with variously-sized micrite clasts. g6a – quartz-bearing ooid grainstone with radial ooids (Pl. 146/9). Not specifically indicated in the sketch is the position of organic buildups, represented by microfacies d1 (Pseudodonezella bafflestone), widely distributed all over the shelf (see signature of Pseudodonezella), and microfacies d2 (bindstone), common in the restricted Facies Zone 8.
Reconstructing a Paleo-Escarpment
displays some of the microfacies types of the limestone boulders. Fig. 16.14 presents a reconstruction of a lost platform that was attached to the Alboran Block, representing a sedimentary realm before the Variscan orogeny. The paleogeographic relationships of that realm were preserved in the Mesozoic ALKAPECA microplate. The break-up and the dispersal of parts of this microplate led to the formation of isolated Paleozoic terranes (now distributed in Spain, Italy and North Africa) whose paleogegraphic relations become more transparent from microfacies-based clast analyses.
16.3.2.2 Case Study: Reconstruction of PaleoEscarpments from Microfacies Data Di Stefano et al. (2002) gave an example demonstrating how microfacies analysis can be used in reconstructing the conversion of a Bahamian-type carbonate platform into a pelagic escarpment that is not recorded by outcrop data. Diagnostic criteria used are • the change from tidalites to oolites, i.e. from restricted, interior lagoonal to a more open marine sandy depositional environment;
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• the superposition of the oolites by micritized skeletal grainstones and wackestones, and by wackestones with radiolarians and sponge spicules indicating the decrease of platform carbonate productivity, recording a dissection of the platform and the subsequent isolation of a submarine topographic high; • the deposition of crinoidal packstones (encrinites) in times of increasing synsedimentary tectonics reflected by neptunian dikes exhibiting infillings of different microfacies types pointing to a stepped slope. Basics: Reconstructing lost platforms Belka, Z., Skompski, S., Sobon-Podgorska, J. (1996): Reconstruction of a lost carbonate platform on the shelf of Fennosarmatia: evidence from Viséan polymictic debris, Holy Cross Mountains, Poland. – In: Strogon, P., Sommerville I.D., Jones, L.L. (eds.): Recent advances in Lower Carboniferous geology. – Geological Society of London, Special Publications, 107, 305-329 Di Stefano, P., Galacz, A., Mallarino, G., Mindszenty, A., Vörös, A. (2002): Birth and early evolution of a Jurassic escarpment: Monte Kumeta, Western Sicily. – Facies, 46, 273-298 Flügel, E., Herbig, H.-G. (1988): Mikrofazies karbonischer Kalkgerölle aus dem Paläozoikum des Rif (Marokko): Ein Beitrag zur Paläogeographie der westmediterranen Paläotethys im Karbon. – Facies, 19, 271-299
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Flügel, E., Kraus, S. (1988): The Lower Permian Sexten Breccia (Sexten Dolomites) and the Tarvis breccia (Carnic Alps): microfacies, depositional environment and paleotectonic implications. – Memoria della Società Geologica Italiana, 34, 67-90 Herbig, H.-G. (1984): Rekonstruktion eines nicht mehr existierenden Sedimentationsraums – Die Kalkgerölle im
Lost Carbonate Platform
Karbon-Flysch der Malagiden (Betische Kordillere). – Facies, 11, 1-108 Pohler, S.M.L., James, N.P. (1989): Reconstruction of a Lower/Middle Ordovician shelf margin: Cow Head Group, Western Newfoundland. – Facies, 21, 189-262 Further reading: K072, K190
Plate 146 Fingerprinting a Lost Carbonate Platform: Early Carboniferous of the Betic Cordillera, Southern Spain The plate displays microfacies types observed in thin sections of limestone boulders of the Late Viséan Marbella Conglomerate deposited in a deep-water basin. The microfacies of the boulders reflect the microfacies types of a carbonate shelf adjacent to the Paleozoic Alboran Block (comprising the Betic Cordillera, the internal zones of the Rif, Menorca, the Chenoua Massif and the Kabylies). Microfacies criteria and the composition and diversity of biota allow the facies differentiation of an attached platform to be reconstructed (Fig. 16.14). 1 Ooid grainstone. Microfacies b1. The well-sorted sediment consists of tangential ooids. Most ooid cores are crinoid debris. Interpretation: Crinoid material has been transported from shallow, moderate energy environments to areas of high energy forming ooid sand bars. The interpretation is supported by the dominance of tangentially structured ooids and the occurrence of other stenohaline skeletal grains acting as cores for the ooids. 2 Oolitic dasyclad grainstone. Microfacies b2. The well-sorted sediment consists predominantly of dasyclads (predominantly Eovelebitella; see Pl. 59/1) with tangential oolitic coatings and some aggregate grains. Interpretation: Dasyclad material was transported from open shelf lagoon areas to high energy areas forming sand bars. 3 Coated crinoid packstone. Microfacies b3. Most grains are variously sized, poorly sorted cortoids. The matrix is pelsparite. Interpretation: Formed in areas of continuous water energy, at the periphery of lime sand bars. Continuous turbulence is indicated by winnowing, dominance of only a few organism groups that were adapted to shifting substrate, and the disintegration of crinoid elements. 4 Oncoid rudstone. Microfacies c1. The large grain is an oncoid formed by cyanobacteria encrusting a shell. Smaller components are aggregate grains. Interpretation: Protected and deeper parts of sand areas, formed behind ooid bars. This interpretation is supported by the occurrence of large cyanoids with sessile foraminifera, the association of oncoids with aggregate grains and ooids, and the scarcity of biota that are not incorporated within oncoids. 5 Fine-grained foraminiferal grainstone. Microfacies c2. The sediment consists of bioclasts and peloids. Skeletal grains are benthic foraminifera and a few calcareous algae and gastropods. Interpretation: Very shallow, moderately energetic sand areas. The interpretation is based on the composition of the highly diverse foraminiferal fauna comprising more than 50 taxa in some samples of this microfacies. 6 Bioclastic wackestone. Microfacies e6. The sediment is characterized by an inhomogeneous, micritic and siltitic matrix and various broken skeletal grains including bryozoans (top and left), shells and some echinoderms. Interpretation: Central part of the open shelf lagoon. Below storm wave base. A relatively deep environment is deduced from the occurrence of open-marine biota (fish remains, conodonts, juvenile ammonites) and the lack of bioturbation. 7 Laminated fenestral mudstone/bindstone. Microfacies g3. Characteristic features are laterally extended spar-filled fenestrae and densely spaced peloids forming irregular layers. Interpretation: Formed in supra- to intertidal restricted bays, indicated by the fenestral mudstone/bindstone fabrics and the almost total absence of skeletal grains both in thin sections and in acid residues. 8 Microfacies boundary. Erosional contact between a grainstone with aggregate grains, amalgamated peloids and skeletal grains (bottom; microfacies f2) and a fine-grained peloidal grainstone with densely packed peloids (top: microfacies g2). Interpretation: The microfacies indicates fluctuations of open and restricted circulation. Microfacies g2 is characterized by biota pointing to higher salinities. Shelf lagoon. Open circulation and well-lighted conditions (microfacies f2) are deduced from common burrowing and the abundance of algae. Restricted conditions (microfacies g2) are indicated by low-diversity biota dominated by a few taxa adapted to salinity fluctuations (e.g. Fasciella, see Pl. 56/4; calcispheres). Micrite laminae point to microbial binding of the sediment. 9 Quartz-bearing poorly-sorted ooid grainstone with radially structured ooids. Microfacies g6a. The irregularly shaped radial ooids are broken and regenerated. The cores of the ooids are terrigenous quartz grains. Interpretation: The abundance of regenerated radial ooids as well as the scarceness of biota point to a protected, hypersaline environment. The quartz grains indicate relatively near-coast areas.
Plate 146: Reconstructing a Lost Carbonate Platform
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16.4 Recognizing Ancient CoolWater Carbonates The criteria and constraints of recent shelf and reef carbonates were discussed in Sect. 2.4.4 (see Pl. 4). A promising method based on grain association analysis to distinguish cold-water from warm-water limestones was described in Sect. 12.2. The present chapter summarizes the diagnostic criteria of ancient cold-water carbonates and presents a case study.
16.4.1 Microfacies Criteria of Non-Tropical Cold-Water Shelf and Reef Limestones The major differences of tropical warm-water carbonates and nontropical cold-water carbonates are listed in Tab. 2.2. Ancient cold-water shelf carbonates are characterized by (1) compositional criteria expressed
Recognizing Cool-Water Carbonates
by characteristic grain associations and supported by paleontological data, and (2) mineralogical and diagenetic criteria. Our knowledge of ancient cold-water limestones is predominantly based on the study of Ordovician, Carboniferous and Permian as well as Cenozoic shelf carbonates, and Tertiary reef carbonates. Compositional criteria. The common compositional criteria of cold-water shelf limestones are the abundance of sand- and gravel-sized bioclasts, sometimes associated with litho- and intraclasts and peloids. Carbonate mud formed by physical and chemical degradation of skeletons or pelagic admixtures occurs in different quantities. Biodegradation of skeletal grains is high. Some grains, characteristic of shallow-marine warm-water carbonates, e.g. ooids and aggregate grains, are almost or completely absent.
Fig. 16.15. Cool-water bryozoan limestone. During the Late Cretaceous and Cenozoic cool-water bryozoan mounds formed in mid- to high-latitudinal settings. An impressive example are extended belts of asymmetric bryozoan mounds consisting of skeletal floatstones, packstones and wackestones. The mounds originated below the photic zone on the outer shelf and upper slope. Composition, facies types and geometry of the mounds exhibit great similarity with Neogene cool-water carbonates of the Great Australian Bight (Surlyk 1997). The sample is a skeletal packstone consisting predominantly of cyclostome bryozoans. Early Tertiary (Early Danian): Stevns Klint, Denmark.
Cool-Water Limestone
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Fig. 16.16. Cool-water coral limestone with deep-water constructional azooxanthellate corals and hydrozoans. The corals occur within mixed carbonate-siliciclastic deposits, including an up to 8 m thick limestone consisting of a lower floatstone unit (interpreted as coral thicket) and an upper bioclastic rudstone unit (interpreted as a coral bank). The sample from the coral bank exhibits slender branching scleractinian corals represented by Dendrophyllia (D) and Oculina (O) and stylasterinid hydrozoans (S) with still preserved gastropores (white arrow). Marine isopachous fibrous cements form only thin rims (black arrows). Most interparticle pores are filled with a bright meteoric-phreatic cement. A deep- and cool-water setting of the coral bank is postulated from the Dendrophyllia-Oculina-community (today flourishing in water depths of about 100 m), low diversity of corals, presence of stylasterinid hydrozoans, presence of planktonic foraminifera, absence of algae, and the inferred presence of deep currents, which might have provided the nutrient supply necessary for azooxanthellate corals. Tertiary (Calcare di Mendicino, Late Miocene): Near Cosenza, Calabria, Italy. After Mastandrea et al. (2002).
Abundant skeletal grains are bryozoans (Fig. 16.15), mollusk shells particularly high-diversity bivalves, benthic low-diverse foraminifera, echinoderms, and in Cenozoic limestones barnacles and rhodophycean algae as well. Azooxanthellate and some zooxanthellate corals started building deep-and cold-water reefs during the Tertiary (Fig. 16.16; Pl. 147). Further groups include brachiopods, serpulid worms and siliceous sponges recorded by spiculites. These fossils contributed to the formation of shelf carbonates characterized by specific skeletal grain associations (Sect. 12.2.1). The associations have been particularly well studied in Tertiary temperate and cool-water carbonates (e.g. Hayton et al. 1995) and used in recognizing paleoclimate changes during the Cenozoic but also during the
Late Paleozoic (Brachert et al. 1993; Beauchamp and Desrochers 1997; Hüneke et al. 2001). Bioerosion, maceration and dissolution of shells may cause considerable taphonomic loss in the fossil record. Important and abundant groups of modern coldwater environments are not known or only rarely or indirectly known from the fossil record because of the absence or rareness of skeletons. Examples are sea weeds (brown algae) that form extended kelp forests in recent high-latitude cold seas, soft sponges, and octocorals. Since these organisms provide substrates for epizoans and epiphytes, the existence of sea weeds, for example, can be derived from the analysis of attached organisms.
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Mineralogical and diagenetic data. Cold-water carbonates are dominated by skeletal grains consisting of Low-Mg calcite (bryozoans, many bivalves) and/or High-Mg calcite (coralline algae, echinoderms). Aragonitic skeletons are rare and usually recorded by molds. Authigenic glauconite and phosphorite grains are common features of cold-water carbonates.
Cool-Water Limestones
A striking feature of carbonate cold-water sediments is their retarded and weak cementation and lithification (Rao 1994). Diagenesis is predominantly destructive. The diagenetic potential is low, but intraskeletal and interskeletal cementation occurs in temperate and cold-water carbonates (Fig. 16.16). Mg-calcite cements dominate in cold-water reefs.
Plate 147 Cool-Water Coral Reef Limestones (Early Tertiary of Fakse, Denmark) The Fakse quarries south of Copenhagen exhibit lowermost Tertiary limestones comprising bryozoan limestones (–> 3) and coral limestones (–> 1). These limestones differ clearly in faunal criteria. The bulk of the grains of the bryozoan facies consists of cheilostome and cyclostome bryozoans with a greatly varying bryozoan/matrix ratio; colonial corals and hydrozoans are missing; octocorals are represented by large colonies, and the associated fauna is low-diverse. The coral facies comprises several subfacies characterized by the dominating coral taxa, including the Dendrophyllia subfacies (–> 1) and the Faksephyllia subfacies (–> 4). Framebuilding organisms were azooxanthellate, ahermatypic corals, stylasterinid hydrozoans (–> 5) and fan-shaped octocorals (–> 6). Associated organisms include bryozoans, bivalves and gastropods, decapod arthropods, brachiopods, echinoderms, serpulids and siliceous sponges. Borings within corals and shells are abundant. The diversity of the framebuilding community is low, that of the associated fauna significantly high. Most of the sediment within the mounds and surrounding the mounds is parautochthonous and consists of broken corals and the debris of other skeletons (–> 5, 6). A cold-water setting for the coral mounds is indicated by the low-diversity azooxanthellate coral community whose generic composition corresponds to that of modern cool-water coral associations, the dominance of colonial corals with arborescent to dendroid corallites and fragile branches, breakdown of coral framework predominantly by bioerosion, occurrence of open-marine planktonic organisms within the micritic sediments of the mound and intermound facies, interfingering of coral mound-bearing strata with pelagic sediments, the predominance of heterotrophic organisms and the complete absence of algae. 1 Coral limestone. The bulk of the mounds consists of micritic limestones with azooxanthellate, dendroid scleractinian corals. Facies types are distinguished by the dominance of specific coral taxa. The Dendrophyllia (D) facies forms together with the Faksephyllia facies (–> 4) the core of the asymmetrical mounds. The originally aragonitic corals are converted to Low-Mg calcite. Note open framework porosity (white arrow) and bioturbation (black arrows) indicating firmground substrate consistency and slow lithification of the lime mud. The bioclastic debris within the matrix was formed by the breakdown of coral skeletons caused by strong bioerosion. 2 Preservation dependent on skeletal mineralogy. Cross section of the azooxanthellate coral Dendrophyllia candelabrum Hennig, encrusted by different bryozoans (B). Note the good preservation of the calcitic bryozoan and the poor preservation of the originally aragonitic coral. The matrix is micrite. 3 Bryozoan limestone. The base of the mounds was built by bryozoan limestone representing a typical ‘bryomol’ facies (see Pl. 106/2, 3). The chalky, poorly lithified bioclastic packstone yields abundant cyclostome and cheilostome bryozoans (white arrows) as well as octocorals (Moltkia isis Steenstrup, black arrows). Note the high open intra- and interskeletal porosity (white areas). 4 Faksephyllia facies. Some corallites of the dendroid colonies are partly dissolved (white areas). Note the fine-grained bioclastic debris. 5 Hydrozoans. Longitudinal and cross sections of hydrozoans. Fan-shaped stylasterinid hydrozoans are associated with scleractinian corals. The branches of the bushy aragonitic skeletons exhibit ramified canals appearing as pores of different size and function. Feeding polyps project through larger gastropores (arrow), prey-gathering polyps occupy dactylopores. 6 Octocorals. Isolated internode of the octocoral Moltkia lyelli Nielsen, exhibiting the characteristic axial channel as well as two exterior calyces. Modern octocorals occur in tropic and non-tropic oceans. They include sea pens, soft corals, sea fans, organ pipe corals and precious corals. The skeleton is commonly internal, spicular, horny or calcitic and aragonitic. Fossil records comprise calcareous spicules as well as axis structures consisting of horny nodes and calcitic internodes. 7 Fine-grained sediment. Most of the sediment within the mounds is parautochthonous and is produced by the breakdown and deposition of corals without transport (left). Allochthonous sedimentation at the slope of the mounds (grain-supported coral debris, right) is uncommon. –> 1, 4, 6, 7: Bernecker and Weidlich (1990)
Plate 147: Cool-Water Coral Reef
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Cold-water reefs. Biotas that may have been important in ancient temperate-water reefs include calcareous red algae, encrusting foraminifera, bryozoans, serpulids and sabellariids, spiculate demosponges, fasciculate ahermatypic and hermatypic corals, vermetid gastropods, and bivalves (e.g. oysters). These groups form reefs within ramp and platform environments, characterized by the low diversity at a species, family or phylum level, the abundance of stress-tolerant taxa (e.g. cyanobacteria, bryozoans), arrested ecologic successions (dominated by pioneering taxa), the absence of distinct zonation patterns and often limited dimensions by comparison to tropical reefs (Copper 1994).
16.4.2 Case Study: Early Tertiary CoolWater Coral Reef Under ‘Basics’ for this section and the keyword K021 (see CD) one can find important references dealing with ancient cold-water carbonates from the Ordovician, Carboniferous and Permian, Cretaceous and Tertiary. Examples of cold-water carbonates from the Ordovician are shown on Pl. 106/1, from the Early Permian on Pl. 106/2, and from Tertiary temperate carbonates on Pl. 106/3 and Pl. 107/1. Pl. 147 shows the microfacies of Early Tertiary coldwater coral reef limestones from Denmark. These limestones represent one of the finest examples of ancient cool-water coral reefs. The Fakse quarry and the nearby Stevns Klint locality south of Copenhagen, Denmark, are classical localities for studying the Cretaceous-Tertiary boundary and its events as well as understanding cold-water mounds and reefs (Pl. 147). In this area uppermost Maastrichtian bryozoan chalky wackestones forming low asymmetrical mounds are overlain by the just centimeter-thick Danian Fish Clay, which passes upward into the Cerithium Limestone exhibiting a prominent hardground at the top. The surface of this hardground is overlain by bryozoan mound complexes of bryozoan floatstone and packstone (Fig. 16.15) with black flint layers (Surlyk and Damholt 2001). These bryozoan mounds as well as ecologically zoned coral mounds were formed below the photic zone on the deep-water shelf of the Danish-Polish Trough. The water depth is assumed to have been 100 to 300 m (Surlyk 1997). The coral mounds with a height up to 30 m, a maximum length of 200 m, and a width of 80 m exhibit an asymmetrical shape that was controlled by paleocurrent directions (Bernecker and Weidlich 1980). The carbon-
Early Tertiary Cool-Water Reef
ate sediment of the mounds is fine-grained. It consists predominantly of bioclastic framestones, packstones and wackestones. Framebuilders are only azooxanthellate organisms (dendroid and arborescent scleractinian corals; fanlike hydrozoans and octocorals). The breakdown of skeletons due to strong bioerosion as well as the baffling effects of coral colonies were responsible for the accumulation of micritic carbonate sediment. Early diagenesis is responsible for arrested marine-meteoric cementation. During late diagenesis the aragonite of the skeletons was dissolved and replaced by calcite (Weidlich and Bernecker 1991). These reefs have been compared with modern coral ‘lithoherms’ formed by ahermatypic corals in the Straits of Florida (Neumann et al. 1977) and off the Little Bahama Bank (Mullins et al. 1981) in deep-water settings between 600 and 700 m, with strong bottom currents providing oxygen and nutrients for the biota. Similarities to Fakse include growth of current-oriented coral thickets, strong sediment trapping and baffling, and extensive bioerosion. Basics: Ancient cold-water carbonates Beauchamp, B. (1994): Permian climatic cooling in the Canadian Arctic. – Geological Society of America, Special Paper, 288, 299-346 Beauchamp, B., Desrochers, A. (1997): Permian warm- to very cold-water carbonates and cherts in Northwest Pangaea. – In: James, N.P., Clarke, A.D. (eds.): Cool-water carbonates. – SEPM, Special Publication, 56, 327-347 Bernecker, M., Weidlich, O. (1990): The Danian (Paleocene) coral limestone of Fakse, Denmark: a model for ancient, aphotic, azooxanthellate coral mounds. – Facies, 33, 103138 Brookfield, M.E. (1988): A mid-Ordovician temperate carbonate shelf: the Black River and Trenton Limestone Groups of Southern Ontario, Canada. – Sedimentary Geology, 60, 137-153 Carannante, G., Esteban, M., Milliman, J.D., Simone, L. (1988): Carbonate lithofacies as paleolatitude indicators: problems and limitations. – Sedimentary Geology, 60, 333346 Hayton, S., Nelson, C.S., Hood, S.D. (1995): A skeletal assemblage classification system for non-tropical carbonate deposits based on New Zealand Cenozoic limestones. – Sedimentary Geology, 100, 123-141 Hünecke, H., Joachimski, M., Buggisch, W., Lützner, H. (2001): Marine carbonate facies in response to climate and nutrient level: the Upper Carboniferous and Permian of Central Spitsbergen (Svalbard). – Facies, 45, 93-136 James, N.P., Clarke, A.D. (eds., 1997): Cool-water carbonates. – SEPM, Special Publication, 56, 440 pp. Mastandrea, A., Muto, F., Neri, C., Papazzoni, C.A., Perri, E., Russo, F. (2002): Deep-water coral banks: an example from the ‘Calcare di Mendicino’ (Upper Miocene, Northern Calabria, Italy). – Facies, 47, 27-42 Willumsen, M.E. (1995): Early lithification in Danian azooxanthellate scleractinian lithoherms, Fakse quarry, Denmark. – Beiträge zur Paläontologie, 20, 123-131 Further reading: K021
Diagnostic Criteria: Seep and Vent Carbonates
16.5 Testing for Ancient Vent and Seep Carbonates Carbonate beds, lenses, mounds (chemoherms) or nodules, occurring within a strongly differing facies (e.g. shales) and yielding specific benthic faunas (predominantly bivalves, tube-worms) and conspicuous amounts of carbonate cement may represent authigenic ‘seepand vent-carbonates’ (see Sect. 2.4.6). There is still much to learn about these limestones, but the evidence for carbonates bearing products of chemosynthesis related to submarine fluid and gas expulsions is growing in the Phanerozoic record. Isolated and areally limited limestones within deep-marine siliciclastic sediments, which are often considered as resedimented blocks, should be tested with regard to this interpretation. Authigenic carbonates and benthic invertebrate communities that cluster around gas, cold brine, petroleum seeps and hot hydrothermal vents are increasingly being reported from the Phanerozoic rocks (Box 16.5). Seep and vent carbonates offer the opportunity to study the effects of bacterial transformation of chemical compounds from the Earth’s interior into energy in shallow and deep sea environments. The recognition of the vent origin of ancient carbonate mounds is important for understanding hydrocarbon formation and mineralizations (Bitter et al. 1992). Pb-Zn-Fe-Ba-Sr mineralizations in Early Carboniferous mounds of southwestern Newfoundland, formerly interpreted as being of late diagenetic origin, have been reinterpreted as sulfides and sulfates introduced simultaneously with mound construction via vents, and fed through fractures in the basement. 16.5.1 Diagnostic Criteria of Ancient Seep and Vent Carbonates Data used to demonstrate an ancient vent-seep origin of carbonate rocks include regional stratigraphic, sedimentologic, paleontological and geochemical criteria. Microfacies criteria are of major importance. Occurrence, geometry and structural features • Anomalous occurrence of autochthonous limestones in deep-water or high-latitude environments. • Isolated occurrence of limestone bodies within noncarbonate sequences (e.g. siltstones, shales, black shales), or carbonate-poor lithologies (marls). • Mound-shaped geometries with dimensions of a few meters to several tenths of meters. • Localized limestones, related to structural features along which methane or other fluids could migrate (syn-
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sedimentary faults, fractures, joints, diapirs; preserved fluid-flow conduits, e.g. carbonate chimneys, pipes, doughnuts). • Fluid related structures: Micritic breccia (explained by the explosive release of methane), pock mark structures, nodular structures, micritic doughnut structures in calcite or aragonite veins explained by bacterial activity. Fabric and carbonate cements • Layered micritic microbial fabrics. • Nodular fabrics associated with cylindrical to encircling concretions. • Laminated, crustose, or void-filling carbonates, in places extensive and thick brecciated carbonate crusts. • Brecciation within centimeter ranges, resulting in autobreccia. • Microcrystalline, radial-fibrous, splayed fibrous, botryoidal cements and irregular yellow calcite cements. Mg-calcite and aragonite cements. • Hydrocarbon inclusions in carbonate cements. • Cements may be volumetrically far more important than fossils. Rocks consisting only of authigenic carbonate (cementstone) are common (Pl. 148/1). Biota • Stratigraphically limited occurrence of extraordinary (eventually chemoautotrophically-based) benthic biota consisting predominantly of tube-worms, large bivalves, gastropods as well as brachiopods (Campbell and Bottjer 1993). Modern cold seeps are characterized by abundant clams (Calyptogena, Solenomya, lucinids) and mussels (Bathymodiolus), all of them linked to symbiotic bacteria; and the ‘tube worm’ hosting chemolithoautotrophic bacteria. • Predominance of sessile, filter-feeding organisms, with a symbiotic relationship to chemosynthetic bacteria or feeding directly on bacteria. • Anomalous low diversity/high abundance faunas in comparison with faunas of the surrounding sediments. • Commonly high density, but also low density faunas. • Strong endemic character of the fauna. • Strong microbial micritic encrustations on fossils. • Size frequency of shells bimodal, but dominated by large individuals. • Shell beds consisting of in-situ assemblages of strongly dissolved, corroded and variously broken, sometimes pyrite-coated shells. Fragmentation and articulation frequency are major distinguishing criteria of cold seep-related autochthonous shell beds and shell accumulations formed in non-seep environments (Powell et al. 1992).
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• The enclosure of fossils within sedimentary precipitates derived from fluid seepage (e.g. sulfide/sulfate minerals, isotopically distinctive authigenic carbonates). Microfacies • Common microfacies types are bivalve-dominated wackestones, floatstones and packstones; intraclastic packstones; worm tube-dominated boundstones, rudstones and cementstones; peloidal grainstones. Some caution is necessary in using these rock names because most of the textures seen in seep and vent carbonates are diagenetic and not depositional in origin. Depending on the volumetric importance of the microbial contribution, an abundant rock type, composed predominantly of calcite cement, micrite crusts and shells might be designated as cementstone or microbial bindstone. • Micritic coatings and crusts. Micritic coatings and micrite crusts (representing microcrystalline cement, possibly microbially triggered) are abundant. Coatings alternate with crusts consisting of calcite cement. • Micrite. Non-ferroan calcite, occurring (a) as internal micrite precipitate in burrows and shells, (b) microcrystalline cement between carbonate or siliciclastic grains, (c) pelmicrite. Often containing clay- and sand-sized grains, pyrite and organic matter. Sometimes recrystallized to microspar. • Micritic peloids are common. Some millimeter- to centimeter-sized peloids are fecal pellets (Pl. 148/2). • Clasts and breccia. Micrite and pelmicrite lithoclasts in autobreccia may be caused by burrowing or fluid circulation in the sediment (Fig. 16.17). • Lamination. Laminated fabrics composed of irregularly arranged layers of micrite and/or calcite cement are common. • Burrows and bioturbation. Burrows filled and lined with fine-grained carbonate containing grains similar in lithology and texture with those of the surrounding strata are common. Pyrite coatings. • Cement. Volumetric contribution of cement high to medium. Carbonate, sulfate and sulfide cements. Common carbonate cements are botryoidal calcite (rolling extinction between crossed nicols) forming botryoids up to 2 cm; fibrous calcite; irregular non-isopachous cement layers. • Organic matter. Dispersed debris, organic inclusions in calcite or black bands. • Pyrite. Common, occurring as linings, coatings, isolated and framboidal crystals. Geochemistry • Authigenic carbonate cements and carbonate shells extremely depleted in δ13C, indicating that the carbonates were derived from bacterial oxidation of methane.
Case Studies: Seep and Vent Carbonates
• Evidence for the presence of sulfide and methane: Fibrous cements derived from mixing of methane-rich fluids with seawater, micritic microbial fabrics, pyritecoated corrosion surfaces, pyrite framboids, pyrite crystals, irregular yellow calcite cements. • Penecontemporaneous sulfide and sulfate mineralizations may be common. • Biomarkers indicating anaerobic oxidation of methane by bacteria (Peckmann et al. 2001).
16.5.2 Case Studies Isolated occurrence of limestone bodies within noncarbonate sequences or carbonate-poor lithologies as well as carbonate mound-like structures characterized by low-diversity but abundant specific benthic faunas (predominantly bivalves, tube-worms as well as brachiopods) and conspicuous amounts of δ13C-depleted carbonate cements may represent authigenic ‘cold-seepcarbonates’ and not sedimentary structures. Some mollusks and worms are able to directly exploit gas seepage and fluid expulsions by hosting chemolithoautotrophic bacteria. Diagnostic criteria of ancient seep and vent carbonates are irregular textures, low-diversity but abundant mollusks, abundance of carbonate cements as well as of pyrite and organic matter. Deposits related to the venting of cold fluids (mainly hydrocarbons, carbon dioxide and water) are known from different geological settings (Box 16.5). Methane-derived and hydrocarbon-derived chemosynthetic carbonates produced at fossil cold-seeps and vents are commonly developed as carbonate masses, mounds and lenses within shales and siltstones. These ‘chemoherms’ form conspicuous structures in outer marine, faultbounded forearc, rift and accretionary prism settings, facilitating conduits for seeping hydrocarbons. Many ancient cold-seep assemblages occurred in water as shallow as the mid-shelf and outer shelf. In contrast, most modern cold-seep communities characterized by large biomass aggregations of animals with chemoautotrophic bacterial symbionts are restricted to bathyal and abyssal depths (Callender and Powell 1999). Seep mounds are associated with specific structures (fractures, diapirs, vents, methanogenic zones) rather than with specific depositional settings. The growth of structural highs may have been an effective mechanism for concentrating gas hydrates by gradually upwardly removing the base of their stability zone. Fluid release may result in sediment instability and large-scaled sliding structures (Conti and Fontana 2002). The internal lithology of seep mounds has little in common with
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Seep Carbonates
Box 16.5. Selected studies of ancient vent and seep communities and carbonates. See Callender and Powell (1999) for an extended list. Tertiary: Aharon and Sengupty 1994; Campbell 1992; Cavagna et al. 1999; Clari and Martire 2000; Clari et al. 1988, 1994; Conti and Fontana 2001, 2002; Goedert and Squires 1990; Goedert et al. 2003; Lutz and Kennish 1993; Niitsuma et al. 1989; Peckmann et al. 2001, 2003; Rigby and Goedert 1996; Squires and Goedert 1991; Terzi 1993; Terzi et al. 1994 Cretaceous: Beauchamp and Savard 1992; Beauchamp et al. 1989a, 1989b; Campbell and Bottjer 1995; Campbell et al. 1993; Kauffman et al. 1996; Lemoine et al. 1982 Jurassic: Campbell and Bottjer 1993; Campbell et al. 1993; Gaillard et al. 1992; Kelly et al. 1995; Lemoine et al. 1982; Little et al. 1999; Rolin et al. 1990; Carboniferous: Bittner et al. 1990, 1992; Peckmann et al. 2001 Devonian: Belka 1998; Kuznetsov et al. 1990; Mounji et al. 1998; Peckmann et al. 1999 Silurian: Liljedahl 1991; Little et al. 1997 Ordovician: Hovland 1989; Lavoie 1997; SteelePetrovich 1988.
the surrounding mound facies. Seep mound faunas exhibit low species richness and high dominance; they are autochthonous with little indication of disturbance. Shell dissolution can be extreme owing to the high CO2 or H2S content in interstitial waters.
16.5.2.1 Late Eocene ‘Whiskey Creek’ Seep Carbonates of Washington State, U.S.A. Carbonate deposits and invertebrate communities associated with areas where seepage fluids enriched in reduced compounds reached the seafloor have a long geologic record in marine strata along the Pacific slope of North America (Campbell and Bottjer 1993). Ancient methane-seep carbonates are widespread within the marine strata of the Olympic Peninsula, western side of Washington State. An example is the Whiskey Creek limestone formed around a methane seep. The limestone consists of micrite, carbonate cement and densely-packed large bivalves (Goedert et al. 2003; Peckmann et al. 2003). These carbonates occur as large limestone boulders exhibiting a chaotic fabric (Fig. 16.17) produced by bioturbation, concretionary carbonate formation, early in-situ brecciation and late fracturing. A nodular fabric is caused by pyritiferous micrite aggregates induced by anaerobic oxidation of methane and carbonate corrosion. Authigenic non-carbonate minerals are common constituents of the limestone. The biota comprises bivalves, worm tubes, benthic foraminifera and some palynomorphs. Abundant bivalve mollusks are represented by six species comprised solemyids, nuculanids, mytilids, lucinids, thyasirinids, and vesicomyids. Lipid biomarkers extracted from the
Fig. 16.17. Seep carbonates. Typical chaotic fabric of a methane-derived limestone. White areas are authigenic carbonate cements. Polished slab. Late Eocene: Washington State. After Peckmann et al. (2003).
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deposits indicate the existence of seep-related microbes and reveal that anaerobic oxidation of biogenic methane was an important component in the biogeochemical cycling of carbon in this ancient seep environment. 16.5.2.2 Early Cretaceous Seep Carbonates in the Canadian Arctic An excellent example of ancient of chemosynthetic cold-seep biota and carbonate mounds was reported from the Early Cretaceous (Aptian to Albian) of the Sverdrup Basin (Beauchamp and Savard 1992; Pl. 148). The circular mounds vary in diameter between 1 and 60 m and in height from 1 to 8 m. They originated in deep (about 400 m) and cold waters on the sea floor of a rifting basin. A cold climatic setting is indicated by glendolite concretions, pebble- and cobble-sized icerafted dropstones and the high paleolatitudinal setting of the siliciclastic sediments yielding the carbonates. The seep communities consist of abundant and densely growing bivalves and serpulid worm tubes, associated with terebratulid brachiopods. The fossils are enclosed in an isotopically-light carbonate groundmass which is volumetrically far more important than the fossils. The mounds formed by a complex sequence characterized by recurrence of fossils and methane-derived carbonate phases (micrite, botryoidal calcite/former aragonite, splayed calcite, yellow calcite), interrupted by episodes of carbonate corrosion and pyrite precipitation. These processes took place in an environment that shifted back and forth between aerobic, oxidizing, H2Sfree conditions and anaerobic, reducing and H2S-rich conditions.
Cold-Seep Carbonates
Basics: Ancient seep and vent carbonates Beauchamp, B., Bitter, P.H. (eds., 1992): Chemosynthesis: geological processes and products. – Palaios, 7, 337-484 Beauchamp, B., Savard, M. (1992): Cretaceous chemosynthetic carbonate mounds in the Canadian Arctic. – Palaios, 7, 434-450 Bitter, P.H.von, Scott, S.D., Schenk, P.E. (1992): Chemosynthesis: an alternative hypothesis for Carboniferous biotas in bryozoan/microbial mounds, Newfoundland, Canada. – Palaios, 7, 466-484 Callender, W.R. (1999): Why did ancient chemosynthetic seep and vent assemblages occur in shallower water than they do today? – International Journal of Earth Sciences, 88, 377-391 Campbell, K.A., Bottjer, D.J. (1995): Brachiopods and chemosymbiotic bivalves in Phanerozoic hydrothermal vent and cold seep environments. – Geology, 23, 321-324 Cavagna, S., Clari, P., Martine, L. (1999): The role of bacteria in the formation of cold seep carbonates: geological evidence from Monferrato (Tertiary, NW Italy). – Sedimentary Geology, 126, 253-272 Clari, P.A., Martire, L. (2000): Cold seep carbonates in the Tertiary of northwest Italy: evidence of bacterial degradation of methane. – In: Riding, R.E., Awramik, S.M. (eds.): Microbial sediments. – 261-269, Berlin (Springer) Gaillard, C., Rio, M., Rolin, Y., Roux, M. (1992): Fossil chemosynthetic communities related to vents or seeps in sedimentary basins: the pseudobioherms of southeastern France compared to other world examples. – Palaios, 7, 451-465 Goedert, J.L., Thiel, V., Schmale, O., Michaelis W., Peckmann, J. (2003): The Late Eocene ‘Whiskey Creek’ methane-seep deposit (Western Washington State). Part I. Geology, Paleontology and Molecular Geobiology. – Facies, 48, 223-240 Hovland, M., Thomsen E. (1997): Cold-water corals - are they hydrocarbon seep related? – Marine Geology, 137, 159-164 Little, C.T.S., Herrington, R.J., Haymon, R.M., Danelian, T. (1999): Early Jurassic hydrothermal vent community from
Plate 148 Early Cretaceous Cold-Seep Carbonates of the Canadian Arctic This plate displays microfacies criteria whose combination is diagnostic of seep carbonates. Diagnostic microfacies criteria are the occurrence of specific bivalves (–> 2) and worms (–> 1, 2), and the complex cementation pattern (–>1, 3). 1 Worm-tube dominated rudstone. Large and small serpulid worm tubes (black arrows) are cemented by several generations of methane-derived authigenic carbonates. Marked differences in the composition of the carbonate (micrite, various types of cements), corrosion features and pyrite precipitation indicate changes in carbonate precipitation and settlement by specialized chemoautotrophic organisms feeding on bacteria. Note the complex sequences of botryoidal and fibrous cement. Early Cretaceous (Aptian to Albian): Ellef Ringness Island, Canadian Arctic. 2 Bivalve shells in methane-derived micrite. The bivalves are lucinids. Numerous fecal pellets demonstrate high biological activity at seep sites. The ovoid pellets consist of dark micritic material with variable proportions of silt-sized quartz, foraminifers and bioclastic debris. Same locality as –> 1. 3 Large serpulids (S) cemented by multiple generations of methane-derived carbonates. Note the distinct difference between authigenic microcrystalline calcite (M) and other cement types (C). Early Cretaceous (Aptian to Albian): Prince Patrick Island, Canadian Arctic. –> 1–3: Courtesy of B. Beauchamp (Calgary)
Plate 148: Cretaceous Cold-Seep Carbonates
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the Franciscan Complex, San Rafael Mountains, California. – Geology, 27, 167-170 Little, C.T.C., Herrington, R.J., Maslennikov, V.V., Zaykov, V.V. (1998): The fossil record of hydrothermal vent communities. – In: Mills, R.A., Harrison, K. (eds.): Modern ocean floor processes and the geological record. – Geological Society London, Special Publication, 148, 259270 Peckmann, J., Gischler, E., Oschmann, W., Reitner, J. (2001): An Early Carboniferous seep community and hydrocarbon-derived carbonates from the Harz Mountains, Germany. – Geology, 29, 271-277 Peckmann, J., Goedert, J.L., Heinrichs, T., Hoefs, J., Reitner, J. (2003): The Late Eocene ‘Whiskey Creek’ methaneseep deposit. Part II. Petrology, stable isotopes and biogeochemistry. – Facies, 48, 241-254 Peckmann, J., Walliser, O.H., Riegel, W., Reitner, J. (1999): Signatures of hydrocarbon venting in a Middle Devonian mud mound (Hollard Mound) at Hamar Laghdad (Antiatlas, Morocco). – Facies, 40, 281-296 Powell, E.N., Callender, W.R., Stanton, R.J. (1998): Can shallow- and deep-water chemoautotrophic and heterotrophic communities be discriminated in the fossil record? – Palaeogeography, Palaeoclimatology, Palaeoecology, 144, 85-114 Rolin, Y., Gaillard, C., Roux, M. (1990): Ecologie des pseudobiohermes des Terres Noires jurassiques liés à des paléosources sous-marines. Le site oxfordien de Beauvoisin (Drôme, Bassin du Sud-Est, France). – Palaeogeography, Palaeoclimatology, Palaeoecology, 80, 79-105 Savard, M.M., Beauchamp, B., Veizer, J. (1996): Significance of aragonite cements around Cretaceous marine methane seeps. – Journal of Sedimentary Research, A66, 430-438, Terzi, C., Aharon, P., Lucchi, F.R., Vai, G.B. (1994): Petrography and stable isotope aspects of cold vent activity imprinted on Miocene age ‘calcari a Lucina’ from Tuscan and Romagna Apennines, Italy. – Geo-Marine Letters, 14, 177-184 Further reading: K211
16.6 Mixed Carbonate-Siliciclastic Settings and Limestone/Marl Sequences The following text deals with carbonates formed in mixed carbonate-siliciclastic environment and discusses limestone-marl sequences.
16.6.1 Carbonate-Siliciclastic Environments Mixtures of siliciclastic and carbonate rocks in shelf settings are a result of lateral facies mixing, or by sealevel changes and/or variations in sediment supply causing vertical variation in the stratigraphic succession. Having long been neglected, the study of modern and ancient mixed carbonate-siliciclastic environments has
Carbonate-Siliciclastic Environments
significantly increased since the 1980s. Many studies deal with the impact of siliciclastic sedimentation on reef growth. Other studies focus on the influence of siliciclastic influx on the development of carbonate platforms and shelves. Overviews: Walker et al. (1983) proposed the first model for carbonate to terrigenous clastic sequences, another model was suggested by Mount (1984). Doyle and Roberts (1988) edited a book describing the aspects of carbonate-clastic transitions in modern environments. Lomando and Harris (1991) presented case studies of ancient mixed carbonate-siliciclastic sequences on the scale of core data. Mount (1985) suggested a classification of sediments that are composed of mixtures of carbonate and siliciclastic material (see Sect. 8.5; Pl. 49). Important papers on carbonate-siliciclastic mixtures were collated in the SEPM Reprint Series by Budd and Harris (1990).
16.6.1.1 Modern Carbonate-Siliciclastic Environments Siliciclastic and carbonate sediments are often mutually exclusive within modern environments because of the biogenic origin of most carbonates. Fine-grained siliciclastics frequently dilute carbonate sediments and negatively effect carbonate production by reducing the light available to autotrophs and/or covering filter-feeding organisms. Transitions from siliciclastic to carbonate sedimentation occur in near-coast and inner shelf environments (Fig. 16.18) where rivers transport eroded terrigenous material to the sea, and on low to high-latitude shelves. The input of siliciclastics affects the carbonate production of platforms and ramps, and is an important control on reef growth. The transition between the two depositional systems is largely controlled by climate, sea level and tectonic setting, influencing in turn hinterland erosion and carbonate production. In equatorial settings, siliciclastic or volcaniclastic influx has a major influence on regional carbonate development, local carbonate deposition and sequence development (see Wilson and Lokier 2002 for an excellent study of causal relationships). Coeval carbonate and siliciclastic sedimentation requires specific conditions. Terrigenous influx has to be low enough to allow the growth of carbonate-producing organisms. This can be achieved by a temporary shift of the siliciclastic depocenter (e.g. abandonment of delta lobes), the winnowing of fine-grained material by currents, or spatial separation of siliciclastic and carbonate depocenters.
Carbonate-Siliciclastic Environments
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shore basins where the carbonates intercalate with basinal shales; • in basins with pelagic carbonate sedimentation.
Fig. 16.18. Mixed carbonate-siliciclastic sediment. The thin section shows a rock consisting of calcareous ooids and terrigenous quartz grains. The latter form the nuclei of many ooids or occur dispersed in the micritic matrix. The angular quartz grains are moderately sorted. The sample (a sandy allochem micrite; see Sect. 8.5) comes from massive to coarse bedded sediments interfingering with coral reef limestones and also with near-coastal sandstones and calcareous sandstones, and was deposited in an inner shelf environment. Late Jurassic (Early Kimmeridgian): Trevijano, Sierra de la Demanada, Province Soria, Spain. After Benke et al. (1981).
Modern low-latitude regions with mixed carbonatesiliciclastic sedimentation are known from the northern Puerto Rico shelf (Pilkey et al. 1988), Southwest Florida (Holmes 1988), the northern Great Barrier Reef (Flood and Orme 1988; Belperio and Searle 1988; Lacombe and Woolfe 1999), the Gulf of Suez and the Gulf of Aqaba, Red Sea (Friedman 1988; Roberts and Murray 1988), and along the coast of Tanzania (Fay et al. 1992). Examples of high-latitude mixed siliciclastic-skeletal carbonate sediments have been described from southern Australia (James et al. 1992) and New Zealand (Gillespie and Nelson 1997). Mixed siliciclastic and carbonate deposits occur • on shelves on which land-derived siliciclastic sediments are trapped in nearshore and inner shelf environments and contemporaneous deposition of carbonate sedimentation takes place in subtidal settings (e.g. eastern Nicaragua coast and Great Barrier Reef); • on siliciclastic shelves on which detrital carbonates are deposited, supplied by the erosion of uplifted carbonate rocks in the hinterland (source mixing model; Mount 1984); • at the margins of carbonate shelves where the shelves interfinger with deep-water shales. In these settings, storms periodically carry carbonates out into distal off-
Lateral transitions from siliciclastic to carbonate facies have been described occurring over different distances from the shoreline, depending on the climate and precipitation rates (Roberts 1987; Fay et al. 1992). The strong influx of siliciclastic sediment into the lowlatitude near-shore carbonate Nicaragua shelf can be followed for up to 20 km. In contrast, under the conditions of an arid climate (Red Sea) and low precipitation siliciclastic influx is restricted to periodic flash floods affecting the near-shore carbonate environments only within distances of tens of meters or a few kilometers; transitions between siliciclastic and carbonate depositional environments are abrupt. Relatively gradual transition patterns confined to a distance of a few kilometers from the shoreline have been observed in the moderately humid climate of the low-relief coast of Tanzania. The siliciclastic and volcaniclastic influx into environments with carbonate production differs with regard to the intensity, grain size and effects on salinity and nutrients (Wilson and Lokier 2002): • Tropical/equatorial, humid areas are characterized by high-continuity clastic influx, clay to cobble clast size (often clay to sand), normal to low salinity, and high nutrient contents. In siliciclastic-dominated environments, regional carbonate production can be promoted by structural or depositional antecedent topography, shelf currents, changes in clastic distribution pathways and relative sea-level change (Wilson 2002). • Subtropical arid areas exhibit moderate and episodic clastic influx, silt to boulder clast size, normal to high salinities, and low nutrient contents. • Volcanic shoreline-attached areas are affected by high to very highly continuous and punctuated clastic influx, clay to boulder clast size, normal to low salinities, and high nutrients. • Volcanic influx on distal parts of isolated platforms is low to high-punctuated, clast size corresponds to clay to cobble, salinity is normal, and nutrients range between low and high. Effects of siliciclastic input on shallow-marine carbonate production. Siliciclastic or volcaniclastic input into an area of carbonate production has a number of potential effects on carbonate producers: • Physical covering or surrounding of the organisms, often resulting in the death and burial of the biota.
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• Increased suspended matter in the water column will result in a reduction of light levels. Turbidity will restrict the type and composition of communities present and reduce the water depth at which different assemblages occur. • The increase of siliciclastic or volcaniclastic sediments can cause changes in water chemistry, such as reduced salinities, variations in pH, and an increase in nutrients. Salinity and pH changes may have a detrimental effect on organic carbonate production. Reefs and siliciclastic sediments. Many reef organisms are adapted to nutrient-poor conditions. The high input of nutrients leading to eutrophication may cause altered skeletal growth rates and changes in biotic composition, and finally the decline and death of reefs. Terrigenous sedimentation, however, must not completely hinder the development of reefs, but may control the position and geometry of reefs and modify ecologic zonations (Roberts 1987; Santisteban and Taberner 1988; Acevedo et al. 1989; Carbone et al. 1999). Modern reefs can survive in areas of significant terrigenous influx and turbidity. For example, deeper-water coral reefs are able to grow despite periodic blanketing by fine siliciclastic sediment introduced during times of heavy rains (Acker and Stearn 1990). Corals reefs are not uncommon in some coarse-grained deltas (see McPherson et al. 1987). The effect of terrigenous input on coral reefs is described by a model based on the relative net rates of terrigenous framework and carbonate non-framework sediment accumulation and and/or removal (Woolfe and Lacombe 1999).
16.6.1.2 Ancient Mixed Carbonate-Siliciclastic Environments The formation of carbonate platforms and reefs within siliciclastic settings requires combinations of adaptive strategies by carbonate-producing organisms and sheltering mechanisms that protect the organisms from unfavorable influences. Sheltering mechanisms include an elevated position (e.g. basement uplift) within areas dominated by siliciclastic sedimentation, longshore currents that screen off suspended siliciclastic grains, local subsidence traps, or sea-level rise resulting in the reduction of siliciclastic influx and favoring carbonate deposition (Blair 1988; Brachert 1992; Garcia et al. 1993; Philip 1993; Leinfelder 1994; Okhravi and Amini 1998; Sanders and Höfling 2000; Khetani and Read 2002).
Describing Carbonate-Siliciclastic Sediments
Ancient reefs and siliciclastic/pyroclastic sedimentation: Reefs formed in environment influenced by siliciclastic influx are known all over the Phanerozoic, but became more frequent from the Late Triassic. Coral reefs occur in different siliciclastic settings in the Jurassic (Leinfelder et al. 2002). Because rudists could cope with sediment-loaden and turbid waters (Steuber et al. 1998; Mitchell 2002), Cretaceous rudist reefs also occur in stressed siliciclastic environments. The number of coral reefs formed in mixed carbonate-siliciclastic environments increased during the Tertiary. Pyroclastic sedimentation can partially or completely surpress growth rates of reefs, depending on the amount and grain size of pyroclastic sediment in the water column. Turbidity caused by the suspension of volcanic ash reduces light penetration and strongly discourages the growth of calcareous algae and corals. Fine ash will plug up filter-feeding reef organisms. Island-arc carbonates and volcaniclastic sediments. Associations of limestones, volcaniclastic strata and volcanics have been described from several island-arc carbonates and tectonic terranes. These limestones occur as (1) carbonate platforms around volcanic edifices (Soja 1990), (2) gravity-displaced blocks embedded within basinal sediments (Blome and Nestell 1991), and (3) small autochthonous limestone lenses enclosed by deltaic volcaniclastic rocks (Watkins 1993; Beltramo et al. 2001).
16.6.1.3 Describing Carbonate-Siliciclastic Sediments: Practical Advice Understanding carbonate-siliciclastic successions requires field work focused on geometrical relationships and lateral and vertical lithologic changes, paleontological studies using fossils as sensitive indicators of environmental changes, and evaluation of the composition of the sediments. Description. Microfacies data are valuable in compositional analysis. In thin sections look for: • Matrix and carbonate grains (categories, frequency, distribution). • Fossils. Many organisms are adapted to siliciclastic influx. Adaption is reflected by the specific morphologies and composition of the assemblages as illustrated by benthic foraminifera (Kumar and Saraswati 1997), corals (Carbone et al. 1999), and bivalves (Best and Kidwell 2000). • Siliciclastic grains (mineralogy, size, angularity/ roundness, texture, abundance).
Mixed Siliciclastic-Carbonate Rocks
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Fig. 16.19. Mixed siliciclastic-carbonate rocks. Terminology of arenites (grain-supported siliciclastic and/or carbonate rocks consisting of more than 50% sand and/or rock fragments). A: Conglomerate composed of arenitic matrix and rounded lithoclasts derived from the erosion of Triassic and Jurassic carbonate rocks. The lithoclasts are partly pitted, indicating surficial dissolution. The arenite of the matrix consists of very angular to very well rounded carbonate rock fragments and a few bioclasts. Shallow-marine environment. Late Cretaceous: Voldöppberg, Brandenberg, Tyrol, Austria. B: Poorly sorted calcilithic hybrid arenite composed of carbonate rock fragments (ooid grainstone, bottom; dolostone; mudstone to wackestone, center) derived from Triassic to Late Cretaceous rocks, as evidenced by microfossils. The rock fragments are subrounded to well-rounded; some are blackened or stained red. Most bioclasts are penecontemporaneous. Isopachous fringes of dogtooth cement are overlain by geopetal infill of lime mudstone, which in turn is covered by blocky cement (center of the picture). Shore zone environment. Same locality as –> A. C: Moderately sorted bioclastic hybrid arenite consisting of carbonate rock fragments, siliciclastic grains and bioclasts. The coarse bioclastic fraction is dominated by nerineid and actaeonellid gastropods (right). Smaller subrounded to wellrounded, less common bioclasts are echinoderms and benthic foraminifera. The siliciclastic fraction consists of subangular to subrounded polycrystalline and monocrystalline quartz grains and serpentine grains, indicating the erosion of crystalline rocks. Lagoonal environment. Late Cretaceous: Nachbergalm, Brandenberg, Tyrol, Austria. D: Transition from a bioclastic grainstone (bottom) into an arenite composed of carbonate rock fragments, siliciclastic grains and bioclasts. The grainstone consists of rounded, coated fragments of mollusks, including rudist fragments, algae, corals, echinoderms, benthic foraminifera and some siliciclastic sand. The grains are cemented by isopachous fringes of dogtooth spar, overlain by blocky calcite spar. The boundary between the grainstone and the overlying arenite is accentuated by pressure solution. Lithification of the arenite proceeded mainly by intergranular pressure solution, producing a tightly packed texture. Shelf arenite. Late Cretaceous: Haidach, Brandenberg, Tyrol, Austria. Scale is 2 mm for all pictures. After Sanders (1998).
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• Relationships between siliciclastic grains and carbonate grains. Are the siliciclastic grains incorporated within carbonate grains (e.g. in oncoids or as nuclei of ooids, Fig. 16.18)? Classification. The classification of mixed carbonatesiliciclastic rocks suggested by Mount (1985; Sect. 8.5) or the differentiation of arenite types (Fig. 16.19) are recommended. A common rock type formed in siliciclastic-carbonate environments is characterized by a framework supported by sand-sized grains. The term arenite (Zuffa 1985) is recommended for a sample containing more than 50% of siliciclastic sand and/or carbonate rock fragments. Arenites consisting of subequal amounts of carbonate rock fragments, bioclastic material and siliciclastic sand are termed bioclastic hybrid arenites. Arenites composed of carbonate rock fragments and siliciclastic sand are designated as calcilithic hybrid arenites. Arenites consisting mainly of carbonate rock fragments are called calcilithic arenites. The terminological classification assists in the genetic interpretation of mixed shelf sediments deposited in front of a mainland with a relatively high, articulated relief. Examples are the Late Cretaceous (Turonian to Santonian) sediments of the Lower Gosau Subgroup in the Tyrolian Northern Calcareous Alps in Austria (Sanders 1998).
16.6.2 Limestone-Marl Sequences: Primary or/and Diagenetic Origin? Classical definitions differentiate limestone from marl using the ranges of 95–100% CaCO3 and 35–65% CaCO3 (Correns 1949), but the definitions of marl vary strongly (Bausch 1997). In rhythmic successions the terms limestone and marl are used in a descriptive sense. Layers that are more resistant to weathering are called limestones; the intercalated less-resistant layers are commonly termed marls. The limestones of limestonemarl alternations are not or only slightly compacted. The highest carbonate content occurs in the central part of the bed or nodule. Marls are strongly compacted and have lower carbonate contents than the adjacent limestone beds. Phanerozoic limestone-marl alternations varying from nodular (mostly marl-dominated) to well-bedded (limestone-dominated) sequences are widespread in ancient shelf and deep-sea environments. These sequences occur from shallow to deeper marine to hemipelagic settings. The striking appearance in the out-
Limestone-Marl Sequences
crop is caused by different weathering as a result of differential diagenesis. Tightly cemented limestone beds alternate with less resistant, less cemented marl beds. The non-compacted nature of the limestones implies an early diagenetic cementation and an import of carbonate from an external source occluding the primary porosity (Bathurst 1987). Controversial discussions continue on whether the limestone-marl alternations are caused by primary variations in the composition of the sediment – often explained as climatically forced differences; e.g. in pelagic sequences see Mount and Ward 1986; Bellanca et al. 1996), or/and by a diagenetic enhancement of primary rhythmic sedimentation patterns (Ricken 1994). Important points are: • For some rhythmic successions a primary, sedimentary nature of limestone-marl alternations is proven by differences between the fossil content of limestones and marls. • Diagenesis strongly enhances differences in the initial sediment composition (Ricken 1986). Initial differences may involve fluctuations in carbonate supply (productivity cycles; Seibold 1952) or periodic increases in clay supply (dilution cycles; Einsele 1982). • Munnecke (1997) and Munnecke et al. (2001) proposed a model that explains the development of limestone-marl alternations by selective dissolution of aragonite in marl beds and reprecipitation of calcite cement in limestone beds. Estimating the primary mineralogy of skeletal grains both in limestones and marls is essential in evaluating this model. • Because the primary sedimentary composition (fossil content, carbonate mineralogy) may be changed by selective dissolution (e.g. of aragonite), redistribution of calcium carbonate and reprecipitation, differences in carbonate content, sedimentary components or trace elements do not necessarily reflect primary differences in the sedimentary succession. • For successions lacking indications of systematic initial differences, a diagenetic origin of rhythmic alternations from homogeneous precursor sediments has been assumed (diagenetic overprint by selective dissolution of aragonite; Munnecke and Samtleben 1996; Munnecke 1997). • Computer modeling shows that diagenetic self-organization alone is not sufficient to produce limestonemarl alternations that can be followed laterally over hundreds of meters or even kilometers (Böhm et al. 2003). The simulation model indicates that the interaction of diagenetic self-organization, an external trigger (e.g. paleoclimate) and minor differences in pri-
Limestone-Marl Alternations
mary sediment composition, porosity or organic content must be considered for unequivocal interpretations. Differential diagenesis in limestone-marl alternations includes redistribution of calcium carbonate from marl layers to limestone beds by dissolution, migration of ions, and reprecipitation. Modern fine-grained shelf carbonates at low latitudes, which represent recent facies equivalents of many ancient limestone-marl successions, are generally composed of a mixture of aragonitic, calcitic and terrigenous material. The precursor mineralogy of limestone-marl alternations can be reconstructed using a model developed by Munnecke and Samtleben (1996), Munnecke (1997), Munnecke et al. (1997), and Westphal et al. (2000). The model assumes that the source of Low-Mg calcite carbonate cement precipitation in the limestone layers is derived from the dissolution of aragonite in marl layers. Since the amount of aragonite present in the precursor sediment limits the amount of calcium carbonate available for redistribution and precipitation, differential diagenesis comes to a halt when the aragonite has been completely dissolved. Mass balances and initial contents of aragonite, calcite and insolubles can be calculated. Mathematical relationships between the thickness of limestone and marl layers and their carbonate content can be visualized in an aragonite-calcite-terrigenous material ternary diagram (Munnecke et al. 2001). The model is supported by the concentration of limestonemarl alternations in warm-water settings (Westphal and Munnecke 2003). This coincides with the high aragonite production in these settings at the present time. Basics: Mixed carbonate-siliciclastic environments, limestone-marl alternations Mixed carbonate-siliciclastic environments Acker, K.L., Stearn, C.W. (1990): Carbonate-siliciclastic facies transition and reef growth on the northeast coast of Barbados. – Journal of Sedimentary Petrology, 60, 18-25 Braga, J.C., Martin, J.M., Alcala, B. (1990): Coral reefs in coarse-terrigenous sedimentary environments (Upper Tortonian, Granada Basin, southern Spain). – Sedimentary Geology, 66, 135-150 Budd, D.A., Harris, P.M. (eds., 1990): Carbonate-siliciclastic mixtures. – Society of Economic Paleontologists and Mineralogists, Reprint Series, 14, 400 pp. Doyle, L.J., Roberts, H.H. (eds., 1988): Carbonate-clastic transitions. – Developments in Sedimentology, 42, 304 pp., Amsterdam (Elsevier) Friebe, J.G. (1991): Carbonate sedimentation within a siliciclastic environment: the Leithakalk of the Weisseneck Formation (Middle Miocene, Styria, Austria). – Zentralblatt für Geologie und Paläontologie, I, 1990, 1671-1687 Leinfelder, R.R. (1993): A sequence stratigraphic approach to the Upper Jurassic mixed carbonate siliciclastic succession of the central Lusitanian Basin, Portugal. – Profil, 5, 119-140
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Leinfelder, R.R. (1994): Karbonatplattformen und Korallenriffe innerhalb siliziklastischer Sedimentationsbereiche (Oberjura, Lusitanisches Becken, Portugal). – Profil, 6, 1-207 Lomando, A.J., Harris, P.M. (eds., 1991): Mixed carbonatesiliciclastic sequences. – Society of Economic Paleontologists and Mineralogists, Core Workshop Notes, 5, 580 pp. McPherson, J.G., Shanmuga, G., Moiola, R. (1987): Fandeltas and braid-deltas: varieties of coarse-grained deltas. – Geological Society of America, Bulletin, 99, 331-340 Mount, J.F. (1984): Mixing of siliciclastic and carbonate sedimentation in shallow shelf environments. – Geology, 12, 432-435 Mount, J.F. (1985): Mixed siliciclastic and carbonate sediments: a proposed first-order textural and compositional classification. – Sedimentology, 32, 435-442 Roberts, H.H. (1987): Modern carbonate-siliciclastic transitions: humid and arid tropical examples. – Sedimentary Geology, 50, 25-65 Steuber, T., Yilmaz, C., Löser, H. (1998): Growth rates of early Campanian rudists in a siliciclastic-calcareous setting (Pontid Mts., north-central Turkey). – Geobios, Mémoir spécial, 22, 385-401 Wilson, M.E.J. (2002): Cenozoic carbonates in SE Asia: implications for equatorial carbonate development. – Sedimentary Geology, 147, 295-428 Wilson, E.J., Lokier, S.L.(2002): Siliciclastic and volcanic influences on equatorial carbonates: insights from the Neogene of Indonesia. – Sedimentology, 49, 583-601 Woolfe, K.A., Lacombe, P. (1999): Terrigenous sedimentation and coral reef growth: a conceptual framework. – Marine Geology, 155, 331-345 Zuffa, G.G. (1985): Optical analysis of arenites: influence of methodology on compositional composition. – In: Zuffa, G. G. (ed.): Provenance of arenites. – 165-189, Dordrecht (Reidel) Further reading: K089, K167, K182 Limestone/marl alternations Bathurst, R.G.C. (1987): Diagenetically enhanced bedding in argillaceous platform limestones: stratified cementation and selective compaction. – Sedimentology, 34, 749-778 Bausch, W.M. (1997): Die Flexibilität der Kalk/’Mergel’Grenze und ihre Berechenbarkeit. – Zeitschrift der Deutschen Geologischen Gesellschaft, 148, 247-258 Bellanca, A., Claps, M., Erba, E., Masetti, D., Neri, R., Premoli Silva, I., Venezia, F. (1996): Orbitally induced limestone/marlstone rhythms in the Albian-Cenomanian Cismon section (Venetian region, northern Italy): sedimentology, calcareous and siliceous plankton distribution, elemental and isotope geochemistry. – Palaeogeography, Palaeoclimatology, Palaeoecology, 126, 227-260 Böhm, F., Westphal, H., Bornholdt, S. (2003): Required but disguised: environmental signals in limestone-marl alternations. – Palaeogeography, Palaeoclimatology, Palaeoecology, 189, 161-178 Eder, W. (1982): Diagenetic redistribution of carbonate, a process in forming limestone-marl alternations (Devonian and Carboniferous, Rheinisches Schiefergebirge, W. Germany). – In: Einsele, G., Seilacher, A. (eds.): Cyclic and event stratification. – 98-112, Berlin (Springer) Einsele, G. (1982): Limestone-marl cycles (peridites): diagnosis, significance, causes: a review. – In: Einsele, G.,
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Seilacher, A. (eds.): Cyclic and event stratification.– 853, Berlin (Springer) Einsele, G., Ricken, W. (1991): Limestone-marl alternations: an overview. – In: Einsele, G., Ricken, G., Seilacher, A. (eds.): Cycles and events in stratigraphy. – 23-47, Berlin (Springer) Hallam, A. (1986): Origin of minor limestone-shale cycles: climatically induced or diagenetic? – Geology, 14, 609612 Mount, J.F., Ward, P. (1986): Origin of limestone/marl alternations in the Upper Maastrichtian of Zumaya, Spain. – Journal of Sedimentary Petrology, 56, 228-236 Munnecke, A. (1997): Bildung mikritischer Kalke im Silur von Gotland. – Courier Forschungsinstitut Senckenberg, 198, 1-71 Munnecke, A., Samtleben, C. (1996): The formation of micritic limestones and the development of limestone-marl alternations in the Silurian of Gotland. – Facies, 34, 159176 Munnecke, A., Westphal, H. (2003): Shallow-water aragonite recorded in bundles of limestone-marl alternations – the Upper Jurassic of SW Germany. – Sedimentary Geology, 164, 191-202 Munnecke, A., Westphal, H., Elrick, M., Reijmer, J.J.G. (2001): The mineralogical composition of precursor sediments of calcareous rhythmites: a new approach. – International Journal of Earth Sciences, 90, 795-812 Ricken, W. (1986): Diagenetic bedding. – Lecture Notes on Earth Science, 6, 210 pp., Berlin (Springer) Ricken, W. (1992): A volume and mass approach in carbonate diagenesis: the role of compaction and cementation. – In: Wolf, K.H., Chilingarian, G.V. (eds.): Diagenesis, vol. III. – 291-316, Amsterdam (Elsevier) Seibold, E. (1952): Chemische Untersuchungen zur Bankung im unteren Malm Schwabens. – Neues Jahrbuch für Geologie und Paläontologie, Abhandlungen, 95, 337-370 Thunell, R., Rio, D., Sprovieri, R., Raffi, I. (1991): Limestone-marl couplets: origin of the Early Pliocene Trubi marls in Calabria, Southern Italy. – Journal of Sedimentary Petrology, 61, 1109-1122 Westphal, H., Head, M.J., Munnecke, A. (2000): Differential diagenesis of rhythmic limestone alternations supported by palynological evidence. – Journal of Sedimentary Research, 70, 715-725 Westphal, H., Munnecke, A. (2003): Limestone-marl alternations: a warm-water phenomenon? – Geology, 31, 263266 Further reading: K182, K194
16.7 Secular Variations in Carbonate Depositional Patterns and Temporal Changes in Microfacies Criteria Carbonate depositional systems as well as microfacies criteria exhibited secular variations during the Phanerozoic. Changes and overall trend patterns are reflected by the composition and distribution of platform carbonates, reef limestones and pelagic sediments (Sect. 16.7.1). Marine benthic carbonate factories, including
Carbonate Environments during Time
the tropical shallow-water, the cool-water and the mudmound production systems differ in composition, geometry and facies patterns, and show differences in spatial distribution and occurrence in time (Sect. 16.7.2). The prevailing mineralogic composition of non-skeletal and skeletal grains, and of micrite and cement has changed over time (Sect. 16.7.3). Temporal differences in the mode, abundance and distribution of some microfacies criteria must be considered using these criteria to interpret paleoenvironmental conditions and depositional patterns (Sect. 16.7.4).
16.7.1 Changes in Major Carbonate Depositional Environments over Time Precambrian carbonates. The rates of carbonate sedimentation varied through time, starting as early as the Archean (Hay 1999; Grotzinger and James 2000). The temporal progression of major carbonate facies types during the Precambrian provides a record of long-term chemical evolution of sea water and of the biological evolution. Archean and Paleoproterozoic carbonates are characterized by abundant sea floor precipitates. Neoproterozoic carbonates were dominated by clastic-textured facies and carbonate mudstones. Mesoproterozoic and Early Neoproterozoic carbonates contain the enigmatic ‘molar-tooth’ structures. Many Neoproterozoic grainstones are unique because of the occurrence of giant, centimeter-sized calcareous ooids. Reefs appeared early in the Archean. Archean to Mesoproterozoic reefs are predominantly stromatolitebased. Lamination textures indicate a shift from in-situ precipitation of aragonite and calcite encrusting the sea floor to textures formed by accretion of loose sediment through trapping and binding in Neoproterozoic stromatolites. This trend reflects a decrease in abiotic factors and an increase in the control of stromatolites by benthic microbial mats. Terminal thrombolitic Proterozoic reefs contain the first calcified metazoans. Snowball Earth. The currently debated Snowball Earth hypothesis postulates a global glaciation during Early and Late Proterozoic time. The Neoproterozoic glacigene sediments are covered worldwide by a carbonate unit, typically dolostone. Some of the best preserved cap carbonates are developed on ‘Marinon’ glacigene sediments in the Mackenzie Mountains, Canada (James 2001). It consists of two units, a lower unit interpreted as a worldwide precipitation event, with laminated, peloidal and stromatolitic dolostones containing conspicuous tepees, and an upper transgressive unit with peloidal grainstones to cementstones, grading upward into thick shale. These carbonates are char-
Phanerozoic Platforms
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acterized by spectacular crystal fans of former aragonite. The crystal stromatolites form extensive bioherms and biostromes, with stacked reefs at the paleoshelf edge. The following discussion is limited to trends in the development of Phanerozoic carbonate sediments.
16.7.1.1 Phanerozoic Carbonate Platforms Carbonate platforms changed substantially in spatial extent, geometry, composition and paleogeographical distribution throughout the Phanerozoic (Kiessling et al. 2003). Trends in paleogeographical distribution and spatial extent are related to continental drift, evolutionary patterns of carbonate platform biotas, climatic change, and possibly, variations in ocean chemistry. A critical analysis of Phanerozoic platforms exhibits the following trends:
Fig. 16.21. Latitudinal distribution of relative carbonate platform area. The Cambrian is not shown because of problematic plate reconstructions. The latitudinal range of platform deposits varied markedly. Note the decline in carbonate production in the core tropics and the increase in the proportion of carbonate production in the higher tropics and lower subtropics in post-Paleozoic times. For the Carboniferous, Permian and Neogene, high latitude carbonate platform positions coincide with global cooling. These platforms exhibit profound qualitative differences in grain associations by comparison to tropical platforms (see Sect. 12.2). After Kiessling et al. (2003).
Fig. 16.20. Composition of Phanerozoic carbonate platforms in 32 supersequences. The names used for supersequences (right) correspond to the time of maximum transgression within the supersequence that has a wider range (e.g. the designation Tremadoc comprises Latest Cambrian to earliest Ordovician; see Golonka and Kiessling 2002). The plot refers to tropical and subtropical platforms. The most prominent groups in the ‘Others’ category are indicated by Br - brachiopods, By - bryozoans, E - echinoderms, M - mollusks, and R - rudist bivalves. Note the difference in the composition of Paleozoic and Mesozoic to Cenozoic platform carbonates. After Kiessling et al. (2003).
• Carbonate platform composition shifted from a microbial-oolitic dominance in the Cambrian and Early Ordovician to an echinoderm-bryozoan-brachiopod dominated composition in the later Early Paleozoic (Fig. 16.20). Late Paleozoic and younger platforms are usually strongly influenced by algal and benthic foraminiferal production. Echinoderms retained their rockbuilding potential until the Jurassic. Rudist bivalves contributed significantly to Late Cretaceous platforms, whereas other mollusks, with a few exceptions in the Late Paleozoic and Early Mesozoic, achieved a rockbuilding potential only in the Cenozoic. • The paleogeographical distribution of carbonate platforms shows two significant trends: The latitudinal range of platform sedimentation widened and the relative coverage of equatorial shelves by carbonate platform sedimentation declined through the Phanerozoic (Fig. 16.21). • Paleozoic carbonate platforms were much larger and less scattered on average than Mesozoic and Cenozoic
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Reef Patterns
Fig. 16.22. Carbonate platforms and reefs. Global area covered by carbonate platforms (left) and integrated reef production at a supersequence stratigraphic resolution (right, data from Kiessling et al. 2000). Error bars indicate statistical 95% confidence intervals based on total values, but do not consider other uncertainties. Reefal carbonate production varied more strongly than platform carbonate production. For most of the time platform and reefal production appears to be decoupled. After Kiessling et al. (2003).
platforms (see maps of platform distribution in Kiessling et al. 2003). • Changes in global carbonate platform areas do not vary systematically with climatic fluctuations or with trends in the availability of shallow shelves at low latitudes. • Carbonate platforms usually survive mass extinction events owing to their global size and commonly also to their accumulation rate, which is not significantly affected because of the potential of platforms to balance taxonomic loss by increased non-enzymatic carbonate production and enhanced production through surviving taxa. • Benthic foraminifera, which achieved their highest productivity in pre-Miocene off-reef settings, are important components of platform carbonates. • Platform production and reef development were decoupled over most of the Phanerozoic time (Kiessling et al. 2000) although parallel temporal trends exist, as shown by the increased development of platform-margin reefs and relatively high platform carbonate production (e.g. in the Devonian and Jurassic). The decoupling is especially evident when the integrated reef carbonate production is compared with the spatial extent of carbonate platforms during times of maximal transgression within Phanerozoic supersequences (Fig. 16.22). Fluctuations in reef production are much more pronounced than fluctuations in platform size and platform accumulation rates. No significant correlation exists between changes in reconstructed reefal carbonate production and changes in platform size. Variations
in calcareous non-reefal benthic organisms (e.g. foraminifera; Fig. 16.20) are apparently an important control on changes in carbonate platform production. • The evolution of carbonate-secreting biota, whose reaction to oceanographic changes is determined by their genetically fixed ecologic capabilities, is a major control on platform development and is as important as sea-level history.
16.7.1.2 Phanerozoic Reef Patterns Many Phanerozoic reef attributes exhibit significant temporal trends. The abundance of reefs, preferential depositional setting, dominant constructional and compositional reef types, prevailing reef biota (Sect. 16.2.3.1), and reefal carbonate production related to reef size and reef debris potential changed markedly during the Phanerozoic. The following discussion of some of the reef attributes is based on Kiessling (2002) and consider the PaleoReef database (Kiessling et al. 1999; Kiessling and Flügel 2002). Note that this database uses a broad reef definition that includes ecologic reefs, reef mounds and mud mounds (see Sect. 16.2). Abundance of reefs. Fig. 16.23 shows the number of reef sites in the PaleoReef database calculated for stages. The graph exhibits a strongly fluctuating pattern with a Phanerozoic maximum abundance of reefs in the Frasnian (Copper 2002 sees the maximum in the Givetian) and Miocene, and high reef abundance in the
Reef Crises
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Fig. 16.23. Abundance of Phanerozoic reefs expressed by the absolute number of reef sites (lumps of age-equivalent reefs from an area of more than 300 square kilometer; gray bars) and the number of reef sites per million years (line, scale at the right) in the PaleoReef database calculated for stages. Global reef abundance was greatest in the Frasnian (mostly stromatoporoid-coral reefs) and very high in the Givetian (stromatoporoid-coral reefs), Oxfordian (coral reefs and siliceous spongemicrobial reefs) and Middle Miocene (coral reefs and algal reefs). Courtesy of W. Kiessling (Berlin).
Wenlockian, Givetian, Viséan, Asselian, Norian, Oxfordian, Albian and Miocene. High reef abundance is independent of the dominant reefbuilders and association of reef-building organisms. Most Wenlockian, Givetian and Frasnian reefs are dominated by stromatoporoids and corals, but Frasnian mud mounds exhibit high microbial contributions, similar to Viséan reef mounds in which microbes and bryozoans are of importance. Most Early Permian Asselian reefs are algal reefs. Guadalupian reefs are rich in coralline calcareous sponges similar to Norian reefs. Oxfordian reefs include reefs formed by siliceous sponges and microbes as well as scleractinian coral reefs. Albian reefs are usually rudist bivalve and/or coral reefs. Seravallian (mid-Miocene) reefs are characterized by coralline algae and corals.
which is not reflected by a reef gap or a global reef crisis. • Major mass extinctions are usually associated with reef crises, but different rankings in the magnitudes of reef crises and extinctions exist. Coincidences exist for the Early-Middle Cambrian, end-Ordovician, Middle Permian-Late Permian and Permian-Triassic events. Reef crises were of a higher magnitude than coeval mass extinctions at the Mississippian-Pennsylvanian, Triassic-Jurassic, Jurassic-Cretaceous, and Cenomanian-Turonian boundaries. • Reef crises affected all reef-building organisms including sponges, corals and bivalves, and occasionally also microbes. A bloom of microbes occurs after reef crises and mass extinctions, but is not a universal rule, as is usually argued.
Global extinction events affected reef abundance in different ways. If global reef crises are defined as significant drops in global reefal carbonate production (Flügel and Kiessling 2002), and not by the measure of diversity, eight first-order reef crises with a carbonate production loss > 90% can be distinguished. The analysis of these reef crises exhibits the following patterns: • Mass extinctions do not necessarily cause reef crises as exemplified by the end-Ordovician, the CarnianNorian turnover, and the Cretaceous-Tertiary event,
Constructional reef types. There is an overall trend to a relative increase in true (ecologic) reefs and reef mounds and a relative decrease in mud mounds during the Phanerozoic. True reefs and reef mounds dominated during the Paleozoic. Mud mounds were common or predominant only in the Paleozoic, particularly in the Early Paleozoic. The most prominent increase in true reefs occurred in the Tertiary. Reef dimensions. Reef thickness is most likely the significant measure of reef dimensions. Fig. 16.24
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Pelagic Carbonates
Fig. 16.24. Reef thickness. Mean thickness values of reefs were calculated from the ratio of the sum of thickness values for all reefs of a stage divided by the number of reef sites. The interval values were translated into metric values to produce expressive means: Each point in the graph corresponds to a stage in the figure. Mean thickness was moderate in the Early and Middle Cambrian and in the Early to early Middle Ordovician reefs and rose sharply in the Ashgillian. The trend towards thicker reefs up in the Late Devonian is interrupted by several setbacks. Mean thickness increased in the Givetian and reached a maximum in the Frasnian, followed by a decline during the Carboniferous with a significant drop in the Early Pennsylvanian. A pronounced and continuous increase took place until the Artinskian, followed by a decrease during the Middle and Late Permian and a significant reduction in the Scythian microbial reefs. A prominent peak is shown by the coralline sponge-microbe reefs of the Ladinian. Subsequent Triassic stages are marked by descreasing mean thickness values. After a sharp decline in the Hettangian, moderately thick reefs developed during the Early Jurassic and Aalenian. Thickness rose sharply from the Callovian to the Kimmeridgian and declined from the Tithonian to the Hauterivian. In the Cretaceous moderate thickness values were achieved only from the Barremian to the Cenomanian, whereas later Cretaceous stages are characterized by low thickness. Tertiary reefs exhibit fluctuating mean thickness values and a distinct trend towards thicker reefs in the Neogene. Modified from Kiessling (2002).
shows that there is no overall trend in reef thickness during the Phanerozoic, but that stages with great mean reef thickness are commonly followed by stages with minor reef thickness. Maximum reef thickness values in the Frasnian and Artinskian are due to stacked reef mounds and reefs and may be biased by the high percentage of subsurface reefs included in the calculation. Subsurface reef studies tend to overestimate reef thickness because of the inclusion of non-reefal carbonates in the reef volume, or because the ‘reefs’ are more likely to seismic reefs. Reef debris. Many ancient reef complexes consist predominantly of debris, whereas the proportion of insitu organic framework can be very low. Because reefal debris may be porous and can contain significant hydrocarbon accumulation, knowledge of secular variations in the amount of reef debris is important. The calculation of debris production in Phanerozoic reefs, taking in consideration reef abundance, reef sizes and debris potential demonstrates extreme variations between times of high and low debris production (Kiessling 2002). The two largest peaks are evident in the Givetian to the Famennian time interval and the latest Middle Jurassic to Late Jurassic interval. A third
peak occurs in the Late Oligocene to Neogene interval. Smaller highs in reef debris production occur in the late Early to Late Silurian, the Middle Permian to Late Triassic, and the mid-Cretaceous.
16.7.1.3 Pelagic Carbonates Pelagic carbonate production today is greater in terms of mass than shallow-water carbonate production (Milliman and Droxler 1986), but the sedimentary record of pelagic carbonates is poor in pre-Jurassic rocks (Kuznetsov 2000). Paleozoic pelagic carbonates deposited in rather shallow water depths are represented by cephalopod and tentaculitid limestone (Sect. 15.8.2.1). The primary locus of global limestone deposition has gradually shifted from shallow cratonic to deep oceanic settings since the Jurassic. This change is generally coincident with the appearance of planktonic foraminifera and the diversification of coccolithophorids (Boss and Wilkinson 1991). The Jurassic planktonic foraminifera (Pl. 69/12) show a dramatic evolutionary radiation throughout the Cretaceous and Cenozoic (Pl. 73) and have been important components of the oceanic plankton ever since.
Benthic Carbonate Factories
The earliest quantitatively important coccolithophorids appeared in the Triassic, reached rock-building quality in the Jurassic and became the main elements of pelagic carbonate rocks during the Cretaceous and Tertiary (see Sect. 4.1.1; Fig. 15.37).
16.7.2 Differences in Phanerozoic Benthic Carbonate Factories Chapter 9 stated that ‘limestones are biological sediments’ and that most limestones are directly or indirectly influenced and controlled by biological processes. Purely physicochemical carbonate precipitation has been discussed for the formation of some micrites or ‘non-skeletal’ components like carbonate cements and ooids. Biological or organic triggers are also under discussion for the formation of these components. Schlager (2003) established a threefold conceptual framework distinguishing the various degrees of biotic influence on the precipitation of benthic marine carbonate sediments (Fig. 16.25): • Abiotic (or quasi-abiotic) precipitation with negligeable biotic effects. • Biotically induced precipitation where organisms trigger biochemical processes. The category includes precipitation driven by energy inherited from life processes (organo-mineralization; see Fig. 4.2), and precipitation driven by living organisms influencing physico-chemical conditions, leading to biomineralization (mainly microbial precipitates). These processes contribute to the formation of mud mounds and are effective in the formation of seep and vent carbonates (Sect. 16.5). • Biotically controlled precipitation where organisms determine the location, beginning and end of the process, and commonly also the composition and crystallography of the mineral. This category includes all skeletal carbonates. These processes are key elements used in a classification distinguishing three benthic carbonate factories (Fig. 16.25) representing end members related by transitions in space: • The tropical factory is characterized by (a) biotically controlled precipitates from photo-autotrophic organisms (e.g. algae) and animals with photosynthetic algae, such as hermatypic corals, certain foraminifera and certain bivalves, and (b) abiotic precipitates in the form of marine cements and ooids. Heterotrophic organisms are common as well as the construction of wave-resistant structures by organic frame-building and/or rapid
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cementation. The factory operates in warm, sunlit waters, high in oxygen and low in nutrients. • The cool-water factory is characterized by biotically controlled precipitates and dominated by heterotrophic organisms. A contribution of photo-autotrophic organisms (red algae, symbiotic larger foraminifera) may, however, be significant. The sediment typically consists of a skeletal hash of sand- to granule size. Shoalwater reefs and ooids are absent. Carbonate mud and abiotic marine cements are scarce. The environment of the cool-water factory is aphotic or photic water that is cool enough to exclude competition by the warmwater tropical factory and sufficiently winnowed to prevent the deposition of terrigenous fines. Water depths range from upper neritic to bathyal. • The mud mound factory is characterized by biotically induced and quasi-abiotic precipitates, usually micrite. The fine-grained carbonate precipitated in situ (automicrite). The typical environment of the mud-mound factory is dysphotic or aphotic, nutrient-rich waters low in oxygen but not anoxic. These conditions often prevail at intermediate depths, below the mixed layer of the sea.
Fig. 16.25. Benthic shallow carbonate factories. Proportions of abiotic, biologically induced and biologically controlled material in the tropical, mud-mound and cool-water carbonate factory. Tropical factory and mud-mound factory are mixtures of the three categories and grade into one another with the mud-mound factory centered on the induced and abiotic material, and the tropical factory centered on biologically controlled material. The cool-water factory consists almost entirely of one category. The graph is based on published compositional data or point counts of published microphotographs of thin sections. Slightly modified fromSchlager (2003).
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Significance of the benthic carbonate factory concept. The concept links properties expressed by a systematic combination of environmental parameters, sediment composition, depositional architecture and growth potential, and thus aids in predicting yet unknown attributes (e.g. porosity development) of carbonate deposits. Some of the properties appear prominently in seismic data, thus facilitating subsurface prediction. Sediment accumulations of the factories differ in composition, geometry and facies patterns (Schlager 2003). The tropical factory is characterized by flat platforms, sharp shelf breaks and steep slopes. Raised rim and empty lagoon (‘empty bucket’) are characteristic of a stress-dominated tropical system. Cool-water accumulations exhibit seaward-sloping shelves and relatively gentle slopes. The geometries of the mud-mound factory are highly variable. Groups of convex mounds are common in deeper-water settings, whereas flattopped mounds can develop in wave-swept shoal-water environments. Basic prerequisites for using the concept as a predictive tool and for interpreting vertical and lateral facies changes are qualitative and quantitative data, most of which can be derived from microfacies analysis, including • sediment composition (Sect. 6.2), • matrix and grain types (Chapter 4), • automicrite (Sect. 4.1.1), • indications of microbial contributions (Box 4.2, Pl. 8., Sect. 9.1, Box 9.1), • environment- and facies-diagnostic fossils (Chapter 10), • dominating skeletal grains and grain associations (Sect. 12.2), • paleoenvironmental constraints including paleowater temperature (Tab. 2.3, Sect. 2.4.4), differentiation of photic, dysphotic and aphotic zones (Sect. 12.1.4.1), paleo-water depth (Sect. 12.3.1), and nutrients (Box 12.8).
16.7.3 Temporal Changes in Non-Skeletal and Skeletal Mineralogy Secular variations in the relative abundance of aragonite (and High-Mg calcite) and Low-Mg calcite nonskeletal constituents (ooids, see Sect. 4.2.5; cement; micrite) of carbonate platforms were recognized by Sandberg (1983; see Sect. 7.1.5; Fig. 7.1). Aragonite (and High-Mg Calcite) were considered to dominate in the Late Precambrian to the Cambrian, the Mid-Carboniferous to the Triassic, and the Tertiary to recent intervals. Low-Mg calcite mineralogy was abundant
Microfacies Criteria in Time
in the Mid-Paleozoic, Jurassic and Cretaceous (Bates and Brand 1990). Fluctuations in marine carbonate mineralogy are generally acknowledged, but different opinions exist with regard to the length and the boundaries of the time intervals characterized by particular mineralogies. The oscillating trend (Sandberg cycle) is usually interpreted as an indication of major variations in paleoceanographic parameters. The temporal distribution of once aragonitic non-skeletal constituents appears to be discontinuous in the Phanerozoic (Sandberg 1991). Controls of nonskeletal mineralogy by pCO2 and covariant factors (e.g. oceanic Ca/Mg) are suggested by the close agreement between sea-level curves and the presence-absence curves for non-skeletal aragonite. Van de Poel and Schlager (1994) studied the mineralogy of Mesozoic to Cenozoic skeletal grains contributing to the formation of carbonate platforms and found that the biogenic aragonite-calcite intervals parallel those observed in nonskeletal constituents. Large-scale patterns in the skeletal mineralogy of major carbonate producers (Wilkinson 1979; Stanley and Hardie 1998) have been linked with plate tectonic evolution and changing Mg/Ca ratios in seawater (Stanley and Hardie 1999; Steuber 2002). The link between the reconstructed chemical composition of seawater and the skeletal mineralogy of reef-building organisms (Stanley and Hardie 1998) seems to be less than originally suggested (see Kiessling 2002 and collected papers in Kiessling et al. 2002). The StanleyHardie model may work better for off-reef carbonate platforms, as indicated by the coincidence of foraminiferal skeletal mineralogy with aragonite-calcite intervals (Martin 1995) and the more frequent occurrence of calcareous algae with High-Mg-calcite and aragonite skeletons in aragonitic intervals (Kiessling et al. 2003). Micritic matrix, grains and fossils, and pore-filling cement of limestones are usually preserved as calcite. An understanding of the primary mineralogy requires a thorough study of microfacies criteria allowing precursor mineralogy (Sect. 7.1.5.2) and skeletal diagenesis (Sect. 4.2.1) to be evaluated.
16.7.4 Temporal Changes in the Abundance and Significance of Microfacies Criteria Although Standard Microfacies Types are similar (compare Fig. 14.28) and the main grain types differentiated in microfacies studies can be observed all over the Phanerozoic, the abundance and significance of some microfacies criteria used in interpreting depositional environments vary through time. The consider-
Microfacies Criteria in Time
ation of these temporal changes is crucial in using microfacies analysis as an essential tool in understanding the formation and properties of carbonate rocks. Micrite. Different possibilities for the origin of micrite were discussed in Sect. 4.1.1 and summarized in Fig. 4.1. Biologically induced automicrite formed in place at the sea bottom or within the sediment and triggered by organomineralization or microbial processes has been known since the Proterozoic, whereby the formation by metabolic processes of cyanobacteria started near the Precambrian-Cambrian boundary and was a common process in the Early Paleozoic (Pratt 2001). The formation of allomicrites through the disintegration of the skeletons of benthic biota started in the Cambrian with micrites formed by disintegrated invertebrate skeleton material. This mode of micrite formation prevailed throughout the Paleozoic. Micrite formation through abrasion of shells increased from the Late Paleozoic (Gischler and Zingeler 2002). Allomicrites originating from the disintegration of calcareous algae (Halimeda model; Fig. 4.3) became important in the Mesozoic and increased in abundance during the Cenozoic (Fig. 10.10). Allomicrite formation caused by bioerosion by microborers was low in the Paleozoic, but became an important process from the Mesozoic (Fig. 9.15). The greatest change in allomicrite formation took place in the Jurassic with the appearance and radiation of calcareous plankton (foraminifera; coccolithophorids and other nannofossils), forming pelagic carbonates throughout the Mesozoic and Cenozoic. Peloids can be formed by various processes (Sect. 4.2.2; Fig. 4.11). In-situ formation (precipitated peloids and microbial peloids) and mud peloids occurred as early as the Precambrian. Peloids of biotic origin (algal peloids, bioerosional peloids and fecal pellets) and bahamite peloids appeared in the Cambrian. Algal peloids are common in Early Paleozoic carbonates, fecal pellets are abundant in Mesozoic limestones. The abundance of peloids in platform carbonates varied. Peloids are usually more common than ooids and were abundant in the Devonian, Early Carboniferous, Middle Triassic, Late Jurassic and Early Cretaceous. Ooids. Ooid abundance varied substantially throughout the Phanerozoic. Times of high ooid contributions to platforms (Late Cambrian-Early Ordovician, Early Carboniferous, Jurassic) do not coincide with times characterized by specific mineralogical compositions or specific geographical distribution patterns of platforms. Giant centimeter-sized ooids are a particular feature of Precambrian carbonates.
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Oncoids. The composition and depositional settings of oncoids have changed over time (Sect. 4.2.4.1). Microbial oncoids and algal oncoids have contributed to subtidal and intertidal platform carbonates since the Precambrian. Spongiostromate oncoids are common constituents of Proterozoic platforms. The first oncoids formed by cyanobacteria (cyanoids) appeared in the Late Vendian. Spongiostromate and porostromate oncoids differ in their temporal distribution. Paleozoic spongiostromate oncoids occurred predominantly in lacustrine and transitional marine environments, whereas porostromate oncoids flourished in marine subtidal environments from the Cambrian to the Jurassic. These oncoids became more frequent in the Late Paleozoic, and peaked in the Late Jurassic (see Box 4.10). A major change in the environmental setting of porostromate oncoids during the Jurassic must be considered, using Late Mesozoic and Cenozoic oncoids as paleoenvironmental indicators: The last normal-marine cyanoids are known from the Cretaceous. Spongiostromate oncoids became the only type in marine settings, but continued to be abundant in lacustrine environments until today. Starting in the Late Cretaceous and Early Tertiary, spongiostromate oncoids were increasingly replaced by coralline algal rhodoids (Sect. 4.2.4.2). Rhodoids occurred as early as the Late Paleozoic, but did not become important constituents of carbonate platforms until the Tertiary. Skeletal grains. The association and abundance of skeletal grain types contributing to the formation of carbonate rocks varied through the Phanerozoic (Box 10.1; Fig. 10.3). The widespread deposition of carbonate rocks coincides with the development and diversification of invertebrates at the beginning of the Cambrian. Common skeletal grains in Early Paleozoic carbonates are bryozoans, brachiopods, echinoderms and ostracods. A long-term increase in the contribution of calcareous algae, foraminifera and mollusks to the formation of shelf carbonates started in the Late Paleozoic and continued during the Mesozoic and Cenozoic. The abundance and carbonate production by calcareous algae varied considerably (Sect. 10.2.1). Different major algal groups flourished at different time intervals (Fig. 10.4) as did benthic foraminifera (Sect. 10.2.2). Late Paleozoic, Late Mesozoic and Cenozoic larger foraminifera exhibit specific environmental distributional patterns that must be considered, using the fossils interpret depositional settings and evaluate standard facies models (Fig. 14.7). Using shell beds as paleoenvironmental proxies must take into account the differences in life habits, taphonomic behavior and skeletal mineralogy of brachiopods and bivalves (Sect.
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10.2.4.1 and Sect. 10.2.4.2). Similarly, the evaluation of crinoid accumulations should regard the changes in preferred environments and associated differences in taphonomy over time (Sect. 10.2.4.9). Burrowing and bioturbation. Using burrowing or bioturbation fabrics (Sect. 5.1.4) as paleoenvironmental proxies, e.g. in recognizing substrate types (Sect. 12.1.3.2) means considering significant changes at the Precambrian-Cambrian boundary and the quantitative increase during the Phanerozoic (Bottjer and Droser 1994). Basics: Secular variations in carbonate deposition Barnes, C.R. (1999): Paleocenography and paleoclimatology: an Earth system perspective. – Chemical Geology, 161, 17-35 Bates, N.R., Brand, U. (1990): Secular variations of calcium carbonate mineralogy: an evaluation of ooid and micrite chemistries. – Geologische Rundschau, 79, 27-46 Boss, S.K., Wilkinson, B.H. (1991): Planktogenic/eustatic control on cratonic/oceanic carbonate accumulation. – Journal of Geology, 99, 497-513 Bosscher, H., Schlager, W. (1993): Accumulation rates of carbonate platforms. – Journal of Geology, 101, 345-355 Burns, S.J., McKenzie, J.A., Casconcelos, G. (2000): Dolomite formation and biogeochemical cycles in the Phanerozoic. – Sedimentology, 47, 49-61 Cowen, R. (1988): The role of algal symbiosis in reefs through time. – Palaios, 3, 221-227 Fischer, A.G. (1982): The two Phanerozoic supercycles. – In: Bergren, W.A., van Couvering, J.A. (eds.): Catastrophes and Earth history. – 129-150, Princeton (Princeton University Press) Fischer, A.G., Arthur, M.A. (1977): Secular variation in the pelagic realm. – In: Cook, H.E., Enos, P. (eds.): Deepwater carbonate environments. – Society of Economic Paleontologists and Mineralogists, Special Publication, 25, 19-50 Flügel, E., Kiessling, W. (2002): Patterns of Phanerozoic reef crises. – In: Kiessling, W., Flügel, E., Golonka, J. (eds.): Phanerozoic Reef Patterns. – SEPM, Special Publication, 72, 691-733 Grotzinger, J.P., James, N.P. (eds., 2000): Precambrian carbonates: evolution of understanding – SEPM, Special Publication, 67, 3-20 Hay, W.W. (1999): Carbonate sedimentation through the late Precambrian and Phanerozoic. – Zentralblatt für Geologie und Paläontologie, Teil I, 1998, 435-445 Kiessling, W. (2001): Paleoclimatic significance of Phanerozoic reefs. – Geology, 29, 751-754 Kiessling, W. (2002): Secular variations in the Phanerozoic reef ecosystem. – In. Kiessling, W., Flügel, E., Golonka, J. (eds., 2002): Phanerozoic Reef Patterns. – SEPM, Special Publication, 72, 625-690 Kiessling, W., Flügel, E., Golonka, J. (2000): Fluctuations in the carbonate production of Phanerozoic reefs. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems; components and interactions. – Geological Soc. London, Special Publication, 178, 191-215 Kiessling, W., Flügel, E., Golonka, J. (2003): Patterns of carbonate platform sedimentation. – Lethaia, 36, 195-225
Basics:Variations in Carbonate Deposition
Kuznetsov, V. G. (2000): Some features of the evolution of carbonate accumulation in earth’s history. Communication 1. Evolution of the intensity, mechanism and setting of carbonate accumulations. – Lithology and Mineral Resources, 35, 32-46 Mackenzie, F.T., Morse, J.W. (1992): Sedimentary carbonates through Phanerozoic time. – Geochimica et Cosmochimica Acta, 56, 3281-3295 Martin, R.E. (1995): Cyclic and secular variation in microfossil biomineralization: clues to the biogeochemical evolution of Phanerozoic oceans. – Global and Planetary Change, 11, 1-12 Milliman, J.D., Droxler, A.W. (1986): Neritic and pelagic carbonate sedimentation in the marine environment: ignorance is not bliss. – Geol. Rundschau, 85, 496-504 Montanez, I.P. (2002): Biological skeletal carbonate record changes in major-ion chemistry of the paleo-oceans. – PNAS, 99, 15832-15854 Sandberg, P.A. (1983): An oscillating trend in Phanerozoic non-skeletal carbonate mineralogy. – Nature, 305, 19-22 Sandberg, P.A. (1985): Aragonite cements and their occurrence in ancient limestones. – In: Schneidermann, N., Harris, P.M. (eds.): Carbonate cements. – Society of Economic Paleontologists and Mineralogists, Special Publication, 36, 33-57 Sandberg, P.A. (1991): Nonskeletal aragonite and pCO2 in the Phanerozoic and Proterozoic. – In: Sundquist, E.T., Broecker, W.S. (eds.): The carbon cycle and atmospheric CO2: Natural variations Archean to Present. – Geophysical Monograph, 32,585-594 Schlager, W. (2000): Sedimentation rates and growth potential of tropical, cool-water and mud-mound carbonate systems. – In: Insalaco, E., Skelton, P.W., Palmer, T.J. (eds.): Carbonate platform systems; components and interactions. – Geological Society of London, Special Publication, 178, 217-227 Schlager, W. (2003): Benthic carbonate factories in the Phanerozoic. – Int. Journal of Earth Sciences, 92, 445-464 Stanley, S.M., Hardie, L.A. (1998): Secular oscillations in the carbonate mineralogy of reef-building and sedimentproducing organisms driven by tectonically forced shifts in sea-water chemistry. – Palaeogeography, Palaeoclimatology, Palaeoecology, 144, 3-19 Stanley, S.M., Hardie, L.A. (1999): Hypercalcification: paleontology links plate tectonics and geochemistry to sedimentology. – GSA Today, 9, 1-7 Steuber, T. (2002): Plate tectonic control on the evolution of Cretaceous platform-carbonate production. – Geology, 30, 259-262 Van de Poel, H.H., Schlager, W. (1994): Variations in Mesozoic and Cenozoic skeletal carbonate mineralogy. – Geologie en Mijnbouw, 73, 31-51 Walker, L.C., Wilkinson, B.H., Ivany, L.C. (2002): Continental drift and Phanerozoic carbonate accumulation in shallow-shelf and deep-marine settings. – Journal of Geology, 110, 75-87 Wilkinson, B.H., Given, R.K. (1986): Secular variation in abiotic marine carbonates: constraints on Phanerozoic atmospheric carbon dioxide contents and oceanic Mg/Ca ratios. – Journal of Geology, 94, 321-333 Wilkinson, B.H., Walker, J.C.G. (1989): Phanerozoic cycling of sedimentary carbonate. – American Journal of Science, 289, 525-548 Further reading: K185
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PRACTICAL USE OF MICROFACIES The third part of the book provides insights into applied aspects of microfacies studies. Chapter 17 considers the role of microfacies in reservoir rock analyses of hydrocarbon reservoirs and carbonate-hosted base metal deposits. Chapter 18 deals with carbonate resources and rock properties and their depositional and diagenetic controls. Chapter 19 emphasizes a very promising approach, the use of microfacies in archaeology.
17 Reservoir Rocks and Host Rocks Reservoir rocks are characterized by petrophysical properties favoring the capacity for the migration and storage of fluids, gas and minerals. Microfacies analysis has been strongly triggered by the exploration and development of hydrocarbon reservoirs and source rocks in limestone and dolomites and is still of interest in understanding depositional controls, facies heterogeneities and reservoir property distribution. Carbonates can also be host rocks for ore deposits. Patterns of syn- and epigenetic mineralizations become more understandable from the use of microfacies data.
17.1 Carbonate Hydrocarbon Reservoirs Estimations of the proportion of carbonate rocks to the total of world’s hydrocarbon production vary between about 40% and more than 60%. Carbonate rocks contain the world’s largest reservoirs in the Middle East (e.g. Ghawar Field, Saudi Arabia) and Kazakhstan (e.g. Kazagan Field). Facies analysis is a critical exploration tool. The objective of reservoir studies is to construct reservoir models that describe or predict porosity and permeability distribution. Carbonate reservoirs can be grouped according to their depositional setting or diagenesis (Roehl and Choquette 1985). Common reservoirs tied to depositional settings are located in • dolomitized platform interior or ramp sequences (e.g. Williston Basin, Dakota), • oolitic and bioclastic sand bodies forming shoals and
sheets on shelves and ramps (Smackover Formation, USA), • reefs and mounds formed on platforms, ramps and shelves, and platform-margin reefs (Malampaya, Philippines), • downslope debris and allochthonous platform and reef material intercalated within basinal sediments (Poza Rica, Golden Lane, Mexico), and • chalks formed on deeper shelves and basins (Cretaceous Chalk, North Sea). Carbonate reservoirs linked strongly to diagenetic processes are related to • subunconformities, subaerial exposure and karst, often associated with leaching processes (Camposta, Spain), • early and late diagenetic dolomitization (Arab ABC Formation, Middle East), • fracturing (Maracaibo Platform, Venezuela), and • burial diagenesis (Devonian, Alberta). Being able to recognize both the depositional facies and the diagenetic history is essential in understanding the creation and preservation of carbonate reservoirs. Facies analysis helps in identifying environments and establishing depositional models including prediction of reservoir extent and rock types. The theme of volume 83/11 (1999) of the AAPG Bulletin was ‘Carbonate Reservoirs: Strategies for the Future’. The papers in this issue highlight the work required for understanding the external envelope and internal architecture of carbonate reservoirs as well as the spatial distribution of porosity and permeability
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_17, © Springer-Verlag Berlin Heidelberg 2010
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within the reservoir units. Current and future work requires that multidisciplinary data, analytical techniques and interpretations be integrated. This includes outcrop studies of field analogs, seismic modeling, wireline log modeling, geochemical stratigraphy, geomechanical and special core analysis. The description and quantification of reservoir heterogeneities require integrating depositional, diagenetic, and petrophysical modeling at various scales. What place does microfacies have? Carbonate reservoirs should be investigated at all scales. Reservoir studies need an understanding of large-scale depositional environments and geometrical patterns and medium- to small-scale variations in lithology, depositional facies and diagenetic overprints responsible for the heterogeneity of petrophysical properties of carbonate reservoirs. Depositional environments are evaluated from the sedimentological and palecological criteria reflected by microfacies criteria. Carbonate geometries and cyclic depositional changes are analyzed by high-resolution sequence stratigraphy. Outcrop analog studies exhibit size, internal architecture, composition and diagenetic history of carbonate bodies. Composition and di-
17.1.1 Distribution of Carbonate Reservoirs during Time
Number of carbonate reservoirs
K
K-TT
Carbonate reservoirs are known from the Precambrian and all through the Phanerozoic, but are concentrated within several intervals (Greenlee and Lehman 1993; Kiessling et al. 1999; Kiessling 2002). Fig. 17.1 shows the temporal distribution of carbonate reservoirs located in various depositional settings (see Fig. 17.3). Fig. 17.2 is limited to reef carbonates with reservoir quality. The principal times of reef reservoirs agree approximately with the maximum reef expansion. Both figures demonstrate a conspicuous heterogeneous distribution, distinct peaks and times with rare reservoirs. Peaks occur in the Late Devonian, Late Carboniferous and Early Permian, Late Jurassic, Mid-Cretaceous and Miocene.
J-K
PC C O S D C P T J
agenesis are recorded by microfacies criteria. Basic differentiation of textural carbonate types requires thinsection studies. Thin sections are the key to understanding the geometry of major pore types and distinguishing effective from ineffective porosity. Today integrating qualitative and quantitative microfacies data, analog studies, petrographic core analysis, seismic analysis and high-resolution reservoir-scale correlation, and 3D modeling form the basis for the characterization of carbonate fields (Grötsch et al. 1998; Grötsch and Mercadier 1999; Grötsch et al. 2003). The currently most challenging hurdle in modeling carbonate hydrocarbon reservoirs is the linkage between the static reservoir description and the dynamic model. With recent advances made in dynamic model resolution (>1 million cells), the question of how to add dynamic rock properties as measured by special core analysis (e.g. drainage-inbibition capillary measure, relative permeability) to each model cell in a consistant fashion became of crucial importance (Grötsch et al. 2003). An example is the Kharaib Formation in the Middle East where a strong increase in permeability towards the top of the formation results in water overriding oil over large distances despite the presence of gravitational forces (Grötsch et al. 1998).
TT
Fig. 17.1. Temporal distribution of carbonate reservoirs. Age of 330 carbonate reservoirs all over the world studied by C & C Reservoirs, Inc.. Reservoirs occur in all time intervals but are abundant in the Devonian, Late Paleozoic, Jurassic, Cretaceous and Tertiary. Note that the height of the columns does not reflect the relative abundance, but rather the general pattern. Based on internet data (www. ccreservoirs.com).
17.1.2 Depositional Setting and Environmental Controls of Carbonate Reservoirs 17.1.2.1 Depositional Setting The distribution of primary porosity is strongly faciescontrolled as illustrated by the Great Bahama Bank model (Sect. 7.4.4; Fig. 7.13). Local facies variations
Reservoirs: Depositional Settings
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Fig. 17.2. Temporal distribution of reefs with reservoir quality calculated for stages. The figure includes subsurface reefs of producing reservoirs as well as outcrop data from reefs with reservoir potential (high porosity, thickness of at least 10 m) but lacking hydrocarbon accumulation. After Kiessling (2002).
result in rapidly changing porosity and permeability values over short distances. Reefs and organic banks and shallow-marine oolitic and bioclastic sand shoals are often cited as common hydrocarbon exploration targets, but reservoirs are known from a wide range of depositional environments including lacustrine and karst settings. Marine reservoirs occur in carbonates formed in lagoonal and open parts of platforms, inner and outer ramps, and slope and basinal settings (e.g. debris deposits, turbidites and chalks; Fig. 17.3). Favored settings for carbonate reservoirs on shelves and on ramps are • subtidal and intertidal complexes with updip facies changes and downdip dolomitization (Paris Basin), • lime sand bodies on shelves with incomplete cementation (Persian Gulf), • organic buildups that may be dolomitized, including reef mounds on the shelf, at platform-margins, outer shelf/ramp pinnacles and atolls (Devonian, Alberta; Malampaya, Philippines), • talus accumulations near reefs and banks (Midland Basin), and • leached zones below unconformities with or without fracturing (Paradox Basin). Reservoir rocks formed in different settings exhibit different depositional criteria that influence initial porosities and permeabilities. These properties are further differentiated and strongly controlled by early and
late diagenetic processes (Chap. 7). The following cursory list gives an idea of reservoir characteristics within different depositional environments: • Near-shore environments, tidal flats. Carbonate muds and sands. Reservoir geometry ribbon- or sheetlike. Grain size fine to medium. Intercrystalline and intergranular porosity. Reservoirs in shoreline sands and high-energy offshore bars of inner ramps and platforms and in tidal flat sediments. Tidal sediments have usually low porosities and permeabilities, but dolomitization can create higher porosities and permeabilities. Associated evaporites may act as seals. Examples are Ordovician reservoirs of the Williston Basin and reservoirs in the Permian Basin of Texas. • Restricted and nonrestricted lagoons. Muddy and grainy sediments. Reservoirs discontinuous, often as stacked lenses. Pores small, permeability low. Intergranular and moldic porosity. Porosity enhanced by dolomitization (e.g. central lagoonal part of Malampaya buildup). • Open platforms. Carbonate sands, muds and small platform reefs and mounds. Variable grain sizes. Intergranular, moldic, vuggy, intragranular and intercrystalline porosity common. Low to good permeabilities. Well stratified reservoirs. Reservoir geometry sheetand ribbonlike. The development of sand bodies requires sand-sized sediments and removing of fines by
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Environmental Controls 5
10
Karst-related detrital wedge Shore line Offshore bar Tidal flat Marine embayment Shelf lagoon Open shelf Skeletal bank Mud-rich skeletal bank Coquina bank High-energy ramp Low energy ramp Patch reef Reef mound Mud-rich reef mound Pinnacle reef Platform/ramp marginal shoal Barrier and fringing reef Debris flows/turbidites Pelagic deposits
15%
Fig. 17.3. Spatial distribution of carbonate reservoirs. Depositional setting of 315 Late Precambrian to Tertiary carbonate reservoirs studied all over the world by C & C Reservoirs, Inc. The figure gives only a rough picture of the relative frequency, but underlines the importance of grainy deposits (offshore bar, high-energy ramp, platform/ramp marginal shoal, debris flows/turbidite), muddy and grainy deposits (tidal flat, shelf lagoon, low-energy ramp) and reefs particularly reef mounds. Note that karst-related reservoirs and pelagic reservoirs (e.g. chalks), are much more important than indicated in this figure! Based on internet data (www.ccreservoirs.com).
wave action or strong tidal currents. Sand bodies occur as tidal bars, storm layers in lagoons, storm washover sediments, in tidal channels, at bank edges, and behind reefs. Carbonate sands have high initial porosities that may be preserved in the subsurface. Secondary porosity may develop shortly after deposition by freshwater diagenesis or later by dolomitization. The best examples of production from sand shoals are the Smackover Formation of the U.S. Gulf Coast and the Arab D Formation forming the reservoir facies of the world’s largest fields, sealed by evaporites (e.g. Ghawar field, Saudi Arabia).
ments which act as updip seal in stratigraphic traps. Examples of reefs acting as excellent reservoirs are the Oligocene reef complex of Kirkuk (Iraq) and the Devonian Leduc reef in western Canada.
• Platform margins. Carbonate sands or reefs. Grain size coarse to medium. Ribbon- and mound-like reservoirs. Intergranular and vuggy porosity in sand shoals. Sand accumulations of reservoir size commonly occur near the seaward edge of banks, platforms and shelves. When dolomitized, intercrystalline and vuggy porosity. Good porosity in backreef sediments and in dolomitized reefs. The reservoir potential of reefs is widely assumed to be high. This is an assumption made during many exploration plays. However it is often not true. Examples with tight margins are the Malampaya/ Camago buildup and the Central Luconia buildups (Sarawak). Often early marine cementation at oceanward slopes occludes porosity in a very early stage. Effective porosity is due to solution and fracture, and preserved framework pores and molds. Proximal backreef sands can retain significant amounts of primary porosity. Good permeability. Forereef deposits and aprons of mud mounds may have good reservoir potential, especially if the reef itself is plugged by ce-
• Basins. This setting is characterized by carbonate muds with intercalations of allochthonous carbonate sands as debris flows and turbidites. The porous parts of turbidites may act as sheetlike reservoirs or conduits for migrating oil (e.g. Cretaceous Golden Lane field, Mexico), Deep-sea micrites generally have low porosities and permeabilities, but are susceptible to fracturing and can form significant reservoirs. Primary porosity due to loosely packed nannofossils and open foraminiferal chambers is very high in chalks, but permeability is low. However, chalks can be significant reservoirs in the Cretaceous and Tertiary of the North Sea and U.S. Gulf Coast due to combinations of preserved primary porosity, moldic porosity and fracture porosity. Chalk reservoirs have proven to be finer grained, more porous, less permeable, more uniform and widespread, and considerably more predictable in terms of petrophysical properties than reservoirs in typical shallowwater limestones. Burial depth is the single most im-
• Slopes. Highly variable sediments. Coarse and fine sediments and breccias. Small but thick downslope reefs (e.g. mud mounds). Proximal reef slopes with intergranular and moldic porosity, distal slope with intercrystalline and moldic porosity. Both parts may exhibit high permeabilities. Reservoirs are ribbon- and mound-like (e.g. Cretaceous, Mexico).
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Environmental Control
portant factor. High porosity is caused by the specific composition (calcitic nannofossils; Pl. 29/5) and micropores (Feazel and Farell 1988). Porosity occlusion is due to overburden stress. Major hydrocarbon accumulations are common in pinnacle reefs and mud buildups formed in deeper-water basinal settings. The porous core facies is usually bounded on all sides by impermeable flank and margin deposits, basinal shales or basinal evaporites. These deposits can form a seal. Examples are known from the Silurian of Michigan and the Devonian of Canada.
17.1.2.2 Environmental Controls Environmental controls primarily affecting the heterogeneity of carbonate reservoir rocks are • hydrodynamic conditions (Sect. 12.1.1) reflected by grain- or mud-supported sediments (Pl. 43) with different primary porosities, • paleowater depth (Sect. 12.3) influencing sedimentation patterns, • high-frequency sea-level fluctuations resulting in cyclic sedimentation (Sect. 16.1) and favoring subaerial exposure and meteoric diagenetic processes, • wind- and storm-related processes shaping the geometry of reefs and producing talus deposits (Sect. 12.1.2), • climatic factors (Sect. 12.2) that contribute to different controls in the development of reservoir rocks in arid and humid regions (Dmietrievsky and Kuznetsov 1992). The influence of prevailing climates producing different styles of diagenesis has been demonstrated for Miocene reservoir rocks (Sun and Esteban 1994): Reefs and banks formed in humid tropical and subtropical settings (Miocene of Southeast Asia) are characterized by common subaerial exposure and meteoric diagenesis. The best reservoir qualities are developed beneath subaerial unconformities, where the effects of freshwater leaching and karstification were most intensive. In contrast, lagoonal ramp carbonates and reefs formed in arid, landlocked temperate-subtropical settings (Mediterranean, Gulf of Suez and Red Sea) with periodically increased salinity are characterized by dolomitization, leaching of metastable skeletal grains, generation of moldic, vuggy and intercrystalline porosity, and widespread anhydrite cement, precipitated over a wide range of burial depths. Porosity development in these reservoirs depends not only on the original depositional fabric, but more importantly on the degree to which the porosity-occluding anhydrite cement was
dissolved or the low-permeability carbonates were fractured. Environmental controls may also be closely linked with diagenetic controls. An example is the early marine cementation of isolated carbonate platform slopes as a result of ocean currents and the enhanced water rock ratio (e.g. Bahama platform; Malampaya).
17.1.3 Diagenetic Controls on Carbonate Reservoirs Meteoric, marine and burial diagenesis may either destroy or enhance original porosity depending on compaction, cementation and dissolution. Pores are often occluded by cements (Sect. 7.4), but can be open due to specific diagenetic processes or cementation is inhibited, as in the case of oil emplacement (Neilson et al. 1998).
17.1.3.1 Reservoir Properties Major reservoir property controls are permeability and porosity. Permeability is often more important than porosity. Porosity was discussed in Sect. 7.3. The terms used to describe visible porosity are summarized in Box 7.4. Fig. 7.5 shows the commonly used Choquette and Pray porosity classification (Sect. 7.3.2). Pore geometry and permeability are discussed in Sect. 7.3.1.2, porosity measurements and pore types seen in thin sections in Sect. 7.3.1.3. The thin-section aspect of common pore types is shown in Pl. 29 and Pl. 30. Porosity has a great influence on the saturation and mobility of hydrocarbons in carbonate rocks. It is determined by conventional core plug analysis, and capillary-pressure curves. SEM studies of three-dimensional resin casts, and the investigation of thin sections exhibit the size and shape of the pores. A correlation of thin-section derived porosity data and physical porosity measurements is possible using digitalized thin-section measurements (McCreesh et al. 1991). Reservoir characteristics depend on the arrangement of the pores and how pores are interconnected by pore throats. Pore throat sizes and distributions depending on the size and shape of grains and crystals are determined by capillary-pressure curves (Wardlaw and Taylor 1976). Pore throats are often too small to be examined visually in thin sections (Wardlaw and Li 1981). The size, shape, arrangement, and distribution of pore throats are major controls on reservoir-rock producibility and recovery efficiency, and must be included in the differentiation of porosity types. Dynamic proper-
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ties of different pore geometries are measured by special core analysis (SCAL), e.g. drainage and inbibition capillary pressure measurements and relative permeability curves of hydrocarbon to water. These pore-geometry dependant parameters are a crucial input for dynamic 3D reservoir models. Pore systems are effective and ineffective with regard to reservoir qualities. Effective porosity is characterized by interconnected pores including interparticle pores caused by leaching of interparticle cement, micropores between calcite or dolomite crystals, and moldic pores associated with interparticle pores. Intercrystalline microporosity is common in many productive Mesozoic and Tertiary reservoirs in the Middle East (Moshier 1989; Budd 1989; Fig. 7.4). Effective porosity depends on pore throat diameter, pore size, reservoir pressure, temperature, and the viscosity of oil or gas. Burrows may cause effective or ineffective porosity. Intraskeletal porosity, e.g. in chambers of foraminifera, is often ineffective. 17.1.3.2 Diagenetic Controls on Reservoir Properties Most reservoir porosity is caused by various diagenetic processes including near-surface dissolution, early and late dolomitization and burial fracturing. This secondary porosity often occurs together with preserved facies-controlled primary porosity. The understanding of the porosity development through time and space is crucial in evaluating reservoir qualities. Studies of porosity development must take into account marked differences in the accumulation of sediments in different carbonate settings. Fast growing platform-margin reefs, e.g., usually have high initial porosities, but resulting reef limestones often exhibit low porosities because of rapid early cementation (e.g. Malampaya). In contrast, low-growing interior platforms may have high porosities due to widespread dissolution and dolomitization. Common diagenetically controlled reservoir types include: • Reservoirs related to near-surface dissolution at subaerial unconformities. Reservoirs associated with unconformities have been estimated to contain 20 to 30% of known recoverable hydrocarbons (Halbouty et al. 1970). Subaerial exposure and meteoric diagenesis are important processes in the development and preservation of porosity. The creation and preservation of pore systems depend on the depositional facies and is related to the duration of subaerial exposure (Hopkins 1999; Saller et al. 1999). Diagenetic processes during subaerial exposure rearrange the pore network, but do
Dolomite Reservoirs
not necessarily increase total porosity. In many cases porosity is reduced below unconformities. Permeability is more strongly changed than porosity. It may increase or decrease. Pore systems evolve with the time of subaerial exposure. Shorter periods of subaerial exposure (10,000 to 400,000 years) are often associated with greater porosity than long intervals (1 to 20 million yr). Many carbonates subjected to subaerial exposure have little or no porosity in the deep subsurface due to compaction, cementation and stratal collapse reducing burial porosity. But unconformity-related diagenesis may also enhance reservoir porosity by creating pore systems that are resistant to deeper burial compaction. Caverns, breccias and vugs can form in the deep subsurface because of dissolution by basinal fluids or the controls of shales overlying carbonates. • Karst reservoirs. Platform carbonates, affected by an interplay of eustatic transgressions and regressions with local tectonic uplifts, are major sites of karstification, often connected with dolomitization. Both processes may lead to the development of karst reservoir rocks (Cater and Gillcrist 1994). Reservoirs produced by near-surface karstic processes, the collapse of paleocaves and later burial compaction and dissolution may be hundreds to thousands of meters across. These reservoirs are complex and heterogeneous and characterized by fractures, breccias and fill sediments. Breccias and cave sediments need specific classifications. Loucks (1999) distinguishes highly fractured crackle breccias and chaotic breccias with rotated and displaced clasts. Predicting the size and distribution of paleokarst features is generally not possible if well data are used. Understanding paleokarst heterogeneity requires outcrop data and the interpretation of karst horizons within a sequence stratigraphic framework. For an overview on karst reservoirs see Fritz et al. (1993). • Dolomite reservoirs. Early dolomitization often results in better preservation of reservoir properties during diagenesis (Arab ABC, Middle East). Dolomitization may increase porosity, but can also decrease it, e.g. in near-surface dolomitization. Dolomite reservoirs are very heterogeneous at different scales (Weber 1986). At a scale of 10 m to 10 km heterogeneities are represented by major lithostratigraphic boundaries, large faults and extensive fractures. At a scale of millimeters to meters heterogeneities are caused by the rock fabric. Reservoir characteristics of the major dolomite reservoirs depend upon the original sediment fabric, the mechanism forming the dolomite, and the extent to which early formed dolomite was modified by postdolomitization diagenetic processes (e.g. karstification and fracturing). Porosity and permeability in platform
Fractured Reservoirs
Fig. 17.4. Fractures and microcracks can create permeability and porosity in low-porosity limestones. The SEM photograph shows a submicroscopic microcrack partly filled with heavy oil, being a part of a network of small fractures within bioclastic wackestones. Identifying cracks is a key point for the characterization of fracture reservoirs. Late Cretaceous (Mishrif Formation): Subsurface, Middle East.
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dolomites depend on the location within the dolomitizing system (Saller and Henderson 1998). Porosity and permeability are low in platform interior, and higher near platform margin settings. Sun (1995), summarizing the published records of hydrocarbon-producing dolomite reservoirs, distinguished four basic types: (1) Peritidal-dominated carbonates, characterized by low matrix porosity and permeability and strong dolomitization. Economic reservoirs are possible if the rocks are karstified and fractured. (2) Subtidal dolomite associated with evaporitic tidal flat/lagoons, typically formed in shallowing-upward cycles. The best reservoirs occur in high-energy, grain-dominated subtidal carbonates forming the lower parts of the cycles, or in low-energy mudstone/wackestone with a coarse crystalline fabric. Primary intergranular porosity is preserved due to the relatively arid climate and because impermeable tidal-flat dolomite and anhydrite, forming the caps of the upward-shoaling cycles seal off the subtidal carbonates. (3) Subtidal dolomite associated with basinal evaporite commonly bears no direct relationship to depositional trends due to the strong overprint of pri-
Box 17.1. Selected carbonate reservoir studies. A list of studies made in the 1950s to1970s can be found in Mazzullo (1980). Lacustrine carbonates: Katz 1990; Lomando et al. 1994 Reef reservoirs: Ahr and Ross 1982; Alam et al. 1999; Shelf carbonates: Burchette and Britton 1985; Chidsey Alsharhan 1995; Amthor et al. 1994; Armstrong et al. 1980; Baria et al. 1982; Crevello et al. 1985; Cys and et al. 1996; Croft et al. 1980; Grötsch et al. 1999, 2003; Jordan et al. 1985; Kopaska-Merkel 1994; KopaskaMazzullo 1985; Gill 1985; Grötsch and Mercadier 1999; Merkel et al. 1994; Lindsay and Kendall 1985; Lomando Herrod et al. 1985; Jardine and Wilshart 1987; Jordan and Abdullah 1988; Klovan 1964; Kuznetsov 1997; et al. 1985; Loucks and Anderson 1985; Madi et al. 2000; Miller 1985; Noack and Schroeder 2003; VahrenLasemi et al. 1994; Longman 1985; Malek-Aslani 1985; kamp and Grötsch 1994 Nayoan et al. 1981; Pomar and Ward 1999; Precht 1986; Roylance 1990; Schmidt et al. 1985; Stemmerik and Shoals: Ahr and Walters 1985; Asquith 1986; Budd 2002; Burchette et al. 1990; Carozzi and Reichelderfer 1987; Larsen 1993; Sun and Wright 1998; Weidlich 2002; WilChoquette and Steinen 1985; Handford 1988; Hovorka son et al. 1997. For further references compare also Box 7.7! et al. 1993; Keith and Zuppann 1993; Kelleher and Smosna 1993; Longacre and Ginger 1988; Moore et al. Turbidites: Duvernory and Reulet 1977; Cazzola and 1986; Parham and Sutterlin 1993; Sellwood and Beckett Soudt 1993; Enos 1985, 1988; Hertig et al. 1994 Chalks: D’Heur 1984; Feazel et al. 1985, 1988; Hardman 1991; Sellwood et al. 1989 Subunconformity reservoirs: Budd et al.1995; Hopkins 1982; Maliva and Dickson 1992; Scholle 1977; Van den 1999; Montgomery et al. 2000 Bark and Thomas 1980 Subaerially exposed carbonates: Budd et al. 1995; Fracture reservoirs: Aguillra 1995; Barchi et al. 1993; Dickson and Saller 1995; Esteban and Klappa 1983, Irish and Kempthorne 1987; Kulander et al. 1990; 1993; Friedman 1980; Hopkins 1999; Kerans 1997; Longman 1985; McQuillan 1985; Narr and Curie 1984; Nelson 1985; Nelson et al. 2000; Reijers and Bartok Longman 1981; Read and Horbury 1993; Saller et al. 1994, 1999; Wagner et al. 1995 1985; Renner 1988; Roehl and Weinbrandt 1985; Van Paleokarst and paleocaves: Cater and Gillcrist 1994; Golf-Racht 1996 Fritz et al. 1993; Hopkins 1999; Hurley et al. 1994; Dolomite reservoirs: Allan and Wiggins 1993; Davies Johnson 1994; Loucks 1999 (contains a list of 35 case 1979; Gosh et al. 1989; Loucks and Anderson 1985; studies ranging from the Ordovician to Early CretaLuo and Machel 1995; Purser 1985; Ruzyla and Friedman 1985; Saller and Henderson 1998; Schauer and ceous!); Loucks and Anderson 1985; Purdy and Waltham 1999; Saller et al. 1994; Tinker and Mruk Aigner 1997; Schmoker et al. 1985; Sun 1995; Weyl 1995; Wang and Al-Aasm 2002 1960; Zeidan 1994
Log Response
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mary porosity and early generated secondary anhydrite cementation. Reservoir potential may be restored by karstification and fracturing. (4) Non-evaporitic carbonate sequences associated with topographic high/unconformity, platform-margin buildups, or faults and fractures. Dolomites are the results of the reflux of slightly evaporated seawater (high microintercrystalline porosity resulting from early diagenetic dolomitization of lime-mud-rich sediments) or burial-derived dolomitizing fluids (leading to coarsely crystalline sucrosic dolomite; solution and intercrystalline porosity is common because of an excess of solution over precipitation during the replacement of limestone by dolomite). This category also includes hydrothermal dolomites that are the focus of current reservoir research. • Fractured reservoirs. Many carbonate reservoirs are hosts to extensive fracture networks characterized by sub-meter scale fracture zones that may penetrate vertically across lithologic and facies boundaries. Connected laterally, these fractures give rise to good reservoir permeability. Fracture and breccia porosity provide the possibility of extracting hydrocarbons for otherwise poor-quality reservoir rocks with low porosities (Nelson et al. 2000). Natural fracture patterns control producibility and permeability systems. Permeability is provided by joints and faults crossed by voids and microcracks (Fig. 17.4). Analysis of natural fractures must regard fracture cementation, size, fracture continuity, orientation and fracture density. The origin of fractures in carbonate reservoirs is related to folding, desiccation or unloading during subaerial exposure. Fracture intensity is higher in dolomites than in limestones and also appears to be related to the crystal size of the matrix. Fine-grained dolomites often are more intensively fractured than coarse-grained dolomites. Examples are the Ordovician dolomite reservoir in Michigan, the Late Cretaceous La Paz field in Venezuela, where strike-slip tectonics associated with faults and fracturing create the producing system, the fractured chalk of the North Sea basin, and the Oligocene Asmari limestone in Iran. • Hydrocarbon migration. Many carbonate reservoirs are influenced by diagenesis which is contemporaneous with hydrocarbon migration (e.g. Kharaib Formation, Middle East; Bab Asab, Sdahail Field, UAE). This is demonstrated by gradual deterioration of reservoir properties from crest to flank across facies boundaries (Grötsch et al. 1998). Diagenetic fluids also cause the formation of bedding-parallel dense diagenetic bodies with abundant stylolites at small-scale sedimentary cycle boundaries.
17. 1.4 Methods Hydrocarbon exploration uses seismic data, log response, the investigation of cores and cuttings and outcrop analog studies. A subsurface depositional model is constructed in three dimensions, using rock type elements, depositional sequence, geometry and bounding lithologies (Cant 1992). The recognition of sequences in seismic data leads to an understanding of facies distribution, which in turn leads to more accurate basin models, making the prediction of criteria required for hydrocarbon entrapment, such as reservoir, seal, source and migration pathway, more realistic and dependable. 17.1.4.1 Seismic Interpretation Interpretation of 3D seismic data allows analyzing the depositional setting of carbonate reservoirs. In case of shallow reservoirs and/or high-fold seismic data acquisition, depositional geometries and a low-resolution sequence stratigraphic analysis can be performed. Comparison of seismic geometries with analog examples is greatly supporting reservoir characterization. 17.1.4.2 Log Response Log responses are used to infer penetrated lithology, properties and the depositional character of rocks. Log character allows correlations and the construction of models showing the vertical trends in stratal thickness (fining or thickening upward) and geometry. Wireline logs provide the most powerful tool for predicting downhole lithology from physical properties. Petrophysical logs are measurements of various electrical, nuclear and acoustical properties, recorded as a function of depth. Commonly used logs are: • Sonic log. Measures the speed at which sound travels in rock units. It is related to both the porosity and lithology of the rock being measured. If the lithology is known, the log can be used in determining porosity. • Neutron log. Measures the porosity of a rock, indicating its response to the quantity of hydrogen present in the rock unit. • Gamma ray log. Records the radioactivity of a formation and shaliness, clay content and presence of organic matter. The most common logs used in carbonate reservoirs are the gamma-ray-neutron log and the compensated neutron-density log that distinguish limestone, dolo-
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Outcrop Analog Studies
mite and anhydrite, and reveal porosity (Longman 1981). The predominance of specific pore-types can be derived from the velocity-deviation log calculated by combining the sonic log with neutron-porosity or density logs (Anselmetti and Eberli 1999). Facies differentiation requires the integration of different logs. Possible relationships between log response and carbonate rock type are evaluated by cross-plots (Pickett 1977). This method can only be used if cores or cuttings from selected wells are available and log responses from control wells can be checked resulting in significant rock type clusters. A log which is not yet run in many wells is the Nuclear Magnetic Resonance (NMR) log. It is particularly suitable for carbonate reservoirs and allows analysis of porosity, permeability and pore type distribution at the same time. 17.1.4.3 Cores and Cuttings Lithofacies identification from logs is frequently difficult in shallow-water carbonate reservoirs and requires the investigation of cores and cuttings. Unlike log data which are plotted on a scale with meter thick marks, core data are plotted on centimeter-scales. Core analysis aims at differentiating lithofacies types and recognizing reservoir compartmentalization. These types are defined on the basis of lithological and textural criteria. Some authors group the lithofacies types according to their position within depositional models, differentiating e.g. a phylloid core facies, a mound-cap facies and a mound-flank facies. Cores are studied by visual inspection that may reveal fracture and faults as well as bedding structures, or in thin sections. The descriptions usually follow the Shell system (Swanson 1981) and should consider lithology, rock color, depositional and diagenetic fabrics, and composition and textures. Many of these criteria are exhibited by microfacies data (see Sect. 17.1.5.2). Rock classifications used in reservoir studies are those of Dunham (1962; Fig. 8.5), Embry and Klovan (1972; Sect. 8.3.2), and Choquette and Pray (1970; Fig. 7.5). Note that the classification of mud-supported rocks (mudstone, wackestone) and grain-supported rocks (packstone, grainstone) is not necessarily congruent with low and high porosities! Packstones should be differentiated into samples whose intergranular areas are completely filled with mud (exhibiting petrophysical data similar to wackestones or mudstones), and samples whose intergranular spaces are incompletely filled (similar in petrophysical data to grainstone), see Kerans et al. (1994).
Cuttings. Thin-section cuttings provide a low-cost source of subsurface data that can be used to generate high-resolution sequence stratigraphic frameworks for wells that lack cores or high-quality wireline logs (Coffey and Read 2002). Reservoir predictions in wells with limited core coverage can be undertaken by integrating semi-quantitative microfacies data from cuttings with core data and log analysis data. This allows us to understand the log response in terms of porosity and microfacies and to reconstruct the rock succession from uncored parts of the wells. The method is relevant if you have to interpret old or incomplete well data or if log responses need to be understood in terms of pores types and complex lithologies. This method can also be applied to identify exploration targets in areas based on cuttings and log data alone. Cuttings should be sufficiently large (>5 mm) in order to show compositional variability. An important requirement for interpreting cuttings is information on the cutting sampling interval and the rate-penetration. Microfacies techniques are applied to well cuttings after they have been calibrated against core data (Rodriguez Schelotto and Carozzi 1981) and after having considered vertical variations in qualitative and quantitative changes in lithotypes (Koch 1991). Changes are expressed by shifts in percentages of microfacies types. Mixed microfacies types and lithologies can result either from (1) actual lithologic variations or interbedding within the cutting sample interval, or (2) downhole mixing. Assessing these possibilities requires comparing the lithologic variations in cuttings with the variations known from cores or outcrops. 17.1.4.4 Reservoir-Related Outcrop Analog Studies Outcrop studies provide compositional, geometrical and dimensional data for carbonate bodies and basic data for 3D modeling (Eggendorf et al. 1999). Common methods used in reservoir-related outcrop studies include measuring stratigraphic sections, determining sequence stratigraphic boundaries, sampling and studying of thin sections with regard to microfacies, and petrophysical measurements conducted on core plugs drilled in the field or from hand specimens to determine porosity and permeability (Doherty et al. 2002). The use of outcrop data in comparing with subsurface data and modeling requires digitalization of outcrop data. Questions that are touched upon by outcrop analogs are: • Large-scale geometry. In which way does the seismic pattern reflect the actual geometries, e.g. of cyclic carbonates, reefs, drowned platforms or slopes (Schlager
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Outcrop Analog Study: Sulawesi
Plate 149 Reservoir-Related Outcrop Analog Study: Late Tertiary of Sulawesi, Indonesia Tertiary shallow-marine tropical carbonates are major hydrocarbon reservoirs in various parts of the world, including Egypt, Iran, Iraq, southeastern Asia (South China Sea, Indonesia, Philippines) as well as Central and South America. The depositional environments of these predominantly Oligocene to Miocene deposits include platforms and ramps, shelf edge and platform reefs, and pinnacle reefs surrounded by deeper-water facies. Many Miocene carbonates, forming stratigraphic traps, exhibit well-preserved primary and secondary porosities. Freshwater leaching of aragonitic bioclasts during stages of emersion, dolomitization and fracturing of partly lithified sediment (enhancing permeability) are common controls on the reservoir rock qualities in the Philippines and various parts of Indonesia (Epting 1980; Longman 1985; Ascaria et al. 1997; Grötsch and Mercadier 1999; Wilson et al. 2000). Important reefbuilders are corals, red algae, bryozoans and encrusting foraminifera. Larger benthic foraminifera, miliolid foraminifera, green algae (e.g. Halimeda) and gastropods prevailed in lagoonal and ramp as well as in peri-reef environments. Microfacies, facies-controlled diagenesis and the distributional patterns of Larger foraminifera and algae offer a possibility for differentiating the various parts of reef complexes in cores and outcrops (Hallock and Glenn 1985; Sect. 14.2, Fig. 14.7). The plate shows the preserved primary and secondary porosity types and biota common in Miocene reef limestones. The impregnation of the samples with blue-dye stained resins reveals open pores. Open biomoldic pores, linked by fracture pores, may form some permeability. Note that particles originally consisting of aragonite (e.g. Halimeda and gastropods) exhibit stronger dissolution and moldic porosity (–> 3, 5) than primary calcitic constitutents (e.g. Larger foraminifera, –> 8). Although visible porosity appears high, effective porosity is low, because most pores are disconnected. The samples are from the Selayar Limestone of the Miocene Walanae Formation in the southern tip of South Sulawesi, Indonesia. Common facies types of the reefal deposits studied are coral limestones (–> 1, 2), foraminiferal limestones (–> 7, 8) and Halimeda limestones (–> 3, 4). Because of volcanism active during the formation of the carbonates, thickness and width of these limestones are limited to a few tens of meters (Personal communication A. Imram, Makassar, 2000). 1 Coral limestone. Corals (C) are encrusted by coralline red algae (RA) and attached foraminifera (F). Matrix is a bioclastic wackestone with algal debris. Porosity: Open macroborings (MB) and solution vugs (SV). 2 The reef framework consists of corals (C) and coralline red algae (RA). In addition, encrusting bryozoans (B) and gastropods (G) occur. 3 Halimeda floatstone. Note the strong selective dissolution of the primary aragonitic algal blades, resulting in high biomoldic porosity (BP). Chambers in foraminiferal (F) and gastropods shells (G) are filled with blocky meteoric cement. 4 Halimeda rudstone. In contrast to –> 3, algal blades are well preserved. Interparticle pores are partly occluded by several generations of submarine cement, consisting of isopachous rim cement (arrows) followed by white fibrous calcite cement. Similar to –> 1 and 2, there is only poor intraparticle porosity within the coral (C) due to the protection by red algal crusts (RA) and rapid submarine cementation. H: Halimeda. 5 Gastropod shell exhibiting intraskeletal and moldic porosity (MP) due to leaching. Note the still preserved cross-lamellar shell structure (arrow). Dissolution vug at right. 6 Oncoid (acervulinid-corallinacean macroid) composed of a coral (C), overgrown by coralline red algae (RA), competing with encrusting foraminifera (Solenomeris, SO). Solenomeris is a reef-building acervulinid foraminifer that formed monospecific reefs in the Early Tertiary, particularly in the Eocene (Plaziat and Perrin 1992). The genus, however, is also known from Miocene non-reef environments contributing to the formation of oncoids. Open porosity is restricted to few dissolution vugs (top right). 7 Foraminiferal grainstone/rudstone with distinct solution voids (SV) that cross the depositional fabric. Larger foraminifera include Amphistegina (A), Myogypsina (MY) and Heterostegina (HET). In Indo-Pacific Miocene carbonates the association of these foraminifera is diagnostic of shallow high-energy conditions (see Fig. 14.7) and perireef and forereef environments. 8 Packed foraminiferal rudstone consisting predominantly of shells of Heterostegina. The primary calcitic foraminifera show no dissolution. Porosity is restricted to solution voids partly cutting the skeletal grains (arrows). Low energy, protected inner shelf facies. Heterostegina requires calm water in sheltered reef parts or in deeper low-energy environments, attached to hard substrates or plants. –> 1–8: Courtesy of A. Imram (Makassar, Sulawesi, Indonesia).
Plate 149: Outcrop Related Study: Tertiary of Sulawesi
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et al. 1991; Campbell and Stafleu 1992; Eisenberg et al. 1994; Kerans et al. 1994; Antonelli and Mollema 2000)? Geometries are derived from mapping, detailed facies analysis and seismic outcrop modeling. Smaller carbonate bodies are evaluated by ground-penetrating radar methods. • Size of carbonate bodies and flow units. At what scale do potential flow units vary? What thickness and lateral extension can be expected for carbonate bodies formed by different processes (e.g. allochthonous and autochthonous oolitic sand shoals, platform-margin reefs, pinnacle reefs and mud mounds)? • Predictive interwell interpretation (Barnaby 1995; Pomar and Ward 1999). What facies and petrophysical characteristics can we expect in interwell areas? Answering these questions requires outcrop studies that consider common well spacing (several hundreds of meters). • Correlation. What stratigraphic units can be used in regional subsurface correlation? • Heterogeneity. Which lateral and vertical variability of petrophysical parameters can be expected, e.g. in reefs formed in different settings, or in different parts of platforms and ramps? At what scale do depositional and/or diagenetic heterogeneities occur?
17.1.5 Microfacies, Lithofacies and Reservoir Rock Types Modeling carbonate reservoirs and bridging the gap from seismic to subseismic heterogeneities requires differentiating rock types that reflect depositional and diagenetic constraints of petrophysical properties. 17.1.5.1 Reservoir Heterogeneity Problems encountered in carbonate reservoirs are due to the inherent heterogeneity and complexity of the system, caused by highly variable in-situ sedimentation and sediment transport as well as subsequent alteration by diagenetic processes. Carbonate reservoirs are characterized by the extreme heterogeneity of porosity and permeability, particularly in cyclic shelf carbonates, as a result of a complex depositional and diagenetic history. The variability of well-log responses and downhole petrophysical parameters and the lateral variability of reservoir quality is related to carbonate textures and pore-system dimensions (Tanguy and Friedman 2001). Different approaches are used to understand reservoir heterogeneities, including reservoir rock type modeling (Sect. 17.1.5.3) and grouping of
Reservoirs and Microfacies
facies data by multivariate methods (Tucker et al. 1998). Understanding heterogeneity requires integrated analysis of cores, cuttings, petrophysical data, logs and seismic data supplemented by reservoir-related outcrop studies (Smith et al. 2003). Outcrops show the common facies types, vertical and lateral facies variations, and offer the possibility for analysis of high-resolution sequence stratigraphy. However, they may have a very different diagenetic history compared to subsurface reservoirs. Cores show sedimentary and diagenetic structures, rock types and their composition, visible porosity, and oil staining. Thin section and microfacies data show depositional fabrics and the frequency and distribution of grain types and matrix responsible for primary porosity, allowing visible porosity to be estimated and porosity reduction by enhanced diagenesis to be evaluated.
17.1.5.2 Relevant Microfacies Data The first part of this book deals with many criteria that are responsible for heterogeneities of carbonate reservoir rocks (Chap. 4, 5 and 7). The following text summarizes these criteria. • Fine-grained matrix Different matrix types (micrite, Sect. 4.1.1; microspar, Sect. 4.1.3; calcisiltite, Sect. 4.1.4) as well as auto- and allomicrites (Pl. 6, Pl. 7) can provide different intercrystalline microporosities. Matrix microporosity may be high (10–20%) due to subaerial exposure or retention of intercrystalline porosity during recrystallization of original lime mud (Moshier et al. 1988). Diagenetic processes, e.g. recrystallization, resulting in pervasive microporosity and leading to microspar formation, may outweigh depositional controls on porosity distribution in micritic reservoirs (Jordan and Abdullah 1988). Take a look at ‘homogeneous’ and ‘inhomogeneous’ micrites (Pl. 6/1) both in the matrix and as constituents of grains and examine SEM microphotographs (Pl. 7/3, 5, 6) for the geometry and connections of micropores! • Grains Grains form the primary depositional framework. Their arrangement, frequency and size determine primary porosity. Grains react differently to subaerial, marine and burial diagenesis, depending on their primary mineralogical composition. Skeletal grains. Higher porosities may occur in intervals with high proportions of bioclastic particles. The evaluation of moldic and dissolution porosity is assisted
Reservoirs and Microfacies
by the differentiation of aragonitic and calcitic bioclasts (Fig. 4.9) and the course of skeletal diagenesis (Sect. 4.2.1; Pl. 29/1; Pl. 30/1). Depositional frameworks and primary porosity are controlled by the abundance and distribution of calcareous algae (Sect. 10.2.1), Larger foraminifera (Sect. 10.2.2), benthic sessile organisms (Sect. 10.2.3) and various shells (Sect. 10.2.4). These fossils can form specific biofabrics that control porosity and permeability development (Sect. 5.1.2). Smallscale heterogeneity can be related to different microstructural types of skeletal grains, e.g. rudists (Pl. 88; Russell et al. 2002). Peloids and cortoids. These grains (Sects. 4.2.2 and 4.2.3) are susceptible to microborings and selective dolomitization. Both processes contribute to microporosity. Oncoids and rhodoids. Oncoids (Sect. 4.2.4.1) are common in lagoonal reservoir rocks, particularly in microbial carbonates. Rhodoids (Sect. 4.2.4.2) are major constituents of Tertiary reservoir rocks. Morphological and systematic differentiation enables paleodepths and subenvironments to be recognized. Ooids. Depending on their primary mineralogy and the composition of their nuclei and cortices, ooids (Sect. 4.2.5) are susceptible to complete or partial leaching, resulting in oomolds (Pl. 13/7). Micritic ooids contribute to microporosity and are susceptible to dolomitization (Pl. 13/8). Many oolitic reservoirs are controlled by paleoenvironmental conditions producing shoal- or sheetlike geometries (Fig. 4.26) and deposition and/or mixing of aragonitic and calcitic grains. Pisoids. Pisolitic limestones (Sect. 4.2.6) can act as reservoirs if the interparticle porosity is enhanced by vadose dissolution of interparticle cements. Aggregate grains (Sect. 4.2.7) associated with ooids and peloids are common constituents of lagoonal and open-platform reservoirs. Lithoclasts. Lithified rock particles contribute to the formation of debris and talus accumulated in proximal or distal positions from the source area. A thin-section based differentiation of clast composition in cores is necessary for distinguishing forereef talus from lithoclasts of debris flows or turbidites. • Fabrics Sedimentary and diagenetic fabrics enhance and/or reduce porosity and permeability. Lamination and bedding (Sect. 5.1.3). Lamination caused by finely distributed clay may reduce permeability. Interbedding of millimeter- to centimeter-thick layers consisting of wackestone or mudstone with grainstone layers can result in different porosity values. Fenestral fabrics (Sect. 5.1.5). Interconnected
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birdseyes may facilitate fluid and gas migration. Keystone vugs (Pl. 29/9) may be responsible for high porosities. Burrowing and bioturbation (Sect. 5.1.4). Interconnected burrows forming networks may have a distinct influence on cementation, permeability and porosity. Nodular fabrics (Sect. 5.1.6). Nodular limestones can exhibit strongly varying porosity values due to the inhomogeneous composition of the rock. Variations depend on the proportion of nodules and matrix, and the mode of origin of the nodules. Fissures (Sect. 5.3.1). Depositional and tectonic fissures related to block-faulting can act as important migration paths for hydrocarbons. The history of fissure networks can be deciphered from the relations of internal sediment, cements and multiple rupturing. Fractures and veins (Sect. 5.3.2). Microfractures associated with shear zones, extensional movements or hydraulic fracturing can be used to predict macrofracture patterns acting as hydrocarbon migration paths and reservoirs. Breccias (Sect. 5.3.3). Fractured talus breccias as well as solution breccias may yield high porosities, depending on the mode of breccia origin. • Pores Evaluation of visible porosity in thin sections is commonly based on thin sections impregnated with blue-dye resin (Yanguas and Dravis 1985) and must consider the following questions: Abundance and distribution of open and closed pores. Are open pores restricted to specific areas? Are the pores occluded by early or late diagenetic cement? Frequency of fabricselective and non-fabric selective pore types using the terminology and classification by Choquette and Pray (1970; Fig. 7.5; Box 7.4). Which pore type dominates? Are primary pores, e.g. interparticle pores, enlarged by solution grading into vugs and enhancing effective porosity (Pl. 29/9)? Are intraskeletal pores of importance (see Bachman 1984)? Are moldic pores restricted to specific grain types (e.g. oomoldic, biomoldic)? Are vugs connected or disconnected (McNamara et al. 1992)? Are pores collapsed (Mowar et al. 1998)? The 3D pore system and dynamic properties of the pore network are measured by special core analysis as input for dynamic modeling. The pore network forms the link between static characterization of the reservoir and its dynamic behavior. • Cement Cementation patterns must be described in regard to cement mineralogy (carbonate, anhydrite, silica etc.), cement types (Sect. 7.4.2.1), cement fabrics (Sect.
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7.4.2.2), and cement sequences (e.g. Pl. 35). Cement sequences are of great importance in evaluating the temporal succession of marine, freshwater and burial diagenesis. • Diagenetic textures Mechanical compaction (Sect. 7.5.1) alters the spatial arrangement of grains and leads to an increase in grain packing paralleled by a decrease in interparticle porosity. For compaction criteria see Box 7.9. Chemical compaction (Sect. 7.5.2) is responsible for pressure solution and stylolites that reduce porosity, but also may act as conduits for fluids (Box 7.10). Fig. 7.18 assists in describing pressure solution features. Of specific interest in evaluating the role of pressure solution for the development of permeability and porosity are the geometry, distribution and abundance of stylolites, and the association of fracture sets with stylolites (Nelson 1981; Grötsch et al. 1998). • Dolomitization Dolomitization can enhance or reduce porosity depending on the mode and the timing of the dolomitization process (Mazzullo 1992). Dolomite textures should be described using the criteria summarized in Sect. 7.8.1.1, including crystal size and crystal textures. Dolomitization may be restricted to the matrix or to specific grain types or may have effected the total rock. Differentiation of dolomitization stages requires ongoing investigations, including cathodoluminescence and isotope studies. Dedolomitization (Sect. 7.8.3), often associated with meteoric diagenesis, may improve intercrystalline and moldic porosity.
17.1.5.3 Reservoir Rock Types and Facies Criteria Lucia (1995, 1999) underlined the importance of pore space related to rock fabrics, which control porosity, permeability, saturation and dynamic behavior of fluid movement through the rock. He introduced the concept of rock/petrophysical classes that aims at a classification of carbonate reservoir rocks based on petrophysical properties and major pore space groups (interparticle and vuggy pore spaces). The classes should reflect unique reservoir characteristics and specific distribution patterns. Lucia’s classification is based on the fact that pore size distribution controls permeability and saturation and that pore-size distribution is related to sedimentary fabric, resulting in different pore types in grain- and mud-supported limestones.
Reservoir Rock Types
Many authors consider depositional fabric and its diagenetic overprinting to be the major factors explaining differences in petrophysical properties. However, complex diagenetic overprinting can prevent depositional facies from being related to petrophysical properties. Recognizing these difficulties and aiming for rock units that can be used in static and dynamic 3D models evaluating reservoir quality, Grötsch and Mercadier (1999) and Grötsch et al. (2003) proposed the concept of reservoir rock types (RRTs). Each reservoir rock type is characterized by its pore geometry, a specific average porosity and average permeability, an average gas saturation, and a saturationheight function. A specific lithofacies may be a definition criteria but not necessarily. RRTs do not always follow the classical texture-based carbonate classification schemes, but instead combine effects of depositional texture and diagenetic overprinting into classes of common petrophysical properties. The rock facies is defined by mercury injection capillary pressure analysis and pore-throat distribution curves as well as other SCAL measurements. The same depositional facies (lithofacies) may exhibit different reservoir properties due to a different degree of diagenetic overprinting. Correlation between depositional fabric and petrophysical data may be good in carbonates that have been affected only by early diagenetic processes, but is usually weaker in carbonates affected by burial diagenesis. Distinguishing RRTs requires integration of depositional lithofacies, including their environmental controls, and diagenetic sequences determined in cores and thin section, porosity and permeability from core plug measurements, high-pressure mercury measurements exhibiting the pore-throat distribution, and the evaluation of wireline log characteristics and formation tests. Microfacies and reservoir rock types. Some of the problems involved in correlating microfacies and petrophysical patterns are caused by the typification of microfacies. Microfacies types based solely on textural classification (e.g. bioclastic wackestone, oolitic grainstone) or defined primarily to the association (e.g. intraclastic skeletal wackestone) or proportion of different grain types (e.g. ooid-rich and ooid-poor grainstone) do not reflect all the constituents of carbonate rocks influencing porosity and permeability. The microfacies types discussed in Chap. 11 of this book are delimited by differences in qualitative and quantitative compositional criteria. These often rather small differences reflect variations within the depositional environment and allow paleoenvironment conditions and depositional settings to be evaluated (Chap.
Basics: Carbonate Hydrocarbon Reservoirs
12; Chap. 14). Because these microfacies types do not consider the diagenetic overprint, a correlation between the types and petrophysical properties is difficult and often not possible (Kostic and Aigner 2004). To relate physical properties and microfacies, major constituents (grains, matrix), their mineralogy, texture and depositional fabric (bedding, layering, burrowing), the visible pore space (pore type, size, distribution and frequency), and diagenetic features must be considered. The latter are of prime importance, particularly indications of dissolution, pressure-solution (stylolites) and fracturing. If these data are included in the characterization of microfacies types, the values and anisotropy of petrophysical data are fairly well related to microfacies criteria visible in thin sections and cores, and their spatial distribution and preferred orientations (Dürrast and Siegesmund 1999; Pöppelreiter 2002). Currently, the main technological challenge in subsurface modeling of carbonate hydrocarbon reservoirs is the development of a petrophysically based classification scheme of carbonate rocks which honors the pore geometries and therefore the dynamic characteristics and which can be at the same time applied for static and dynamic high-resolution reservoir modeling. Basics: Carbonate hydrocarbon reservoirs Aigner, T., Heinz, J., Hornung, J., Aspirion, U. (1999): A hierarchical process approach to reservoir heterogeneity: examples from outcrop analogues. – Bulletin de la Centre des Recherche, Elf Exploration Production, 22, 1-11 Alsharhan, A.S., Scott, R.W. (eds., 2001): Middle East models of Jurassic/Cretaceous carbonate systems. – SEPM, Special Publication, 69, 365 pp. Asquith, G.B. (1985): Handbook of log evaluation techniques for carbonate reservoirs. – Methods in Exploration Ser., American Association of Petroleum Geologists, 5, 47 pp. Campbell, A.E., Stafleu, J. (1992): Seismic modeling of an early Jurassic, drowned carbonate platform: Djebel Bou Dahar, High Atlas, Morocco. – American Association of Petroleum Geologists, Bulletin, 76, 1760-1777 Cant, D.J. (1992): Subsurface facies analysis. – In: Walker, R.G., James, N.P. (eds.): Facies models. Response to sealevel change. – 27-45, Geological Association of Canada Chilingar, G.F., Mannon, R.W., Rieke, H.H. (eds., 1972): Oil and gas production from carbonate rocks. – 408 pp., Amsterdam (Elsevier) Chilingarian, G.F., Mazzullo, S.J., Rieke, J.H. (eds., 1992): Carbonate reservoir characterization: a geologic-engineering analysis. Part I. – Developments in Petroleum Science, 30, 639 pp., Amsterdam (Elsevier) Chilingarian, G.F., Mazzullo, S.J., Rieke, J.H. (eds., 1996): Carbonate reservoir characterization: a geologic-engineering analysis. Part II – 1010 pp., Amsterdam (Elsevier) Coffey, B.P., Read, J.F. (2002): High-resolution sequence stratigraphy in Tertiary carbonate-rich sections by thinsectioning well cuttings. – American Association of Petroleum Geologists, Bulletin, 86, 1407-1415 Dmietrievsky, A.N., Kuznetsov, V.G. (1992): Climatic con-
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trol on Carboniferous-Permian Carbonate reservoirs from the Southeast Russian Platform, USSR. – In: Spencer, A.M. (ed.): Generation, accumulation and production of Europe’s hydrocarbons II. – Special Publication, European Association of Petroleum Geoscientists, 2, 225-229 Dürrast, H., Siegesmund, S. (1999): Correlation between rock fabrics and physical properties of carbonate reservoirs, – International Journal of Earth Sciences, 88, 392-408 Elf-Aquitaine (1982): Exploration for carbonate petroleum reservoirs. – 213 pp., New York (Wiley) Ehrlich, R., Crabtree, S.J., Horkowitz, K.O., Horkowitz, J.P. (1991): Petrography and reservoir physics, I: objective classification of reservoir porosity. – American Association of Petroleum Geologists, Bulletin, 75, 1547-1562 Fritz, R.D., Wilson, J.L., Yurewicz, D.A. (eds., 1993): Paleokarst related hydrocarbon reservoirs. – Society of Economic Paleontologists and Mineralogists, Core Workshop, 275 pp. Greenlee, S.M., Lehmann, P.L. (1993): Stratigraphic framework of productive carbonate buildups. – In: Loucks, R.G., Sarg, J.F. (eds.): Carbonate sequence stratigraphy: recent developments and applications. – American Association of Petroleum Geologists, Memoir, 57, 43-62 Grötsch, J., Mercadier, C. (1999): Integrated 3D reservoir modeling based on 3D seismic: The Tertiary Malampaya and Camago buildups, offshore Palawan, Philippines. – American Association of Petroleum Geologists, Bulletin, 83, 1703-1728 Grötsch, J., Suwaina, O., Ajlani, G., Taher, A., El-Kassawneh, R., Lockier, S., Coy, G., Van der Werd, E., Masalmeh, S., Van Dorp, J. (2003): The Arab Formation in central Abu Dhabi: 3D reservoir architecture and static and dynamic modeling. – GeoArabia, 8, 47-86 Halley, R.B., Loucks, R.G. (eds., 1980): Carbonate rock reservoirs: notes. – Society of Economic Paleontologists and Mineralogists, Core Workshop, 1, 186 pp. Harris, P.M., Kowalik, W.S. (1994): Satellite images of carbonate depositional settings. Examples of reservoir- and exploration scale geologic facies variation. – AAPG Methods in Exploration Series, 11, 147 pp. Homewood, P., Mauriaud, P., Lafont, F. (2000): Best practices in sequence stratigraphy for explorationists and reservoir engineers. – Mémoir de la Centre des Recherche, Elf Exploration Production, 25, 81 pp. Hughes, G.W.G., Siddiqui, S., Sadler, R.K. (1993): Shu’aiba rudist taphonomy using computer tomography and image logs, Shaybah field, Saudi Arabia. – GeoArabia, 8, 585-596 Jardine, D., Wilshart, J.W. (1987): Carbonate reservoir description. – Society of Economic Paleontologists and Mineralogists, Special Publication, 40, 129-152 Jones, R.W. (1996): Micropaleontology in petroleum exploration. – 432 pp., Oxford (Clarendon Press) Katz, B.J. (1990): Lacustrine basin exploration. – American Association of Petroleum Geologists, Memoir, 50, 340 pp. Kerans, C., Lucia, F.J., Senger, R.K. (1994): Integrated characterization of carbonate ramp reservoirs using Permian San Andres Formation outcrop analogs. – American Association of Petroleum Geologists, Bulletin, 78, 181-216 Kiessling, W., Flügel, E., Golonka, J. (1999): Paleoreef maps: Evaluation of a comprehensive database of Phanerozoic reefs. – American Association of Petroleum Geologists, Bulletin, 83, 1552-1587 Koch, R. (1991): Faziesanalyse aus Spülproben. – Zentralblatt für Geologie und Paläontologie, Teil I, 1990, 1029-1043
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Lidz, B.H. (ed., 1987): Reservoir sedimentology. – Society of Economic Paleontologists and Mineralogists, Special Publication, 40, 357 pp. Longman, M.W. (1981): Carbonate diagenesis as a control on stratigraphic traps (with examples from the Williston Basin). - American Association of Petroleum Geologists, Education Course Note Series, 21, 159 pp. Lucia, F.J. (1983): Petrophysical parameters estimated from visual descriptions of carbonate rocks: a field classification of carbonate pore space. – Journal of Petroleum Technology, 35, 629-637 Lucia, F.J. (1995): Rock-fabric/petrophysical classification of carbonate pore space for reservoir characterization. – American Association of Petroleum Geologists, Bulletin, 79, 1275-1300 Lucia, F.J. (1999): Carbonate reservoir characterization. – 226 pp., Berlin (Springer) Maliva, R.G., Dickson, J.A.D. (1992): Microfacies and diagenetic controls of porosity in Cretaceous/Tertiary chalks, Eldfisk Field, Norwegian North Sea. – American Association of Petroleum Geologists, Bulletin, 76, 1825-1838 Mazzullo, S.J. (ed., 1980): Stratigraphic traps in carbonate rocks. – American Association of Petroleum Geologists, Reprint Series, 23, 217 pp. Mazzullo, S.J. (1992): Geochemical and neomorphic alteration of dolomite: a review. – Carbonates and Evaporites, 7, 21-37 Moore, C.H. (2001): Carbonate reservoirs. Porosity evolution and diagenesis in a sequence stratigraphic framework. – Developments in Sedimentology, 55, 460 pp., Amsterdam (Elsevier) Pickett, G.R. (1977): Recognition of environment and carbonate rock type identification. – Formation Evaluation Manual Unit II, section Exploration Well, 4-25, Oil and Gas Consultants International Inc. Pöppelreiter, M. (2002): Facies, cyclicities and reservoir properties of the Lower Muschelkalk (Middle Triassic) in the NE Netherlands. – Facies, 46, 119-132 Purdy, E.G., Waltham, D. (1999): Reservoir implications of modern karst topography. – American Association of Petroleum Geology, Bulletin, 83, 1774-1794 Reeckman, A., Friedman, G.M. (1982): Exploration for carbonate petroleum reservoirs. – 213 pp., New York (Wiley) Roehl, P.O., Choquette, P.W. (eds., 1985): Carbonate petroleum reservoirs. – 622 pp., New York (Springer) Schlager, W., Biddle, K.T., Stafleu, J. (1991): Picco di Vallandro (Dürrenstein) - a platform-basin transition in outcrop and seismic model. – Dolomieu Conference on Carbonate Platforms and Dolomitization, Guidebook, Excursion E, 22 pp. Scholle, P.A. (1977): Deposition, diagenesis and hydrocarbon potential of ‘deeper-water’ limestones. – American Association of Petroleum Geologists, Continuing Education Course Note Series, 7, 25 pp. Vahrenkamp, V., Grötsch, J. (1994): 3D architecture model of a Lower Cretaceous carbonate reservoir: Al Huwaisah field, Sultanate of Oman. – Geo’94, The Middle East Petroleum Geosciences, vol. II, 901-913 Zempolich, W.G., Cook, H.E. (eds., 2003): Paleozoic carbonates of the Commonwealth of Independent States (CIS): subsurface reservoirs and outcrops analogues. – SEPM, Special Publication, 74 Further reading: K073, K074, K075, K076 (porosity), K160 (log evaluation), K200 (carbonate hydrocarbon reservoirs)
Ore Deposits in Carbonate Settings
17.2 Carbonate-Hosted Mineral Deposits Carbonate rocks are host sediments for various ore deposits occurring within sedimentary rocks, particularly for lead-zinc, copper, manganese and iron mineralizations. Some of these mineralizations have been already discussed including iron ooid deposits (Sect. 4.2.5), Pb-Zn-Fe mineralizations occurring at subaerial exposure surfaces of carbonate platforms (Sect. 5.2.3), mineralizations within carbonate breccias (Sect. 5.3.3.5) and those associated with stromatolites and microbial carbonates (Sect. 9.1.5.2). The following text concentrates on the controls of carbonate depositional facies on mineral deposits and underlines the potential of microfacies in the study of sedimentary carbonate-hosted mineral deposits. These deposits are strata-bound (ores that are bound to specific lithologies or specific stratigraphic units without a textural correlation with the host rock), stratiform (ore bodies showing textural correlation with the host rock), or sedimentary (ores occurring within the sediment, e.g. in cavities and exhibiting geopetal structures). Age relationships between host rocks and mineralization may show that the ore is younger (epigenetic) or synchronous (syngenetic respectively syndepositional/diagenetic) as compared with the host rock.
17.2.1 Ore Deposits and Carbonate Settings Settings. Syn- and epigenetic mineralizations in carbonates occur preferentially in: • Paleokarst settings related to emersion phases and recorded by major unconformities (e.g. Assereto et al. 1976; Boni et al. 1981; Rhodes et al. 1984; Sangster 1988; Glazek 1992; Sass-Gustkiewicz 1999). • Platform carbonates. Syn- and epigenetic mineralizations in platform carbonates are known from various parts of the world (e.g. Spalletta et al. 1981; De Voto 1988). Syngenetic mineralizations (usually PbZn and Fe-Mn) took place in tidal and subtidal carbonates. The specific paleogeographic positions of the platforms, the input of connate waters and brines and decay of organic matter are discussed as major controls on metal accumulations in platform carbonates (e.g. Fuchs 1981; Parnell et al. 1990). The source of fluids containing metals in solution is explained as being bottom waters of saline or dysoxic and anoxic intra- or periplatform basins floating as metal-enriched brines into adjacent platform carbonates and porous reef limestones (e.g. Fan et al. 1999). Another possibility is the remobilization and migration of older mineralizations.
Microfacies and Ore Deposits
• Reefs (Monseur 1973; Etminan et al. 1984; Lagny 1984) are common host rocks, because of their high primary porosity and the possibility that ongoing cementation forms a seal. • Hardgrounds associated with condensation horizons in basinal settings (e.g. Mn mineralizations; Krainer et al. 1994). • Slopes of shelf-basin transitions provide settings with reduced sedimentation favoring e.g. manganese mineralizations (Maynard 1990). Pb-Zn deposits in carbonate rocks. Carbonatehosted Pb-Zn ores known from Cambrian to Cretaceous limestones and dolomites are excellent examples of synand epigenetic carbonate-hosted mineral deposits. Facies controls are especially evident for MississippiValley Type and Bleiberg Type deposits (Fontboté and Amstutz 1983; Buelter and Guillemette 1988; Mazzullo and Gregg 1989; Cerny 1989; Bechstädt and Boni 1994; Zeeh et al. 1995; Dejonghe 1998; Sass-Gustkiewicz and Dzulynski.1999; Zeeh et al. 1999; Liaghat et al. 2000). The epigenetic (diagenetic) Mississippi-Valley Type ores occur near platform margins or in epeiric carbonates. (e.g. Missouri, North West Canada, Poland, Sardinia, North Africa). The input of metal solutions took place via saline connate waters from adjacent basins. The precipitation of metals occurred in solution breccias and paleokarst structures and in reef zones. The distribution of the syngenetic Bleiberg Type ores is distinctly controlled by facies and regional tectonics. Strata-bound and stratiform Pb-Zn deposits occur in different paleogeographic positions: (a) at the base of transgressive surfaces above the basement near basin margins (e.g. Mississippi-Valley Type deposits and deposits in Europe), and (b) within thick carbonate sequences far from transgressive horizons (e.g. Bleiberg or Spain). Both types correspond to shallow-marine sediments, the former with partly high clastic input. The joint criteria of both types are mineralizations of tidal and lagoonal sediments, evaporitic environments, and abundance of dolomitization (Cerny 1989). Many mineralizations are associated with bedded platform interior limestones exhibiting cyclic sedimentation patterns (Bechstädt 1975) and indications of subaerial exposure, and with reef limestones. Most reefs are shallow-water platform reef mounds and platformmargin reefs, both of which are prone to sea-level fluctuations causing karstification and solution and cavern porosities. In contrast, Pb-Zn deposits within mud mounds formed in deeper outer ramp positions are often fault-controlled (e.g. Early Carboniferous mounds in Ireland). The mineralization postdates mud mound formation. The precipitation of Pb-Zn sulfides in reefs
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may have been accelerated by the availability of abundant organic matter and microbes as postulated for the Middle Devonian Pine Point deposits. The lithofacies of many Pb-Zn host rocks is characterized by the predominance of carbonates formed in very shallow depths, tidal and lagoonal sediments, the abundance of organic matter in limestones and dolomites indicating reducing diagenetic conditions, abundant microbial textures, and evaporitic minerals. The multi-phase diagenesis of the Pb-Zn ores often starts within an open system where precipitation can take place due to sulfatereducing bacteria or triggered by seepage and evaporative pumping mechanisms. The second phase characterized by diagenesis of both carbonate grains and ore minerals occurs within a closed system and is connected with the formation of diagenetic crystallization rhythmites and syngenetic breccias.
17.2.2 Microfacies and Ore Deposits Microfacies studies of host rocks and mineral deposits are relevant in evaluating the depositional environment of the sediments affected by mineralizations, understanding the spatial and temporal relationships between host rocks and ore bodies, and answering metallogenetic questions. Ore minerals must be studied in polarized reflected light of polished sections. Reconstruction of sedimentary environments. Microfacies data described in Chap. 4 and Chap. 5 allow paleoenvironmental controls (Chap. 12) to be identified and syngenetic mineralizations to specific facies zones and facies models (Chap. 14) attributed. These data assist in evaluating the overall depositional framework affected by mineralizations. Host rock/mineralization relationships. Answering this question requires a thorough thin-section analysis of carbonate beds below and above the mineralized horizon (e.g. Hagenguth 1984) and the differentiation of microfacies types (e.g. Cisternas 1986). Metallogenetic problems. Sedimentary fabrics of carbonate rocks (Chap. 5) and fabrics observed in syngenetic ores exhibit many coincidences. These coincidences are used to prove the syngenetic origin of stratiform deposits. Relevant sedimentary ore fabrics have been summarized by Schulz (1976): • Stratified conformable arrangement of ore bodies within bedded carbonate sediments. • Parallel bedding of mineralized strata and carbonate strata.
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• Nonparallel oblique bedding that might indicate cross-bedding structures. • Gravity-dependent mechanical deposition of grains or ore detritus. • Vertically graded bedding. • Geopetal deposition of minerals within sedimentary cavities or in intraskeletal pores, e.g. shells filled with micrite at the base and ore at the top. • Repeated sequences exhibiting an increase and decrease in the abundance of ore minerals within individual strata forming rhythmic graded beds. • Resediments that provide evidence of erosion of orecontaining strata. • Mechanical deposition of mineralized grains within an erosional bottom relief. • Syngenetic deformation of both ore bodies and host carbonates due to sliding and slumping. Basics: Carbonate-hosted mineral deposits Bechstädt, T., Boni M. (eds., 1994): Sedimentological, stratigraphical and ore deposits field guide of the autochthonous Cambro-Ordovician of southwestern Sardinia. – Memorie descrittive, Carta geologica d’Italia, 48, 432 pp. Bechstädt, T., Dohler-Hirner, B. (1983): Lead-Zinc deposits of Bleiberg-Kreuth. – In: Scholle, P.A., Bebout, F.G., Moore, C.H. (eds.): Carbonate depositional environments. – American Association of Petroleum Geologists, Memoir, 1, 55-63 Cerny, I. (1989): Die karbonatgebundenen Blei-Zink-Lagerstätten des alpinen und außeralpinen Mesozoikums. Die Bedeutung ihrer Geologie, Stratigraphie und Faziesgebundenheit für Prospektion und Bewertung. – Archiv für Lagerstättenforschung der Geologischen Bundesanstalt Wien, 11, 5-125 Cisternas, M.E. (1986): Stratigraphische, fazielle und lithogeochemische Untersuchungen in der Unterkreide der Region Atacama: Metallogenetische Bedeutung am Beispiel der schichtgebundenen Eisen-Lagerstätte Bandurrias. – Heidelberger Geowissenschaftliche Abhandlungen, 2, 268 pp. De Voto, R.H. (1988): Late Mississippian paleokarst and related mineral deposits, Leadville Formation, Central Colorado. – In: James, N.P., Choquette, P.W. (eds.): Paleokarst. – 278-305, Berlin (Springer) Diaz, L.L. (1986): Strata-bound Zn-, Ba-, Ag-ore deposits of the Lower Cretaceous in the Atacama region, Northern Chile. – Heidelberger Geowissenschaftliche Abhandlungen, 3, 186 pp. Fan, D., Hein, J.R., Jye, J. (1999): Ordovician reef-hosted Jiaodingshan Mn-Co deposit and Dawashan Mn deposit, Sichuan Province, China. – Ore Geology Reviews, 15, 135-151 Ferguson, J., Burne, R.V., Chambers, L.A. (1983): Iron mineralization of peritidal carbonate sediments by continental groundwaters, Fisherman Bay, South Australia. – Sedimentary Geology, 34, 41-57 Flügel, E. (1989): Typen und wirtschaftliche Bedeutung von Riffkalken. – Archiv für Lagerstättenforschung der Geologischen Bundesanstalt Wien, 10, 25-32 Friedrich, G.H., Herzig, P.M. (eds., 1988): Base metal sul-
Basics: Carbonate-Hosted Mineral Deposits
fide deposits. – 290 pp., Berlin (Springer) Hagenguth, G. (1984): Geochemische und fazielle Untersuchungen an den Maxerbänken im Pb-Zn-Bergbau von Bleiberg-Kreuth/Kärnten. – Mitteilungen der Gesellschaft der Geologie- und Bergbaustudenten in Österreich, Sonderheft, 1, 110 pp. Hentschel, B., Kern, M. (1992): Ein vererzter unterkarboner Paläokarst in den Zentralen Karnischen Alpen (Italien/ Österreich). - Jahrbuch der Geologischen Bundesanstalt Wien, 135, 225-232 Klemm, D.D., Schneider, H.-J. (eds., 1977): Time- and stratabound ore deposits. – 444 pp., Berlin (Springer) Krainer, K., Mostler, H., Haditsch, J.G. (1994): Jurassische Beckenbildung in den Nördlichen Kalkalpen bei Lofer (Salzburg) unter besonderer Berücksichtigung der Manganerz-Genese. – Geologische Bundesanstalt Wien, Abhandlungen, 50, 257-293 Lagny, P.H. (1984): Milieu récifal et minéralisation plombozincifères. – 3ème Cycle de la Terre, 32.1-32.16 Mazzullo, S.J., Gregg, J.M. (1989): Mississippi-Valley type mineralization and relationships to carbonate diagenesis: introductory remarks. – Carbonates and Evaporites, 4, 131135 Parnell, J., Ye, L., Chen, C. (eds., 1990): Sediment-hosted mineral deposits. – International Association of Sedimentologists, Special Publication, 11, 227 pp. Rhodes, D., Lantos, E.A., Lantos, J.A., Webb, R.J., Owens, D.C. (1984): Pine Point ore bodies and their relationship to the stratigraphy, structure, dolomitization, and karstification of the Middle Devonian Barrier Reef complex. – Economic Geology, 79, 991-1055 Ruffell, A.H., Moles, N.R., Parnell, J. (1998): Characterisation and prediction of sediment-hosted ore deposits using sequence stratigraphy. – Ore Geology Reviews, 12, 207223 Sangster, D.F. (1988): Breccia-hosted lead-zinc deposits in carbonate rocks. – In: James, N.P., Choquette, P.W. (eds.): Paleokarst. – 102-116, Berlin (Springer) Sangster, D.F., MacIntyre, D. (eds., 1983): Sediment-hosted stratiform lead-zinc deposits. – Mineralogical Association of Canada, Short Course Handbook, 8, 308 pp. Sass-Gustkiewicz, M., Dzulynski, S. (1999): On the origin of strata-bound Zn-Pb ores in Upper Silesia, Poland. – Annales Societatis Geologorum Poloniae, 68, 267-278 Saunders, C.M., Strong, D.F., Sangster, D.F. (1992): Carbonate-hosted lead-zinc deposits of western Newfoundland. – Geological Survey of Canada, Bulletin, 419, 78 pp. Sawkins, F.J. (1984): Metal deposits in relations to plate tectonics. – 325 pp., Berlin (Springer) Schneider, H.J. (1964): Facies differentiation and controlling factors for the depositional lead-zinc concentrations in the Ladinian geosyncline of the Eastern Alps. – In: Amstutz, G.C. (ed.): Sedimentology of ore deposits. – Developments in Sedimentology, 2, 29-45 Schneider, H.J. (ed., 1983): Mineral deposits of the Alps and the Alpine Epoch in Europe. – 402 pp., Berlin (Springer) Schulz, O. (1976): Typical and nontypical sedimentary ore fabrics. – In: Wolf, K.H. (ed.): Handbook of strata-bound and stratiform ore deposits, I. – vol. 3, 295-338, Amsterdam (Elsevier) Wolf, K.H. (ed., 1976-1986): Handbook of strata-bound and stratiform ore deposits. – 14 volumes, Elsevier (Amsterdam) Further reading: K201
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18 Carbonate Rock Resources, Facies, Weathering, Preservation
The industrial use of carbonate rocks depends on physical and chemical properties. These properties are the result of the sedimentary and diagenetic history of the rocks. In many cases, the ‘primary (depositional) facies’ influencing the course of diagenesis are important controls for specific physical-chemical rock criteria. Microfacies analysis assists in understanding these controls. The study of causal relationships between microfacies data and specific rock properties will be an important tool of future work. This chapter demonstrates some examples illuminating the potential of applied facies analysis in this field.
18.1 Industrial Use of Carbonate Rocks The industrial use of carbonate rocks is maxifold, including chemical and metallurgical industries, construction and concrete materials as well as agricultural uses (Boynton 1980; Leidner 1994). Limestones and dolomites have the third rank in the amount and value of globally extracted mineral resources. Classifications. The specific uses of carbonate rocks need specific classifications related to the chemical and mineralogical composition and petrophysical properties (porosity, pore size and pore throat distribution, permeability; see Sect. 7.5.3). Industrial use requires categories describing compositional intergradations caused by differences in the relative amounts of calcite and dolomite and by various amounts of non-carbonate minerals (e.g. Bissell and Chilingar 1967; Correns 1968; Bentz and Martini 1968; Flügel and Haditsch 1975; Langbein et al. 1982; Lobitzer and Surenian 1984). These categories are determined by national and international norms (e.g. Gotthart and Kasig 1996; Oates 1998). Differences in chemical composition can be expressed on the basis of Ca/Mg weight ratios or using percentages of calcite, dolomite and clay. Different limits are used in order to define subgroups within
the calcite/dolomite/non-carbonate triangle. Subdividing carbonate rocks according to their chemical composition (e.g. CaO, MgO, SiO2, and P2O5) is essential for the economic use of limestones and dolostones. Textural classifications (Chap. 8) are of value in characterizing petrophysical rock properties related to fabric and grain/matrix ratios (e.g. weathering resistance of building stones). The incorporation of additional criteria (rock color, grain size, sorting, porosity) in the Dunham classification is recommended for distinguishing carbonate rocks of different physical and chemical qualities. Dearman (1981) proposed an engineering classification of carbonate rocks based on a combination of textural and compositional criteria with material and mass properties. Understanding geotechnical properties (e.g. compressive strength) requires classifications that combine petrography, textural and diagenetic features and technological data (Rashed and Sediek 1997; Bell et al. 1999).
18.2 Exploration and Exploitation of Carbonate Rocks Economic exploitation of carbonate rocks for industrial use should consider • relations between depositional facies, diagenetic history and rock properties, • differentiation of facies types related to differences in physical and chemical criteria, and • transformation of these facies types into limestone and dolomite categories that are of interest to companies and clients. Practical steps involved are: • Field work. Facies mapping and sampling of outcrops exhibiting lithofacies and microfacies variations of the carbonates. • Geochemical prospecting. Chemical data from outcrops and wells are important in the evaluation of the industrial use of carbonate rocks. Common com-
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_18, © Springer-Verlag Berlin Heidelberg 2010
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pounds considered in geochemical prospecting are CaCO3, MgCO3, SiO2, SO3, Fe2O3, Al2O3, Mn2O3, K2O as well as various minor elements. Combinations of geochemical prospecting and microfacies analysis of samples has proved quite successful (e.g. Städter 1989). Thin-section studies: Differentiation of facies types based on textural criteria and diagenetic characteristics (recrystallization, dolomitization, dedolomitization, stylolitization; see Chap. 7). Assessment of a ‘facies model’ describing the lateral and vertical spatial relationships of the facies types. This model is used in further interpreting samples studied during the exploration phase. This phase demands investigation of cuttings or cores. Exploration wells: Study of cores and cuttings. Cuttings are washed, differentiated into lithotypes using a binocular, and characterized by microfacies criteria. In addition, the total carbonate content as well as mineralogical and geochemical properties of core samples are determined (clay minerals, dolomite, main and rare elements). The number of samples analyzed depends on the rock qualities demanded (Langbein et al. 1982). Samples should be grouped statistically according to their spatial occurrence or according to particular features. This approach leads to the visual presentation of rock bodies consisting of carbonate facies units with particular physical and/or chemical properties. Mapping and differentiation of facies-controlled carbonate bodies with regard to boundaries, extension and volumes.
Exploration of pure limestones. Pure limestones with high calcium carbonate and very low non-carbonate contents are used in the glass, paper and chemical industries for water desulfurization of flue gasses, and for producing fertilizers and animal food (Kimmig et al. 2001). Pure carbonate rocks are represented by marbles and by platform and reef limestones (see Fig. 14.22). Distinct correlations exist between total carbonate content and depositional conditions: Open-marine bioherms and reefs, sand piles and open-marine platform carbonates often correspond to ‘pure limestones’, because of low or missing siliciclastic input or deposition in high-energy environments. Reasons for the high purity of many limestones are winnowing of finegrained carbonate and non-carbonate impurities in high-energy environments and rapid early cementation (Flügel 1977; Bertle 1982; Koch et al. 1984, Dimke 1997). White marbles are often pure carbonates and are industrially sought after raw materials.
Facies and Physical-Chemical Properties
Ultra-poor limestones are usually white limestones characterized by a very high calcium carbonate and very low iron contents (>99% CaCO3, <0.5% Fe2O3). An example are Late Jurassic limestones (particle-rich sponge boundstones, packstones and grainstones) occurring in Swabia in lenses that extend laterally several hundreds of meters to several kilometers across. Quality variations are controlled by diagenesis, secondary alterations and karstification. Exploration includes detailed mapping related to the lithology, facies, alterations and structure, and requires continuous sampling along drill cores with a close spacing (Koch et al. 1997; Aigner and Schauer 1998).
18.3 Facies and Physical-Chemical Properties of Carbonate Rocks Physical properties important for the industrial use of carbonate rocks include the textural composition, porosity, permeability, specific surface and the distribution of pore throat diameters, water sorption and bulk density, rock hardness; compressive, tensile and shear strength; rigidity and elasticity, refractive properties and thermal conductivity (Harben and Purdy 1991; Bellanger et al. 1993; Winkler 1994). Both physical and chemical properties are at least partly facies-controlled. Specific interrelationships exist between depositional facies, rock color and the amount of non-carbonate contents and assist in the differentiation of the purity of limestones (Dimke 1997). The relationships between depositional facies, diagenesis and porosity explain weathering types (see Sect. 18.4). Frost damage is the result of tensile stresses caused by water freezing in the pores of limestones. The saturation state and water distribution are a function of pore geometries. Depositional facies, porosity, mineralogical and chemical composition influence burning behavior and quick lime production (Gotthardt et al. 1967; Gotthardt and Wilder 1981, 1984; Gosh 1981; Butenuth et al. 1993; Ellmies 1995; Rossner and Michel in Koch et al. 1997). Different lithotypes exhibit specific burning behaviors. High quality soft quick lime is produced by pure calcilutites and calcarenites. Grain size and porosity do not affect burning or slacking. In contrast non-carbonate contents and clay admixtures reduce the quality of quick lime (Ellmies 1995; Koch et al. 1997; Frey 1998). The controls on density, porosity, permeability, hardness and strength by the original depositional texture and the course of cementation are responsible for different physical properties, which in turn influence the technological properties of limestones and dolo-
Weathering of Carbonate Rocks
mites (Bell 1981; Hozman 1983). Some micritic limestones exhibit exceptional purities and high intergranular porosities and therefore have a high aptitude for disintegration and grainy destruction. This is a good base for the production of high quality technical limestones. Rock color and color stability. The colors of limestones are influenced by environmental and diagenetic factors. Interconnected environmental factors are water energy, sedimentation rates, influx of non-carbonate material, pore water flux and the position of the oxidation/reduction boundary. The differences in these factors govern the distribution of coloring agents such as iron and its oxides, manganese and organic substance and favor the development of specific colors in different parts of marine environments. The iron content, rock color and depositional facies can exhibit distinct correlations as shown by Upper Jurassic carbonates from Southern Germany (Grunenberg 1992; Engelbrecht 1992; Dimke 1997). The recognition of these relations during an early phase of prospecting assist in a guided exploitation of the carbonate rocks.
18.4 Weathering, Decay and Preservation of Carbonate Rocks 18.4.1 Weathering of Carbonate Rocks Physical, chemical and biological weathering of subaerially exposed carbonate rocks lead to • abrasion connected with disintegration, chipping, crumbling, exfoliation, • corrosion connected with the solution of carbonate minerals, changes in the non-carbonate minerals and organic matter, causing changes in the roughness of the surfaces and in rock colors, • the formation of weathering crusts, • changes in the hardness, strength and stress patterns. These alterations are evident in material used for the construction of buildings, pavements or works of art. Carbonate rocks used as outdoor building stones (e.g. dimension stones forming walls or monuments) are affected by natural and man-made weathering agents, leading to stone decay caused by moisture and salts, frost action, chemical weathering and biological contributions (Winkler 1994). Weathering resistance depends on many, partly interconnected factors. Porosity and permeability and the fabric (expressed by grain size, grain distribution, grain form, grain type, depositional texture, mud-support and grain-support) are of primary importance, because these factors control the storage capacities of water and gasses.
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Facies-dependent controls on weathering connected with fluid sorption, dissolution and freezing are: • depositional structures (bedded carbonates are generally more susceptible to weathering by freezing than massive limestones) and inhomogenities of carbonate bodies (variations in texture, mineralogy as well as joints), • depositional texture (mud- or grain-support, grainor crystal fabric), • effective porosity and permeability, • the amount and distribution of non-carbonate material (freezing potential can increase with increasing clay mineral content or with the concentration of clay minerals within layers). These generalizations should be regarded with some caution, because of the scarceness of studies on causal relationships between the facies criteria of carbonate rocks and those physical and chemical data that control weathering patterns. Case studies. Grimm (1990) provided a wealth of information on the weathering potential of rocks commonly used for monuments in Germany. The data include geological characterization, fabric, microfacies criteria (composition, texture, diagenetic criteria) and SEM observations as well as physical parameters (specific gravity, porosity, water storage capacity and permeability). Grimm classifies weathering resistance as very good, moderate and poor. Very good and good weathering resistance are characterized by the absence or rarity of abrasion criteria. Samples with moderate resistance exhibit selective weathering of parts of the rock as well as common chipping, exfoliation and crumbling. Poor resistance is marked by strong disintegration. Using these data, three categories (I: very good and good; II: good to moderate, intermediate weathering resistance and wide resistance range; III: moderate and poor) can be differentiated, allowing the textural composition of limestones and weathering to be compared. There is a common assumption that micritic, mudsupported limestones are more heavily affected by weathering than sparitic grain-supported carbonates, because of the higher specific surface of fine-grained limestones. This is only partly supported by the data. Very good and good weathering resistance occurs with about equal amounts in mud- and in grain-supported rocks. Intermediate resistance seems to be more common in micritic limestones, whereas low weathering resistance appears restricted to grain-supported carbonates.
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Category I includes predominantly bioclastic wackestones followed by grainstones, floatstones, rudstones and mudstones. Category II also includes a broad textural spectrum due to • irregularly distributed stylolites that enhance weathering, • differences in interparticle porosities and the degree of cementation in bioturbated and non-bioturbated parts of limestone, • selective dissolution of fossils and ooids, and • selective dissolution of the micritic matrix. Category III includes limestones used as building stones in monuments, which are highly susceptible to rapid weathering in humid climates. Primary controls of the poor weathering resistance are: • differences in the lithological composition (differential weathering of carbonate and non-carbonate or hard and soft particles), • differences in grain sizes and primary inter- and intragranular porosities within various depositional layers, • differences in the dissolution of bioclasts due to different primary skeletal mineralogies, • strong solution along circum- and intragranular stylolites. In summary, depositional texture is an important but not the only control on weathering resistance, as shown by a comparison of microfacies criteria and physical data: Effective porosity and permeability, but more important, the size and size distribution of the pores (determining water saturation), act as important controls on soundness and slake durability, which in turn may influence the degree of weathering. The effective porosity of the building stones studied by Grimm (1990) varies between 0.20 and 24 vol.% for category I limestones, 0.20 and about 9% for category II, and 0.8 and 35% for category III. Most mudstones, wackestones and floatstones have porosity values <5%; values of <1% are most frequent. Most grainstone and rudstone samples exhibit porosities of >3%, often within a range between about 5 and 25%. The size of intragranular pores and solution voids can strongly affect a susceptibility to freezing. Because larger pores, e.g. within calcareous tufa or travertine, become only partly filled with water, freezing may not destroy their fabric. The pore size distribution of many mudstone and wackestone samples with low porosities is characterized by a high percentage (about 55%, up to nearly 100%) of intercrystalline pores with a size
Weathering of Carbonate Rocks
of <1 µm. High amounts of very small pores within a range of between 100 and >1 µm may impede the infill of free pores with water and therefore, the disintegration of the limestones by freezing. However, these differences are modified by pore geometry. A high water saturation potential of grainstones (depending on primary texture, selective dissolution, incomplete cementation) will negatively affect frost resistance. Although these examples can not be regarded as statistically significant for specific limestone or dolomite types, the importance of texture and diagenetic characteristics in causing weak or strong physical, chemical and biochemical weathering of carbonate rocks is evident. Further studies must focus on data leading to quantitative interpretations (using defined weathering-potential indices, e.g. the ‘microfracture index’ describing the number of microcracks in a 10 mm traverse of the thin section) and on the study of the weathering and decay of carbonate rocks by dissolution and biochemical actions (bacteria, algae, fungi, lichens, macroplants, borings). Biological weathering and the interaction of carbonate rock surfaces with bacteria, cyanobacteria, fungi and lichens play an enormous role in the weathering of buildings and works of art. It is estimated that about 70% of the biodeterioration of limestones and marbles is caused by microbes forming crusts and biofilms (Danin 1982; Saiz-Jimenez 2002). Another topic is a better understanding of the timespan involved in destructive weathering. Well-dated monuments (e.g. archaeological sites, churches or tombstones on churchyards) offer an elegant possibility for this approach (Marinos and Koukis 1988; AiresBarros et al. 1990; Düppenbecker and Fitzner 1991; May 1994). Microfacies, texture and weathering of building stones. Relations between depositional texture and weathering are evident for bioclastic and oolitic grainstones (Koch et al. 1999; Andriani and Walsh 2002). Microporosity and hygroscopic porosity are responsible for atmospheric moisture condensation in fine and medium calcarenites. Coarse calcarenite with larger pores enables rapid fluid transfer into pores, causing rapid cementation and a reduced weathering potential. Different weathering types, intensity and distribution can be also related to variations in the grain composition and clay content as demonstrated by Tertiary pelagic Globigerina limestones (Fitzner et al. 1997). Weathering indices. Precise diagnoses of weathering damage are required for evaluating and preserving
Weathering of Carbonate Rocks
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Fig. 18.1. Weathering of the Great Pyramid in Gizeh near Cairo, Egypt. The building blocks are almost pure Eocene limestones of the Mokkatam Group with only a low quartz and low organic matter content. The porosity of the material varies between 8.5 and 37.5%; most values are between 20 and 25%. The limestones comprise various textural types including fossiliferous to sparse wackestones, skeletal packstones with nummulitid foraminifera and skeletal packstones to grainstones. The few samples studied show no correlation between microfacies, porosity values and the proportion of pore throat classes. Relations may exist between grain fabric, the clay content of limestones (Rodriguez-Navarro et al. 1997) and weathering types. The photograph shows the lower part of the pyramid. Note the strongly varying weathering types of the dimension stones within the individual rows. Following Fitzner et al. (2002) the photograph displays different weathering categories. Fissures cutting cross-bedding structures (arrows 1, 4), crumbly disintegration (arrow 2), loss of stone material by backstepping and breakout (arrow 3) and detachment of flakes and platy pieces of stones (arrows 5 and 6) are evident. Height of blocks up to 1 m.
stone monuments (Fitzner and Heinrichs 2001). Based on monument mapping and the visual inspection of weathering forms, six damage indices characterizing subsequent ratings of weathering were distinguished by Fitzner and Heinrichs (2001): 0 - no visible damage, I - very slight damage, II - slight damage, III - moderate damage, IV - severe damage, V - very severe damage. These categories have been modified and adapted to specific climatic conditions (Fig. 18.1). Fitzner et al. (2002) studying ancient Egyptian monuments distinguished four weathering categories: 1 - loss of stone material characterized by back weathering, relief and breakout; 2 - discoloration characterized by the discoloration of the original rock color, dirt deposits on stone surfaces and loose salt deposits, dark crusts; 3 - detachment causing granular or crumbly disintegration, detachment of flakes and platy larger pieces of stone, detachment of crusts; and 4 - fissures independent or dependent of the stone structure.
18.4.2 Preservation The effect of carbonate stone conservation using surface sealers penetrating stone consolidants (Perth et al. 1992; Winkler 1994) or biological agents (Toepfer 1997) can be increased if the texture, porosity and permeability of the stones are studied prior to conservation and restoration (Lorenz 1998). A case study is explained in Pl. 150. Surface sealing of stone surfaces aims to prevent the access of moisture. However, pressure may develop behind the sealing by efforescence, crystal growth and/or freezing, causing new spalling and flaking. Effective conservation needs the complete drying of the stones followed by an impregnation that affects all open pores at and below the surface. A total impregnation with low viscose methylmetacrylate under vacuum pressure conditions in an autoclave, followed by a thermic treatment causing polymerization and formation of polymethylmetacrylate has proved to be a very useful method (Acrylic-TotalImpregnation-Process, Ibach Stone Conservation, Bamberg).
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Limestones: Decay and Conservation
Plate 150 Microfacies, Weathering and Conservation of Limestones The oolitic limestones, exploited near Savonnières-en-Pertois in eastern France, have been widely used as sculptural stones in Central Europe (e.g. for the Gothic Cathedral in Cologne). These grainstones consists predominantly of ooids with a high porosity (30–35%) caused mainly by the dissolution of the nuclei. Porosity distribution however, assists in the workability of the stone, the combination with low permeabilities (about 30 mD vertical and 100 mD parallel to bedding) and conspicuous differences of interparticle and intercrystalline pore neck diameters facilitates on the other hand weathering and destruction (Lorenz and Ibach 1996). The plate shows the microfacies of the not-weathered stone (–> 1), destruction types of sculptural stones –> (2–6) as well as the effect of a conservation method (–> 7, 8). Figs. 2–6 are samples of building stones used for the outer front of the Thomas Church in Leipzig, Germany. The Ibach Total Impregnation method involves soaking well dried limestones and marbles with liquid methyl-methacrylate, using vacuum and pressure. 1 Oolitic grainstone. The nuclei of most ooids and mollusk shells are dissolved, causing high oomoldic and biomoldic porosities (blue staining). Intercrystalline porosity (within ooid cortices and cement crystals) also occurs. Intercrystalline pores may act as capillary systems that hold fluids and cause long-term rock destruction. The ooids have been cemented by marine-phreatic isopachous cement crystals. Early Cretaceous (Berriasian): Savonnières-en-Pertois, eastern France. 2 Building stone. Rain-exposed surfaces of building stones are prone to freezing damage. Water intruded via intercrystalline ooid and cement pores, causing volumetric expansion and leading to cracking (arrow). The detachment of cortices is caused by pronounced leaching along growth stages of the ooid. The ooid nucleus is completely leached. Savonnières oolite used as external building stone for the Thomas Church in Leipzig, Germany. 3 Industrial pollution contributes to the deposition of gasses and various particles in rain-protected niches, leading to the formation of crusts (upper part). These 0.5 to 1 mm thick crusts consist of gypsum (white) and various soot particles (black and brownish). Gypsum is formed by the reaction of SO2 with water, producing sulphurous acid, which in turn reacts with CaCO3 (red, alizarine staining). Microbial activity has a great influence on crust formation. Gypsum also occurs within oomolds several millimeters below the stone surface (–> 5). The crusts do not act as protecting cover. Differences in gypsum and calcite with regard to thermic expansion and the reaction with fluids lead to a loosening of the uppermost centimeters of the rock. Removal of the gypsum crusts accelerates rapid weathering. Open pores are blue. Same locality as –> 2. 4 Surface of a building stone. Porosity variations (bottom, with original porosity; top, with altered porosities) are responsible for vertical and horizontal permeability differences. Red: calcite; blue: open pores; white: gypsum, dark: industrial soot particles. The yellowish color (top and corner) is produced by the thin-section glass. Same locality as –> 2. 5 Oomoldic pore, partly filled with idiomorphic gypsum crystals (formed during weathering; arrow) and brownish organic matter (produced by hydrocarbons having migrated through the limestone). Same locality as –> 2. 6 Vertically cut stone surface. The uppermost ‘weakened zone’ (–> 3) is affected subsequent to the loss of the outer gypsum crust by frost weathering, causing splitting via microcracks (arrows; blue) parallel to the surface. Same locality as –> 2. 7 Conservation. Various methods (surface sealing, acrylate total impregnation) are in use in order to protect stone sculptures and building stones from further decay. The thin section shows the totally impregnated oolite. The normally colorless acrylate (A) was stained in this sample to show that the impregnation has affected all the intra- and interparticles pores. Isolated vacuoles (VA, white) within the acrylate are caused by volumes decreasing during polymerization. Because these bubbles are enveloped by acrylate, they are not accessible for gases or fluids. Note the complete compound of the acrylate with the surface of the calcite cement crystals and within the leached ooids. This compound hinders further intrusion of moisture and gases. Dark patches are organic matter (OM; see –> 5). Broadening and rounding of cement crystal tips indicate meteoric overprint of marine-phreatic cements. Same locality as –> 2. 8 Conservation. Inhomogenities within the oolite are caused by shell layers (SL) that were more strongly affected and leached by rain and weathering, because of their primary aragonitic mineralogy. However, total impregnation also affects these microfacially different parts of the rock. Note the complete impregnation (here white) of all molds and interparticle pores with acrylate. The total adhesion of acrylate to the pore walls prevents the carbonate from taking up moisture and pollutant and results in good consolidation of the whole rock. Vacuoles within the acrylate have been stained subsequent to the impregnation and appear blue. Same locality as –> 2. –> 1–8: Courtesy of H. Lorenz (Erlangen)
Plate 150: Microfacies, Weathering and Conservation of Limestones
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Basics: Carbonate Rock Resources, Facies Control and Rock Properties Amoroso, G.G., Fassina, V. (1983): Stone decay and conservation. Atmospheric pollution, cleaning, consolidation and protection. – Materials Sci.Monographs, 11, 446 pp., Amsterdam (Elsevier) Boynton, R.S. (1980): Chemistry and technology of lime and limestone. Second edition. – 578 pp., New York (Wiley) Cucchi, F., Gerdol, S. (1989): Der Naturstein aus dem Triester Karst. – 195 pp., Trieste (Camera di Commercio Industria Artigianato e Agricultura) Fitzner, B., Heinrichs, K. (2001): Damage diagnosis of stone monuments - weathering forms, damage categories and damage indices. – In: Prikryl,R., Viles, H.A. (eds.): Understanding and managing stone decay. – Proceedings of the International Conference ‘Stone weathering and atmospheric pollution network (Swapnet 2001)’, Praha Fitzner, B., Heinrichs, K., La Bouchardiere, D. (2002): Limestone weathering of historical monuments in Cairo, Egypt. – In: Siegesmund, S., Weiss, T., Vollbrecht, A. (eds.): Natural stone, weathering phenonema, conservation strategies and case studies. – Geological Society London, Special Publication, 205, 217-239 Flügel, E., Haditsch, J.G. (1975): Vorkommen hochreiner und reinster Kalke im Steirischen Salzkammergut. – Archiv für Lagerstättenforschung, 15, 65-84 Gotthardt, R., Kasig, W. (1996): Karbonatgesteine in Deutschland. Rohstoff, Nutzung, Umwelt. – 420 pp., Düsseldorf (Beton-Verlag) Grimm, W.D. (1990): Bildatlas wichtiger Denkmalgesteine der Bundesrepublik Deutschland. – Arbeitsheft, Bayerisches Landesamt für Denkmalspflege, 50, 450 pp. Kieslinger, A. (1932): Zerstörungen an Steinbauten, ihre Ursachen und ihre Abwehr. – 346 pp., Wien (Deuticke) Kieslinger, A. (1949): Die Steine von St.Stephan. – 486 pp., Wien (Herold) Kieslinger, A. (1956): Die nutzbaren Gesteine von Salzburg. – 436 pp., Salzburg (Bergland-Buch) Koch, R., Flügel, E., Dimke, M., Hasselmeyer, B., Michel, U., Rossner, R., Sobott, R. (1997): ‘Angewandte Faziesforschung’ am Institut für Paläontologie in Erlangen. – Zentralblatt für Geologie und Paläontologie, Teil I, 1996, 130-174 Koch, R., Sobott, R., Lorenz, H.G. (1999): Der Schaumkalk (Trias, Unterer Muschelkalk) am Naumburger Dom als Baustein: Einfluß von Fazies und Diagenese auf die Gesteinsqualität. – In: Hauschke, N., Wilde, V. (eds.): Trias. Europa am Beginn des Erdmittelalters. – 449-471, München (Pfeil) Langbein, R., Peter, H., Schwahn, H.J. (1982): Karbonat- und Sulfatgesteine. Kalkstein-Dolomit-Magnesit-Gips-Anhydrit. – Nutzbare Gesteine und Industrieminerale, Monographienreihe, 5, 335 pp. Lorenz, H.G., Ibach, H.W. (1999): Marmorkonservierung durch Ibach-Volltränkung: Qualitätskontrolle und Optimierung durch mikroskopische und petrophysikalische Untersuchungen. – Zeitschrift der Deutschen Geologischen Gesellschaft, 150, 397-406
Basics: Carbonate Rocks: Resources, Facies, Properties
May, A. (1994): Microfacies controls on weathering of carbonate building stones: Devonian (Northern Sauerland, Germany). – Facies, 30, 193-208 May, A. (1997): Verwitterungsbeständigkeit und Verwitterung von Naturbausteinen aus Kalkstein. – Geologie und Paläontologie in Westfalen, 48, 185 pp. May, A. (1998): Die Verwitterung von karbonatischen Naturbausteinen in Mitteleuropa. – Geoökodynamik, 19, 71-88 Mirwald, P.W., Brüggerhoff, S. (1995): Verwitterungsvorgänge an Kalksteinen unter mitteleuropäischen Klimaund Immissionsbedingungen. – Geologisch-Paläontologische Mitteilungen Innsbruck, 20, Festschrift H. Mostler, 207-220 Oates, J.A.H. (1998): Lime and limestone. Chemistry and technology. Production and uses. – 455 pp., Weinheim (Wiley-VCH) Parwellek, T., Aigner, T. (2002): Fazies, Petrophysik und Rohstoffeigenschaften von Karbonatgesteinen des Schwäbischen Jura - ein Atlas. – Jahrbuch und Mitteilungen der Oberrheinischen Geologischen Vereinigung, 84, 257-321 Rodriguez-Navarro, C., Hansen, E., Sebastian, E., Ginell, W.S. (1997): The role of clay in the decay of Egyptian limestone sculptures. – Journal of the American Institute of Conservation, 36, 151-163 Schiele E., Behrens, L.W. (1972): Kalk. Herstellung - Eigenschaften - Verwendung. – 627 pp., Düsseldorf (Verlag Stahleisen) Schroeder, J.H. (ed., 1999): Naturwerksteine in Architektur und Baugeschichte von Berlin. Gesteinskundlicher Stadtbummel zwischen St.Marienkirche und Siegessäule. – Führer zur Geologie von Berlin und Brandenburg, 6, 230 pp. Siegel, F.R. (1967): Properties and uses of carbonates. - In: Chilingar, G.V., Bissell, H.J., Fairbridge, R.W. (eds.): Carbonate rocks. – Developments in Sedimentology, 9B, 343393, Amsterdam (Elsevier) Smith, M.R. (1999): Stone: building stone, rock fill and armourstone in construction. - Engineering Geology, Special Publication, 16, 480 pp. Snethlage, R. (1997): Leitfaden Steinkonservierung. Planung von Untersuchungen und Maßnahmen zur Erhaltung von Steindenkmälern aus Naturstein. – 215 pp., Stuttgart (Frauenhofer IRB) Warren, J. (2000): Dolomite: occurrence, evolution and economically important associations. – Earth Science Reviews, 52, 1-81 Winkler, E.M. (1994): Stone in architecture. Properties, durability. Third edition. – 313 pp., Berlin (Springer) www.abraxas-verlag.de: CD ROM. Natural Stones Worldwide. Contains 17,000 commercial names, 10,000 natural stones, and 2,000 pictures. Determination, technical and physical characteristics, as well quarry location and supplier’s addresses are given. Further reading: K202, K209
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19 Microfacies and Archaeology
19.1 Questions and Methods
19.2 Building Stones
Archaeologists are faced with the problem of the provenance of materials used in ancient buildings, sculptures or ceramics. Archaeometric methods applied to source studies include analyses of element distributions, stable isotope patterns and studies of the mineralogical and petrographical composition (Herz 1987; Riederer 1987; Herz and Waelkens 1988; Gibson and Woods 1990; Walsh 1990; Waelkens et al. 1992; Orton 1993; Rapp and Gifford 1995; Herz and Garrison 1998; Pollard 1999). ‘Geoarchaeology’ (Rapp and Gifford 1998) is an exciting and promising interdisciplinary approach. Archaeometric studies of limestones and marbles consider petrographical data, stable isotopes and strontium isotopes, trace elements, and cathodoluminescence signatures.
Building stones are rocks suitable for use in constructions. They are chosen for their properties of durability, economy and attractive beauty. Building stones that are quarried and prepared in regularly shaped blocks are called dimension stones. Provenance analysis of carbonate and non-carbonate stone material used in the construction of ancient buildings assists in deciphering the exploitation history of antique quarries and provide information on changing local and regional sources. Knowledge of the source of marbles and other precious stone material used for buildings and works of art gives information on former trading patterns. Specific questions that could be answered by microfacies-based studies of building stones are: Did ancient builders only exploit local sources or did they use different sources for buildings of different social rank? What trading patterns can we discern from the distribution of particular materials? How did the available material affect architectural styles? Are building stones repeatedly reused? Some of these questions are touched on in a faciesbased provenance study of the building stones of Roman Sagalassos in southwestern Turkey (Degryse et al. 2003) who identified distinct relationships in time changing local sources, changing architecture styles and the attempt to use material exhibiting better technological properties (e.g. strength).
Microfacies analysis is playing an increasingly more important role. The microfacies approach can be used both for carbonates and materials made of cherts. Archaeological objects consisting of limestone or dolomite are typified according to textures, composition and paleontological criteria. The combination of these criteria with geological and facies data provides a highly valuable means for provenance determinations of antique building stones, mosaics, works of art and ceramics. Chapter 19 summarizes several case studies. Microfacies studies contribute to • differentiating the source of material used as building stones, mosaics and works of art, • recognizing the source area of raw materials used in the manufacture of pottery as well as the production sites, • defining time-dependent technologies used in the fabrication of ceramics, • understanding trade routes, and • evaluating provenance analyses based on isotopic and geochemical data.
19.2.1 Building Stones: Methods The provenance analysis of ancient building stones is based on the investigation of lithology and petrography, stable isotopes and trace elements as well as physical properties. Carbonate building stones are studied with regard to: • Rock color. Weathered and unweathered.
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_19, © Springer-Verlag Berlin Heidelberg 2010
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• Lithology. Limestone? Dolomite? Marly limestone? Cherty limestone? • Sedimentary structures. Massive limestones or indications of bedding? Cross-bedding structures? Reef limestone? • Microfacies. Matrix-dominated or grain-dominated limestones? Mud-support or grain-support? Dunham-classification types? Dominating grain types? • Paleontological criteria. Dominating fossils (e.g. shells, foraminifera, calcareous algae). Fossils that give indications of paleoenvironment and age. • Sedimentary fabrics. Biofabrics, lamination, burrowing, fenestral fabrics, birdseyes, nodular fabrics, discontinuity surfaces. • Diagenetic features. Cement, recrystallization, diagenetic fabrics (compaction, stylolites).
In the case of quarries located within major stratigraphic units, the chemical composition (e.g. SiO2Al2O3-CaO plots) and the occurrence and distribution of minor and rare elements allow, together with thinsection criteria building stones to be attributed to specific source areas (Harrell 1992; Klemm and Klemm 1973). A combination of textural, geochemical and petrophysical criteria (water absorption, ultrasound transmission velocity, porosity, density and compressive strength) has been successfully applied for locating ancient quarries used for constructing historical buildings (Galán et al. 1999).
These criteria or at least some of these criteria will characterize facies types that can be compared with facies types of adjacent outcrops and ancient quarries. Many ancient buildings used material from domestic sources, but also preferred specific materials, often exported from far distant regions. That is true particularly of marble (see Sect. 19.6).
There are many examples relating materials used as building stones to potential or proved source areas, but rather few investigations based on the differentiation of limestones in relation to facies criteria. The following examples demonstrate how detailed studies can contribute to answering specific archaeological questions.
19.2.2 Building Stones: Examples
Fig. 19.1 Building stones of the giant outer wall of Hattusa, the capital of the Hittite reign. The blocks of the wall are limestones, particularly reddish coral limestones. Especially the stones in the lowermost row show abundant reef corals. Microfacies and paleontological criteria prove that the material used was derived from Early Cretaceous limestones outcropping in and near Hattusa. The width of the largest block in the center is 80 cm. Bogazkale about 150 km east of Ankara, central Anatolia, Turkey.
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The Giant wall of Hattusa (Turkey). Hattusa, the capital of the Hittite reign (situated near Bogazkale in central Anatolia), is surrounded by a giant 6 km long wall built in the middle of the second millenium BC (Neve 1996; Fig. 19.1). The blocks of the wall are between 40 and 80 cm high. Most blocks are gray and red limestones. The latter exhibit conspicuous coral structures and correspond to coral framestones (Flügel and Flügel-Kahler 1997). Thin sections show dendroid corals affected by strong diagenetic alterations characterized by cementation, dissolution and internal sedimentation (Pl. 36/4; Fig. 7.15). The corals exhibit thick biogenic encrustations. The sediment between the corals is a bioclastic peloidal packstone (Pl. 151/8). Microfossils indicate an Early Cretaceous age of the limestones. Outcrops of reddish and gray limestones occur within the boundaries of Hattusa. The material used for the wall is of domestic origin.
e.g. that most mosaic material came from sources near or very near to the site of the mosaic. Microfacies provides a mean of distinguishing between local, regional and imported material, or recycled material, e.g. tesserae made from off-cuts of building stones. Local or domestic material is defined as material derived from rocks existing at the mosaic site. Regional material comes from sources within a distance of few to several tens of kilometers. The term ‘imported’ refers to material imported from far distances.
Roman building stones on the Magdalensberg, Carinthia (Austria). Starting in the first century BC and terminating in the middle of the first century AD, a Roman city developed on the slopes of the Magdalensberg near St.Veit in Carinthia (Piccottini 1989). The building stones used for houses and the public buildings were investigated by F. Thiedig and his team with regard to lithology and source (Thiedig 1998; Thiedig and Wappis 2003). A detailed survey of more than 47,000 building stones, including mapping and petrographical studies showed that more than 95% of the excavated walls consist of volcanic rocks outcropping directly next to or within a distance of a few kilometers to the Roman city. 5% of the building stones are marbles, travertines and Pleistocene pebbles, including crystalline rocks, limestones and sandstones. Marbles used in public building and for grave materials, came from marble quarries located within a distance of a few to some tens of kilometers. Travertine quarries were located at a distance of about 30 kilometers.
• Color. Macroscopic assessment and definition of rock color groups. The principal colors of mosaic stones are white and yellow, black and gray, pink and red as well as orange and brown. Blue and also some green tesserae are usually glasses. Many red tesserae were made of fragments or ceramic tiles. Rock colors can be altered by burning, an effect probably also known to Roman mosaicists. The assessment of colors of carbonate tesserae must take into account that broken, cut, ground, smoothened or polished tesserae can exhibit considerable differences in color gradations and shades. • Sampling based on rock-color categories. Because color effects are essential for the beauty of polychrome mosaics, samples must consider the principal rock color categories. The composition of non-lithic tesserae (ceramics, bricks and tiles, glass) is studied by microscopical, chemical and isotopical criteria. Sampling has to consider the relative abundance of differently colored tesserae within the mosaic. Abundant tesserae may be of local origin. Rare rock types may indicate the use of specifically selected and possibly imported material. For statistical reasons at least five to ten tesserae of each color group should be taken. Tesserae identical in color are not a priori identical in microfacies and age of the limestone! • Microfacies analysis. Texture (Dunham- or Folk classification), matrix and grain types, and fabrics (lamination and fine bedding, burrows, nodular structures, calcite-filled open-space structures, calcite veins, microbreccias) allow a preliminary grouping of limestone samples within the categories defined by rock color. Dolomite samples should be differentiated according to the size, shape and boundaries of crystals and the type of crystal fabrics. Microfossils are impor-
19.3 Mosaic Material Mosaics consist of tailored stones and other material (e.g. glass or ceramics) juxtaposed to produce specific patterns and designs. The individually tailored stone ranging in size less than one centimeter and a few centimeters is called tessera. Most mosaic studies concentrate on stylistic groupings, design development and art history. Provenance analyses of mosaic stones are rare (see review in Flügel and Ch. Flügel 1997), although thin-section based studies can be used for testing some common assumptions,
19.3.1 Mosaic Material: Methods Microfacies-based provenance analysis of tesserae (stone or other hard material cut in a cubical or some other regular shape (Pl. 151/1-5) used for floor and wall mosaics, comprises eight stages:
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tant in determining the age and depositional environment of limestones used for tesserae. • Microfacies types. Because of the difficulties involved in estimating lithological and textural variations in small mosaic stone samples, microfacies types of tesserae are usually more strictly defined than microfacies types used in geological facies analysis. Comparison with outcrop samples may lead to a reduction in the number of microfacies types of tesserae. These types may be lumped together, but also subdivided by additional criteria (e.g. stable isotope composition). • Standard Microfacies Types. Microfacies types can be summarized into Standard Microfacies that facilitate the estimation of carbonates produced in different depositional environments. • Comparison with outcrop data. Comparison of geological and paleontological data and the microfacies inventory of carbonates within a distance of at least 50 kilometers to the mosaic sites with the microfacies types of the tesserae. • Evaluation of bio- and lithostratigraphic units. Assessment of carbonate tesserae to specific geological time/rock units followed by suggestions of potential provenance strata. • Discussion of the results with archaeologists. Provenance assessments may satisfy geologists, but archaeology colleagues may tell you that your provenance proposal can not work, because the region where the mosaic material is believed to have come from was not accessible to Romans, because Germanic tribes had occupied the region.
Mosaic Material
position and fossils, and potential provenance areas derived from geological data and comparative samples. Important results are: • Punians and Romans used different mosaic materials. The Late Punian pavement mosaics (‘pavimenta punica’) consisting of mosaic stones or stone fragments irregularly embedded into a mortar floor) are dominated by only a few limestone types occurring relatively near to the mosaic sites (Pl. 151/4, 5). • Roman tesserae exhibit a significantly higher diversity of microfacies and limestones types (8 microfacies types) than Punian mosaics (3 microfacies types). This reflects the geographically wider extension of quarrying (within a distance of up to about 100 km) and directed exploitation during Roman times. • Some of the Carthage tesserae were made from offcuts of building stones, as indicated by identical microfacies types in limestones used for building stones. • Most of the tesserae used for the sixth century Byzantine Rotunda at Damous-el-Karita (Dolenz 2001; Pl. 151/1-3), the largest Christian basilica in North Africa, correspond in microfacies types and age to limestone material already used in Roman times.
The following examples (Flügel and Ch. Flügel 1997) involve simple pavement mosaics consisting of black and white stones and polychrome mosaics exhibiting different colors and designs. The main targets of these investigations were the adaption of microfacies methods to mosaic studies, the estimation of the contribution of local, regional, imported and recycled material, and the evaluation of time trends in the use of different carbonate materials for tesserae.
Roman mosaic in Kraiburg, southern Bavaria, Germany. The polychrome pavement mosaic exhibits black, gray, white, yellowish, reddish and red colors. The tesserae consist of limestones, siliciclastics and a few ceramics. Carbonate mosaic stones exhibit 18 microfacies types, sandstones can be attributed to 8 lithotypes. All the rock types used for the Kraiburg mosaic can be found in Triassic, Jurassic and Cretaceous sediments of the Northern Calcareous Alps. The material of coeval Roman mosaics in South Bavaria has been interpreted as imported from the Salzburg region. This interpretation can not be ruled out, but several criteria support another interpretation of provenance. Most rock types also occur in glacial limestone pebbles and boulders transported from the Alps to the mosaic site during the Pleistocene and distributed by Holocene river systems. The tesserae exhibit curved outlines and bent surfaces, which means they may have been produced from rounded pebbles. Several microfacies types correspond closely to microfacies of boulders occurring in river terraces adjacent to Kraiburg. The most interesting result of the case study is that tesserae of a specific color can represent up to 8 microfacies types that indicate carbonate rocks of different age and geological position.
Late Punic, Roman and Late Byzantine mosaics from Carthage near Tunis. Mosaic stones were studied with regard to microfacies types defined by texture, com-
Late antique mosaics of the Hemmaberg near Globasnitz, southern Carinthia, Austria. The churches of Roman Iuenna built in the fifth and sixth century AD
Microfacies studies of tesserae can be supplemented by geochemical analysis (e.g. Mampbelli et al. 1988) and stable isotope data (Bollin and Maggetti 1996; Flügel and Ch. Flügel 1997).
19.3.2 Mosaic Material: Examples
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Fig. 19.2. Polychrome mosaic floor of a late antique church from the Hemmaberg near Globasnitz, southern Carinthia. Most white and pink mosaic stones are Triassic dolomites. Black stones include Triassic and Jurassic as well a few Carboniferous limestones. No distinct correlation exists between rock color and specific microfacies types. Most of the material was derived from outcrops exposed within a distance of a few kilometers adjacent to the mosaic site. In addition some glacial boulders have been used in the production of mosaic stones. The average size of these stones is 1.5 cm.
were a center of early Christian pilgrimage (Fig. 19.2). The churches were decorated by rich polychrome mosaic floors. Stylistic analyses of the mosaics prove that the mosaicists knew the repertoire of mosaics in the northern Adria region (Glaser 1996). Did the mosaicists import material from the Adria region or did they use local material? Pink mosaic stones are fine to medium-crystalline dolomite. White mosaic stones consist of dolomite and limestone and exhibit 5 microfacies types. Light gray tesserae are limestones exhibiting 4 microfacies types. Most microfacies types observed in the Hemmaberg mosaics can be attributed to Middle and Late Triassic carbonates exposed along a distance of a few kilometers. The idea that material was imported from the Adria region can be rejected. A central on-site production of the Hemmaberg tesserae is also supported by the enormous amount of rock material necessary for the large mosaics. The size of the mosaic area, the number of mosaic stones necessary to cover this area, and the average size and volume, means about 4000 kg of material must have been available for the tesserae. Stable isotope data of Hemmaberg and Carthage tesserae can be used in evaluating the differentiated microfacies types. Close-spaced δ18O/δ13C patterns of apparently different microfacies types indicate a too rigorous definition and the necessity of uniting these types.
19.4 Works of Art Many antique sculptures were made of marbles, others of limestones or other rock material. An examination of the material may be necessary to clarify the production area or recognize falsifications (Riederer 1994).
19.4.1 Works of Art: Methods Limestone sculptures are studied by means of thin sections, marbles by geochemical methods and thin sections (Sect. 7.9; Pollini et al. 1998). Available thin-section samples are usually very small, because archaeologists fear that the sculpture could be injured. However, even very small thin sections can provide a great deal of information as demonstrated by the following example. 19.4.2 Works of Art: Example Fig. 19.3 shows a Roman sculpture from the middle of the first century AD, tentatively interpreted as the head of the Goddess Iuno. The head was said to have been found in the Roman province of Raetia, near Augsburg in southern Bavaria, but this assignment has been questioned in the light of stylistic arguments (Eingartner
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and Ch. Flügel 1999). The study of thin sections of four very small samples taken from the back side of the sculpture prove that the sculpture was made of a limestone common in the Salzburg area but not in southern Bavaria near Augsburg (Flügel 1999). Microfacies and fossils are characteristic of Late Triassic reef limestones (Fig. 19.3). The ‘Upper Rhaetian Reef Limestone’ is typified by organic frameworks composed of corals, sponges, solenoporacean red al-
Microfacies and Archaeology
gae and strong biogenic encrustations, and by foraminifera and microproblematica bound to small cryptic cavities between reef-building organisms. All these criteria can be found in thin sections made from the tiny samples of the Iuno head. The identification of the sculpture material as Late Triassic reef limestone supports the interpretation of the head as part of a divine sculpture originally located in Iuvavum/Salzburg (Eingartner and Ch. Flügel 1999).
Plate 151 Microfacies and Archaeology: Mosaics, Ceramics and Building Stones Provenance analysis of stone artefacts, ceramics and ancient buildings is an essential tool of archaeology aiming for an understanding of ancient life, trade and technology. A number of archaeometric techniques have been employed comprising microscopic characteristics as well as mineralogical and geochemical criteria. Microfacies studies can be successfully applied to problems related to source, technological properties as well as the decay and conservation of carbonate rock materials. The plate illustrates some of the uses of microfacies data in the investigation of mosaics (–> 1–5), ceramics (–> 6–7) and building stones (–> 8). 1 Mosaic. Late antique-byzantine mosaic from a subterranean rotunda near the Damous-el-Karita in Carthage (Tunisia), the largest Christian basilica of Roman North Africa. Sixth century AD. The polychrome mosaic consists of yellowish gray (bottom row), white (central row), black and dark gray (upper row) and red (brown ochre) pelagic limestones representing seven microfacies types. Note the differences in texture and microfracture patterns. The material corresponds in microfacies and the Upper Cretaceous age to material already used in Roman mosaics of the 1st century AD. Thin section. Size of individual mosaic stones 1.0 to 1.5 cm. 2 Detail of –> 1. Black foraminiferal wackestone with Cretaceous globigerinids. Carthage is flanked by bedded Late Cretaceous limestones on the west, south and southeast, at a minimum distance of ten kilometers. 3 Detail of –> 1. Another limestone type used in the mosaic is a black sponge spicula-radiolaria wackestone with fusiform spicules that may appear as spar-filled circles in cross sections. Larger spheres are radiolarians. 4 Late Punic mosaic (pavimentum punicum) composed of monochrome white limestone tesserae embedded in a mortar floor. The limestone used is an echinoderm packstone, probably derived from turbiditic intercalations of shallow-marine material within Upper Cretaceous pelagic limestones exposed within a distance of about 20 km. Central Carthage. Excavation of the German Archaeological Institute Rome. Thin-section. 5 Detail of –> 4. Echinoderm packstone with the foraminifer Omphalocyclus Bronn indicating a Maastrichtian age of the source rock. Carbonate rocks used in late Punic mosaics and pavements are microfacially less diverse than Roman mosaics. Central Carthage. 6 Ceramic. Thin-section cut from a sherd of Punic pottery. Microfossils (diverse globigerinid foraminifera) and fragments of crystalline rocks (arrows) are incorporated into the clay matrix. Fossils as well as the mineralogical composition of the rock fragments indicate that the clays used came from Tertiary sediments of the coastal plains near the ancient city of Amathus. Good preservation of the calcitic tests of the foraminifera indicates low-temperature firing of the ceramic piece. Excavation northeast of Limassol, southern Cyprus. 7 Ceramic. Microfacies analysis of ceramics assists in the search for the area in which the pottery was produced. Fabric and microfacies of lithic inclusions within the clay matrix can be used in distinguishing local production and imported ceramics, as demonstrated by the Roman ‘Auerberg’ ceramic (1st century AD). This ceramic is characterized by marble inclusions (M) whose source area in the Southern Alps was identified by textural and stable isotope data. Marble grains have been added to the clays in order to improve the quality of the material. Other inclusions are wood and charcoal (CH). The arrows point to shrinkage structures, indicating low-temperature firing. Lorenzberg near Kempten, southern Germany. 8 Building stone. The provenance of building stones can be recognized from characteristic geological data, including microfacies and paleontological criteria. Thin-section studies show that parts of the large wall of the Hittite capital Hattusa (now Bogazkale) in central Anatolia (14th century BC) were built with limestones derived from local carbonate rocks outcropping within the area of the ancient city. The rock is a coral bafflestone. Encrusting microfossils indicate an Early Cretaceous age of the material used as building stones. –> 1–5: E. Flügel and Ch. Flügel 1997, 8: E. Flügel and Flügel-Kahler 1997
Plate 151: Microfacies and Archaeology
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Fig. 19.3. Roman sculpture interpreted as the head of Goddess Iuno. Maximum height 28.2 cm. The sculpture was made of a yellowish-white limestone. It was said to have been found near Augsburg. Microfacies criteria indicate a provenance in Iuvavum (now Salzburg, Austria). Courtesy of Archäologische Staatssammlung München. Thin section of material taken from Iuno’s head. Late Triassic reef limestone. The framestone fabrics exhibit a succession of three different coralline sponges (SP 1 and SP 2, inozoan sponges; SP 3, spongiomorphid sponge) overgrown by solenoporacean red algae (SO). Small facies-characteristic foraminifera and echinoderms also occur. After Flügel (1999).
19.5 Ceramics
19.5.1 Ceramics: Methods
The major aims of archaeometric studies of ancient ceramics are to assist in typological classifications and to provide qualitative and statistical classifications of sherds (Francovich 1992; Schneider 1995). Other questions relate to the provenance of the raw materials and the techniques used for the fabrication of the ceramics (Olcese 1991; Orton et al. 1993; Daszkiewicz and Schneider 2001; Tite 2002). These data are important in the discussion of trade routes of ancient pottery. Guidelines for field work and laboratory methods were published by Schneider (1989), current developments in ceramic petrology are summarized by Middleton and Freestone (1991).
Thin-sections of ceramic sherds exhibit: • A fine-grained homogeneous or inhomogeneous matrix representing the fired clay material. Various grain size boundaries (0.002, 0.02 or 0.1 mm) are used to discriminate matrix and grains. • Isolated temper grains (also called inclusions: rock fragments, minerals, fossils) floating within the matrix and corresponding to natural constituents of the clay or to artificial admixtures that control the tempering processes of the pottery. • Pores reflecting differences in firing temperatures and firing atmosphere or caused by a later dissolution of parts of the sherds (e.g. during burial of the archaeo-
Ceramics
logical material). Frequency, size, shape and the orientation of macro- and micropores assist in understanding production technologies. Round pores may suggest high firing temperatures, irregular and elongated pores relatively low firing temperatures (Middlestone et al. 1995). The identical porosity values of typologically uniform ceramic sherds can indicate that comparable techniques were used. Pore structure analysis is important in grouping ancient potteries with regard to firing temperatures and techniques (Gosselain 1992). The mineralogical composition of the matrix can not be sufficiently resolved by petrographical means. Matrix and temper grains are studied by means of X-ray fluorescence analysis and X-ray diffraction as well as neutron activation analysis and Mössbauer spectroscopy (Riederer 1987; Wagner et al. 1986). The temper grains are investigated in thin sections of ceramic sherds or in artificial grain fractions derived by crushing and sieving of the samples (Hagn 1985). The mineralogical composition of temper grains and matrix reflects burning techniques. Temper grain analysis directed to the understanding of the provenance and technology of ceramics must consider the following points. • Mineralogical composition: Common natural admixtures of pottery clays are quartz, feldspars, micas, calcite and cherts. In addition, other minerals and rock fragments (e.g. amphiboles, pyroxenes, epidote, opaque ore minerals; quartzite, granite, crystalline rocks) can
Fig. 19.4. Roman ceramics with white marble tempering. ‘Auerberg pots’. These Roman pots from the first century A.D. were imported from the border region between Austria and Italy and found in a Roman settlement on the Auerberg near Bernbeuren in southern Bavaria. White dots visible on the pots are marble grains which had been mixed into the clay. The largest pot is 19 cm high.
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occur, depending on the local geological situation and the genetic history of the clay. Artificial admixtures, added by the potters to improve the tempering process, include quartz, various minerals, marble and rock fragments and crushed shells as well as ash, slacks, glass, ore grains, crushed pottery and organic material, e.g. straw, wheat and plant stems. Thin section studies of argillaceous inclusions require specific descriptive methods (Whitbread 1986). • Shape: Differences in the angularity and roundness of temper grains especially quartz grains, may provide a means of discriminating natural and additional temper material. • Size and sorting: The maximum size of grains provide a first approximation in differentiating ceramic types (Riederer 1985; Schubert 1986). Distinct grain size hiati between matrix and temper components can point to an addition of temper grains, whereas a gradual transition of temper grain sizes into matrix grain sizes should indicate the lack of additions. However, caution is necessary, because of the possibility of a natural mixture of clay matrix and coarse temper grain fractions (e.g. in lake sediments). • Orientation of temper grains: Orientation patterns seen in ceramic thin sections provide clues to the techniques used in fabrication (Philpotts 1994). Pronounced orientation patterns reflect the mode of the manufacture of the pottery. • Frequency: Important data describing variations in temper composition are the amount of temper grains.
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Ceramics
Fig. 19.5. Thin section of an Auerberg pot sherd. The brown clay matrix shows parallel blueish shrinkage pores. The pores as well as some temper grains (left upper corner) are open indicating ‘lime blowing’. The marble temper grains (center and left bottom corner) have partially lost their fabric due to the burning process. The altered marble grains exhibit a skeletal structure that still reflects more or less the original twin and cleavage lamellae.
Frequency can be measured by (a) the number of grains within a defined area, usually one square millimeter, (b) the number of grains along transect lines of 10 cm length, and (c) the percentage of grain size fractions (<0.1, 0.1–0.5 and >0.5 mm). Maximum grain size is another often used measure. Riederer (1985) proposed a semiquantitative classification of ceramic fabric types based on increasing grain size and an increasing number of temper grains per sq.mm. These data may reflect differences in provenance but also different production techniques. Point-, line-, ribbon- and areacounting are applied to ceramics (Middleton et al. 1985; see Sect. 6.2). • The microfacies of small carbonate grains occurring as inclusions in coarse ceramics provide important clues to the provenance of the raw material (Pl. 151/6-7). Carbonate rock fragments and fossils (e.g. foraminifera, mollusk shells) occurring as isolated skeletal grains within the matrix or enclosed within rock fragments are of prime importance. The dating of fossils and occurrence of specific microfacies types in combination with the regional geology data facilitate the location of potential source areas of the raw material. This assists in distinguishing between locally produced and imported pottery. The combination of petrographic and microfacies characteristics of carbonate temper grains with carbon and oxygen isotope data is a highly successful tool for answering archaeological questions, as shown by the following example.
19.5.2 Ceramics: Example A specific black coarse pottery typified by a triangular rim profile was in use in several Roman provinces (Ch. Flügel 1999) during the first century AD. Originally
described from the Auerberg near Bernbeuren in southern Bavaria, this type has been found between Kempten in Bavaria and Carnuntum near Vienna in Austria and between the Danube and South Tyrol and Friuli in northern Italy. Archaeological analyses proved that the coarse ware named ‘Auerberg pots’ (Fig. 19.4) includes imported material as well as local production. Both products exhibit a conspicuous tempering by white marble grains (Pl. 151/7). The addition of marble and limestone as temper grains facilitates the plasticity and workability of clay, influences shrinkage processes during drying and after burning, improves porosity development during burning, and increases resistance to thermic shocks (Gibson and Woods 1990; Santoro et al. 1988) The marble inclusions were studied by means of thin sections (Fig. 19.5) and stable isotopes of the temper grains. The composition of the ceramic matrix was investigated using neutron activation analysis and Mössbauer spectroscopy. The latter methods indicate distinct differences between local and imported Auerberg pots (Ch. Flügel et al. 2000). The isotope analyses of the marble temper grains (Ch. Flügel et al. 1997) resulted in the recognition of different δ13C/δ18O patterns for samples from different areas. Isotope patterns of black Auerberg pot sherds found in southern Germany agree with patterns of marbles occurring in South Tyrol. Thin sections of these marbles show differences in crystal fabric, crystal size and accessory minerals, which point to the addition of a specific marble type (Sterzing marble) to the clay used for the production of the pots. Most probably the fine temper grains (average size between 0.5 and 1.5 mm) represent off-cuts of the production of marble building stones located near Lienz/Aguntum in eastern Tyrol, Austria (Ch. Flügel et al. 2004).
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Fig. 19.6. Confirmation of a legend. A: Ooids of Cleopatra’s beach, Sedir Island, Turkey. The ooids are characterized by well-defined concentric lamination and nuclei consisting of skeletal grains and clastic grains. White particles are siliciclastics. B: Ooids from the beach sand at the coast west of Alexandria, Cleopatra’s capital. The great similarity in size and microfabric as well as mineralogical composition and geochemical data support the old legend that the ooid sand at Cleopatra’s beach on Sedir Island was shipped by the Roman ruler Marcus Antonius from Egypt to the Turkish island to construct a bathing area for Cleopatra. After El-Sammak and Tucker (2002). Portrait of Cleopatra VII: Courtesy of Ägyptisches Museum Berlin.
19.6 Marble Studies Marbles in the archaeological sense include all polishable sedimentary, volcanic and metamorphic stones (e.g. carbonate rocks, porphyry, granite, basalt). Calcite and dolomite marbles are studied in polished slabs and thin sections with regard to crystal fabrics, orientation and size of crystals as well as noncarbonate minerals, which are often responsible for variations in rock colors (Walsh 1990; Dodge 1991; Ward-Perkins 1992; Sect. 7.9). Postcrystalline deformation criteria (curved twin lamellae, intertwinned grain boundaries, undulose extinction), accessory minerals and the orientation of calcite crystals may be of help in distinguishing marbles from different regions. Stable isotope patterns, rare element data, cathodoluminescense and crystal fabric analysis are useful in discriminating ancient marbles of specific provenance (Walker and Matthews 1990; Moens et al. 1990; Riederer 1990; Margolis and Showers 1990; Newman 1990; Barbin et al. 1992; Cabanot et al. 1995; Maniatis et al. 1995; Cramer et al. 1998; Schmid et al. 1999, 2000; Lapuente et al. 2000; Fant 2001). These methods should be used in combination
with the study of petrographical criteria, because of possible coincidences of isotopic patterns of marbles from different localities (Germann et al. 1988; Fiorentini and Hoernes 1996). Correct provenance assignment also helps in detecting forgeries related to copies of antique sculptures.
19.7 Antony and Cleopatra: Tracing a Famous Love Affair Legend tells us that the Roman ruler Mark Antony shipped white sand from Alexandria in Egypt to Sedir Island in the southeastern Aegean Sea sometime around 35 BC to create a pleasant beach for his lover, the famous Egyptian queen Cleopatra VII. ‘Cleopatra’s beach’ is one of the most famous in Turkey for its distinctive white sand, composed largely of calcareous ooids. This type of sand is a complete anomaly on the northeastern and eastern shores of the Mediterranean. El-Sammak and Tucker (2002) have found good evidence that lends credence to the Antony and Cleopatra story.
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Cleopatra’s beach on the small Sedir Island in Gökova Bay not far from Marmaris in the southeastern Aegean Sea is a small pocket beach separated by a man-made wall from the back-shore vegetation and soil. The thickness of the ooid layer increases from zero near the stone wall to about 80 cm near the water edge. The bay is bordered by Tertiary conglomerates. The ooids range in diameter from 0.1 to 1.0 mm. Most ooids are spherical, but there is some variation in shape depending on the shape of nuclei (Fig. 16A). The latter are microcrystalline carbonate grains, skeletal fragments and siliciclastic grains. The cortex consists of tangential, predominantly aragonitic lamellae. To check the old legend, the ooids of Cleopatra’s beach were compared with ooids from oolitic sands on the coast west of Alexandria (Fig. 16B). Size, shape, the laminated cortex and the predominance of aragonite correspond to criteria found in Cleopatra’s ooids. Were the ooids of Cleopatra’s beach formed in situ? Low-energy hydrodynamic conditions, low average water temperature, and the lack of CaCO3 oversaturation and carbonate bedrocks in the vicinity contradict an in-situ formation as high-energy ooids. A generation as low-energy ooids is extremely unlikely because of the absence of characteristic shapes, sizes and radial cortical structures (see Sect. 4.2.5). The conclusion is that Cleopatra’s ooids were not formed near Cleopatra’s beach but were brought to Sedir Island from Cleopatra’s capital Alexandria. Detailed investigation of carbonate grains using microfacies methods fully support this episode in a famous love story. Basics: Archaeology and microfacies Attanasio, D., Armiento, G., Brilli, M., Emanuele, M.C., Platania, R., Turi, B. (2000): Multi-method marble provenance determinations: the Carrara marbles as a case study for the combined use of isotopic, electron spin resonance and petrographical data. – Archaeometry, 42, 257-272 Barbin, V., Ramseyer, K., Decrouez, D., Burns, S.J., Chamay, J., Majer, J.L. (1992): Cathodoluminescence of white marbles: an overview. – Archaeometry, 34, 175-183 Barbin, V., Ramseyer, K., Fontignie, D., Burns, S., Decrouez, D. (1992): Differentiation of blue-cathodoluminescing white marble. - In: Waelkens, M., Herz, N., Moens, L. (eds.): Ancient stones: Quarrying, trade and provenance. Interdisciplinary studies on stones and stone technology in Europe and Near East from the Prehistoric to Early Christian period. – Acta Archaeologica Lovaniensia, Monographiae, 4, 231-235 Benea, M., Müller, H.W., Schwaighofer, B. (1993): Korngrößenuntersuchung zur Differenzierung von Marmoren. – Studia Univ., Babes-Bolyai, Geologia, 38, 71-74 Bergamini, M.L., Fiori, C. (1995): Characterization of limestones, marbles and other stones used in ancient mosaics. – ASMOSIA IV, Proceedings, 1999-207 Cabanot, J., Sablayrolles, R., Schenck, J.-L. (eds., 1995): Les
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marbres blancs des Pyrénées, approaches historiques et scientifiques. Entretiens d’archéologie et d’histoire. – 312 pp., Saint-Bertrand-de-Comminges (Musée archéologique départemental) Chavalas, M.W. (1999): Building in Egypt: Pharaonic stone mansonry (review). – 316 pp., Oxford (University Press) Craig, H., Craig, V. (1972): Greek marbles: determination of provenance by isotope analysis. – Science, 176, 401-403 Degryse P., Muchez, P., Loots, L., Vandeput, L., Waelkens, M. (2003): The building stones of Roman Sagalassos (SW Turkey): facies analysis and provenance. – Facies, 48, 9-22 Dodge, H. (1991): Ancient marble studies; recent research. – Journal of Roman Archaeology, 4, 28-50 El-Sammak, A., Tucker, M. (2002): Ooids from Turkey and Egypt in the eastern Mediterranean and a love-story of Antony and Cleopatra. – Facies, 46, 217-228 Eingartner, J., Flügel, Ch. (1999): Der Kopf einer weiblichen Gottheit in der Prähistorischen Staatssammlung München. – Bayerische Vorgeschichtsblätter, 64, 377-379 Fiori, C., Barboni, R., Saragoni, L. (1998): Marmi e altre pietre nel mosaico antico e moderno. – Quaderni Istituto Ricerche Tecnologiche per la Ceramica, 8, 302 pp. Flügel, Ch. (1999): Der Auerberg III. Die römische Keramik. – Münchener Beiträge zur Vor- und Frühgeschichte, 47, 237 pp. Flügel, Ch., Flügel, E., Häusler, W., Joachimski, M., Koller, J., Baumer, U., Wagner, U. (2004): Roman coarse ware from Bavaria, Austria and Northern Italy. – In: Wagner, U. (ed.): Mössbauer Spectroscopy in Archaeology. – Hyperfine Interactions Topical Issue, 154, 231-251, Dordrecht (Kluwer) Flügel, Ch., Flügel, E., Joachimski, M., Wagner, U. (1997): Auerberg black fabric – an archaeological and archaeometric approach to Roman rough ware. – Rei Cretariae Romanae Factorum Acta, 35, 85-88 Flügel, Ch., Joachimski, M., Flügel, E. (1997): Römische Keramik mit Marmormagerung: Herkunftsbestimmung mit Hilfe von stabilen Isotopen (Auerbergtöpfe aus Süddeutschland). – Archäol. Korrespondenzblatt, 27, 265-284 Flügel, E. (1999): Microfacies-based provenance analysis of Roman imperial mosaic and sculpture materials from Bavaria (Southern Germany). – Facies, 41, 197-208 Flügel, E., Flügel, Ch. (1997): Applied microfacies analysis: Provenance studies of Roman mosaic stones. – Facies, 37, 1-48 Flügel, E., Flügel-Kahler, E. (1997): Der rote Korallenkalk in der Hethiter-Mauer von Bogazköy (Anatolien): Mikrofazies und Herkunft. – Geologische Blätter für NordostBayern, 47, 321-338 Frankovich, R. (ed., 1992): Archeometria della Ceramica. Problemi di Metodo. – Atti 8 SIMCER – Simposio Internazionale della Ceramica, Rimini, 150 pp., Bologna (Centro Ceramico) Gálan, E., Carretero, M.I., Mayoral, E. (1999): A methodology for locating the original quarries used for constructing historical buildings: application to Málaga Cathedral, Spain. – Engineering Geology, 54, 287-298 Gibson, A., Woods, A. (1990): Prehistoric pottery for the archaeologist. – 293 pp., Leicester (Leicester Univ. Press) Hagn, H. (1985): Angewandte Geologie und Mikropaläontologie in der Keramikforschung. – Razprave IV, Razreda SAZU, 26, 185-198 Herrmann, B. (ed., 1994): Archäometrie. Naturwissenschaftliche Analyse von Sachüberresten. – 212 pp., Berlin (Springer) Herz, N. (1985): Isotopic analysis of marble. – In: Rapp, G.,
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Jr., Gifford, J.A. (eds.): Archaeological Geology. –331351, New Haven (Yale University Press) Herz, N. (1987): Carbon and oxygen isotopic ratios: a data base for classical Greek and Roman marble. – Archaeometry, 29, 35-43 Herz, N. (1992): Provenance determination of Neolithic to classical Mediterranean marbles by stable isotopes. – Archaeometry, 34, 185-194 Herz, N., Garrison, E.G. (1998): Geological methods for archaeology. – 343 pp., New York (Oxford University Press) Herz, N., Waelkens, M. (eds., 1988): Classical marble: Geochemistry, Technology, Trade. – NATO Advanced Science Institute Series, E, Applied Sciences, 153, 481 pp., Dordrecht (Kluwer) Kars, H., Broekman, A. (1981): Early-mediaval Dorstad, an archaeo-petrological study. Part IV: The mortars, the sarcophagi, and other limestone objects. Petrography and provenance of the limestone material. – Berichten Rijksdienst Oudheidkundig Bodemonderzock, 31, 415-151 Klemm, R., Klemm, D.D. (1973): Steine und Steinbrüche im alten Ägypten. – 465 pp., Berlin (Springer) Lamprecht, H.O. (1993): Opus caementitium. Bautechnik der Römer. 4. Auflage. – 264 pp., Düsseldorf (Beton Verlag) Maniatis, Y., Herz, N., Basiakos, Y. (eds., 1995): The study of marble and other stones used in antiquity. – ASMOSIA III Athens. Transact. 3rd Symp. Assoc. Study of Marbles other Stones used in Antiquity, 302 pp., London (Archetype) Margolis, S.V., Showers, W. (1990): Ancient Greek and Roman marble sculpture: Authentication, weathering and provenance determinations. - In: Walsh, J. (ed.): Marble. Art, historical and scientific perspectives on Ancient sculpture. – 283-299, Malibu, Ca. (J. Getty Museum) Matarangas-Varti, M. Matarangas, D. (1994): Lithofacies determinations and geological origin of building stones of the Asklepieion, at Epidavro: their correlation to ancient quarries. – Pact, 45, 15-24 Matthew, A.J., Woods, A.J., Oliver, C. (1991): Spots before eyes: new comparison charts for visual estimation in archaeological material. – In: Middleton, A., Freestone, I. (eds.): Recent developments in ceramic petrology. - British Museum London, Occasional Paper, 81, 211-263 Middleton, A., Freestone, I. (eds., 1991): Recent developments in ceramic petrology. – British Museum London, Occasional Paper, 81, 410 pp. Moens, L., Roos, P., De Rudder, J., De Paepe, P., van Hende, J., Waelkens, M. (1990): Scientific provenance determinations of Ancient white marble sculptures using petrographical, chemical and isotopic data. - In: Walsh, J. (ed.): Marble art, historical and scientific perspectives on Ancient sculpture. – 111-124, Malibu, Ca. (J. Getty Museum) Mommsen, H. (1986): Archäometrie. Neuere naturwissenschaftliche Methoden und Erfolge in der Archäologie. – 306 pp., Stuttgart (Teubner Studienbücher) Noll, W. (1991): Alte Keramiken und ihre Pigmente. Studien zu Material und Technologie. – 334 pp., Stuttgart (Schweizerbart) Orton, C., Tyers, P., Vince, A. (1993): Pottery in archaeology. – Cambridge Manuals in Archaeology, 269 pp., Cambridge (University Press) Pollard, A.M. (1998): Archeological reconstructions using stable isotopes. – In: Griffith, H. (ed.): Stable isotopes. Integration of biological, ecological and geochemical processes. – 285-301, Oxford (Bios Scientific Publisher) Pollard, A.M. (ed., 1999): Geoarchaeology; exploration, en-
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vironments, resources. – Geological Society of London, Special Publication, 165, 180 pp. Pollini, J., Herz, N., Polikreti, K., Maniatis, Y. (1998): Parian Lychnites and the Prima Porta Statue: new scientific tests and the symbolic value of the marble. – Journal of Roman Archaeology, 11, 275-284 Rapp, G., Jr., Gifford, J.A. (eds., 1985): Archaeological geology. – 435 pp., New Haven (Yale University Press) Rapp, G., Jr., Gifford, J.A. (1985): A selective bibliography of archaeological geology. – In: Rapp, G., Jr., Gifford, J.A. (eds.): Archaeological geology. – 377-431, New Haven (Yale University Press) Riederer, J. (1987): Archäologie und Chemie. Einblicke in die Vergangenheit. – 276 pp., Berlin (Rathgen-Forschungslabor) Riederer, J. (1994): Echt oder falsch. Schätze der Vergangenheit im Museumslabor. – 332 pp., Berlin (Springer) Rottländer, R.C.A. (1983): Einführung in die naturwissenschaftlichen Methoden in der Archäologie. – Archaeologica Venatoria, 6, 604 pp., Tübingen (Institut für Vorgeschichte Univ.) Schmid, J., Ambühl, M., Decrouez, D., Müller, S., Ramseyer, K. (1999): A quantitative fabric analysis approach to the discrimination of white marbles. – Archaeometry, 41, 239-245 Schneider, G. (1978): Anwendung quantitativer Materialanalysen auf Herkunftsbestimmungen antiker Keramik. – Berliner Beiträge zur Archäometrie, 3, 63-122 Schneider, G. (ed., 1989): Naturwissenschaftliche Kriterien und Verfahren zur Beschreibung von Keramik. – Acta Praehistorica et Archaeologica, 21, 7-39, Berlin Schneider, G. (1995): A short note on project planning and sampling for laboratory analysis of archaeological ceramics. – KVHAA Konferenser, 34, 23-27 Schubert, P. (1986): Petrographical modal analysis - a necessary complement to chemical analysis of ceramic coarse ware. – Archaeometry, 28, 163-178 Thiedig, F., Wappis, E. (2003): Römisches Bauen aus naturwissenschaftlicher Sicht in der Stadt auf dem Magdalensberg. – Carinthia II, 193, 33-128 Tyers, P. (1996): Roman pottery in Britain. – 228 pp., London (Batsford) Waelkens, M., Herz, N., Moens, L. (eds., 1992): Ancient stones: Quarrying, trade and provenance. Interdisciplinary studies on stones and stone technology in Europe and Near East from the Prehistoric to Early Christian period. – Acta Archaeologica Lovaniensia, Monographiae, 4, 295 pp. Wagner, U.(ed.): Mössbauer spectroscopy in Archaeology. – Hyperfine Interactions, 154, Dordrecht (Kluwer) Walsh, J. (ed., 1990): Marble. Art historical and scientific perspectives on Ancient sculpture. – 299 pp., Malibu, Ca. (J. Getty Museum) Wenner, D.B., Herz, N. (1992): Provenance signatures for classical limestones. – In: Whitbread, I.K. (1986): The characterization of argillaceous inclusions in ceramic thin sections. – Archaeometry, 28, 79-88 Williams, D.F. (1983): Petrology of ceramics. - In: Kempe, D.R.C., Harvey, A.P. (eds.): The petrology of archeological artefacts. – 301-329, Oxford (Clarendon Press) Important references dealing with provenance, restoration and conservation studies of ancient limestones and marbles used for works of art can be found in Conference Proc. Assoc. Study Marble Other Stones In Antiquity (ASMOSIA); http://www.asmosia.org Further reading: K209
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20 Adding Some Samples by Axel Munnecke
Due to the work by Erik Flügel, who passed away far too early in 2004, microfacies investigations of carbonate rocks have a long tradition in our institute in Erlangen (Germany). He stimulated several generations of geologists in the field of carbonate sedimentology, and it is a great honor for me to provide a small additional chapter to the present book. The chapter presents examples from a variety of depositional environments and ages. The thin sections have been prepared for the international microfacies courses held in Erlangen regularly since 1974. The chapter has two goals. On one side, thin sections have been chosen which complement the previous chapters, and provide some additional information, on the other, this chapter is intended as some sort of self-test. Before reading the figure captions you might try to study the pictures and build your own interpretation. After reading the book many of the features visible in the following pictures are ‘old friends’, but some are new documenting that carbonate microfacies analysis is a fascinating, living scientific field, and each thin section tells its own story.
Fig. 20.1. The cold-water origin of this modern carbonate sediment (with some siliciclastic SI and carbonate rock fragments CR) is dominated by various types of herbivorous gastropods from the Fucus-Ascophyllum belt (common intertidal brown seaweed species - SW). Remnants of other epifaunal organisms (e.g. balanids - B, serpulids - S) are also typical for this environment. The rock fragments have been derived from the rocky shore. The black, mostly porous components represent remnants of the brown seaweed, which in diagenetically mature carbonate rocks are not preserved. Please note that carbonate deposits from sublittoral kelp forests (Laminaria) have a different appearance. These large seaweed plants use phenolic compounds as chemical defense against herbivores, and as a result, sediments from these environments usually lack herbivorous gastropods. Thin section of a recent beach deposit near Galway, Ireland. B - Balanids, CR - Carbonate rock fragment, SC - Light-gray siliciclastic component, SP- Serpulids, SW - Seaweed
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_20, © Springer-Verlag Berlin Heidelberg 2010
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Fig. 20.2. Thin section of a recent warm-temperate carbonate sediment deposited on a volcanic seamount (Gettysburg Seamount, Atlantic Ocean, W of Gibraltar, 86 m water depth). This highly diverse sample contains fragments of herbivorous gastropods, bryozoans (mainly cheilostome bryozoans), sessile benthonic foraminifera (Miniacina miniacea), serpulids, bivalves, echinoids, coralline algae, and volcanic rocks. The coralline alga indicates that at least part of sediment was formed within the photic zone. BI - Bivalves, BR - Bryozoans, CA - Coralline red algae, ECH - Echinoid spine, FOR - Foraminifera (Miniacina), GA Gastropods, SERP - Serpulids, SI - Siliciclastic debris
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Fig. 20.3. Thin section of a recent beach deposit near Tromsø (N-Norway), rich in fragments of coralline algae (Lithothamnium glaciale), balanid plates, bivalves, gastropods, echinoderms (sea urchins), serpulids, bryozoans, and siliciclastic rock fragments. Beach deposits very often represent a mixture of different littoral environments. High abundance of organisms needing a hard substrate for their metamorphosis and growth (balanids, coralline algae) is typical for rocky shores as e.g. in northern Norway. BI - Bivalves, BA - Balanids, BR - Bryozoans, CA Coralline red algae (upper corner: bored), ECH - Echinoid spines, GA Gastropods, SI - Siliciclastic grains
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Fig. 20.4. Archaeocyathid bafflestone from the Early Cambrian of Canada (Jasper National Park). Two different archaeocyathids are present: (1) the irregular form (Archaeocyathus sp.) and (2) the more regular form in the center (?Mackenziecyathus). The matrix contains trilobite fragments. Quartz grains are extremely abundant indicating a close vicinity to a landmass. Archaeocyaths (now assigned to sponges) had a calcareous skeleton and represent the first metazoan reef builders in Earth History. ARCH - Archaeocyathus, MAC - Mackenciecyathus, SHELL Tiny shelly material indet., TRIL - Trilobite
Fig. 20.5. Poorly sorted, Late Ordovician reef debris limestone (SMF 5) with large fragments of a labechiid stromatoporoid (?Pachystylostroma) with two growth increments, and a heliolitid tabulate coral (?Acidolites). High initial intraskeletal porosity of the stromatoporoid. Bioclasts are abundant, mostly consisting of debris of cryptostomate (?Pachydictya) and trepostomate (Hallopora sp.) bryozoans, but also of trilobites, ostracods, and crinoids. Cemented by clear, blocky calcite spar. Micrite is absent. Crinoid fragments show syntaxial cement growth. Stromatoporoids, tabulate corals, bryozoans and crinoids represent most important carbonate-producing organisms especially in reefs and reef-related environments from the Late Ordovician to the Late Devonian. Late Ordovician: Anticosti Island, Canada. CR-BR - Cryptostomate bryozoan, HE - Heliolitid coral, STR Stromatoporoid, TR-BR - Treptostomate bryozoan, TRILTrilobite, CRI - Crinoids
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Fig. 20.6. Parautochthonous debris from a small coraldominated patch reef, with large solitary rugose corals (Streptelasma sp.), and abundant gastropods of varying size, fragments of crinoids, trilobites, and ostracods forming a rudstone with packstone matrix. Geopetally filled cavities inside and under fossils are common and exhibit graded infilling. After rapid deposition of the sediment the cavities were filled with suspension which settled according to grain size. The primary calcitic organisms (corals, brachiopods, crinoids, ostrocods) are well preserved whereas the originally aragonitic gastropod shells have been completely replaced by clear blocky calcite spar. Shell abrasion-derived material in the matrix. Abundant tiny quartz grains indicate a nearby land mass. Late Ordovician: Anticosti, Canada. BRACH - Brachiopod shell (note the geopetal underneath), GA - Gastropod. RUG-C - Rugose corals
Fig. 20.7. Poorly sorted reef limestone (SMF 5) from a stromatoporoid-dominated biostrome rich in fragments of stromatoporoids and crinoid ossicles. In addition, brachiopods and bryozoans are present. The sediment is poorly washed, geopetal structures are frequent. The components are not abradedand delicate particles such as fenestrate bryozoans are not destroyed. This, in combination to the poor sorting and the high amount of finegrained material, points to a very short transport distance. Silurian (Ludlow): Gotland, Sweden. BRACH - Brachiopod, BRY - Bryozoans, encrusting, F .-BRY - Fenestrate bryozoans, CRIN - Crinoids, STR - Stromatoporoids, STYLO - Stylolithic contact underneath the stromatoporoid colony
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Fig. 20.8. Rudstone with reworked oncoids, intraclasts, and bioclasts. Bioclasts are represented by fragments of crinoid stems, tabulate corals, bryozoans, and brachiopods. Different types and stages of encrustations are present. Some of the crinoids exhibit only a very thin crust, whereas the large oncoid with the fragment of a ramiform tabulate coral in the center of the picture shows a thick crust formed by concentrically arranged faint cyanobacterial layers. The small components in the grainstone matrix are mostly crinoid fragments and peloids, the latter mostly resulting from complete micritization of small bioclasts. In this sample, components of a backreef facies (oncoids) are mixed with remnants from organisms which lived in reef-related environments (crinoids, ramiform tabulate corals). Silurian (Ludlow): Gotland, Sweden. BRY - Bryozoans, CRIN - Crinoids, mainly stems, ONC - Oncoids, TAB COR - Tabulate coral
Fig. 20.9. Poorly sorted, Late Cretaceous reef debris limestone (SMF 5) with large fragments of (1) of a radiolitid rudist (reef-building bivalves) with typical meshwork and (2) of an oyster (upper right corner) with laminar and vesicular shell layers both (rudstone with packstone matrix). Rudists exhibit a fast growth rate because they probably housed algal symbionts, and formed abundant reef-like structures in the Tethyan realm. The shell is thick, but porous, and has an outer calcitic and an inner aragonitic (usually recrystallized) layer. The poorly washed matrix contains abundant fragments of bivalves and echinoderms, benthonic foraminifera (e.g. the miliolid Quinqueloculina), and peloids. Late Cretaceous, Coniacian: Lattengebirge, Northern Calcareous Alps. FOR - Miliolid foraminifera
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Fig. 20.10. Bioclastic layer (packstone) rich in fragments of trilobites, brachiopods, bryozoans, ostracods, crinoids, and black pebbles (small black grains) sharply overlain by a micritic layer (mudstone). The bioclastic layer was not yet lithified when the overlying sediment was deposited as the two layers have been slightly mixed by bioturbation. Some components are completely recrystallized due to their originally aragonitic composition. The large intraclast with a similar lithology as the overlying mudstone layer contains a sharply truncated trilobite fragment indicating lithification of the sediment before reworking. The dark, pyrite-rich upper surface of this clast indicates that it represents a reworked piece of a hardground (cf. Fig. 20.11). This sample represents a storm deposit where both fragments of various biogenic components and a hardground have been reworked. Silurian (Ludlow): Gotland, Sweden. BRACH - Brachiopod, CLAST - Two types of clasts: Originally aragonitic components and micritic clasts, BIOTURB Bioturbation, BRY - on top: Bryozoa (Ptilodictya), on bottom bryozoa (?Helopora), TRIL - Trilobite
Fig. 20.11. Hardground on top of the Laframboise Member (Late Ordovician: Anticosti Island, Canada). The irregular surface and the impregnationwith iron are typical features of hardgrounds. The lower layer represents a floatstone with various components surrounded by oncolithic crusts. The centers are bryozoans, gastropods, and some recrystallized components (probably remnants of algae). At the hardground surface components are truncated. The overlying sediment shows grading with larger components deposited in the depressions of the hardground and finer-grained, quartz-rich material above. Dark lense-like structures in the overlying sediment represent filled burrows. 1 Truncated component, 2 - Truncated oncoid, 3Bryozoan-Wetheredella oncoid, 4 - Originally aragonitic clast (alga?), 5 - Larger components (crinoids) in a depression, 6 - Gastropod
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Fig. 20.12. Parautochthonous coral rubble on a Miocene escarpment showing the deep-water coral Desmophyllum cristagalli and large gastropods. Due to the extremely fine-grained matrix (with rare vugs), the corals and the gastropods still exhibit their primary aragonitic shell structure. Although the large components apparently are floating in a wackestone matrix (floatstone) the fabric is grain-supported (rudstone) because the coral exhibits a branched growth form with only few grain contacts when accumulated as debris. In the matrix, fragments of echinoderms (sea urchins and brittle stars), benthonic and planktonic foraminifers, serpulids and bryozoans (encrusting the corals) are present. The different orientation of the geopetal structures indicates multiple reworking, probably resulting from seismic activity in this area (Calabria, southern Italy). Due to the narrow shelves and the absence of a pronounced oceanic shelf front in the Mediterranean area pelagic components (e.g. planktic foraminifera) are often associated with nearer-shore deposits. 1 - Desmophyllum cristagalli, 2 - Coral rubble, 3 - Gastropods, 4 - Gastropod, marginal cut, 5 - Bivalve shell
Fig. 20.13. Parautochthonous coral rubble from a late Pleistocene deep-water mound structure in the Florida Strait. The deep-water coral (Lophelia pertusa) is heavily bored by a clinoid sponge but still consists of aragonite. The matrix contains abundant planktonic foraminifera (4) and pteropods (planktonic gastropods - 3). In contrast to Fig. 20.12 all geopetal structures exhibit more or less the same orientation. Some geopetals show inverse grading (coarsening upward) resulting from microbial precipitation of peloids. Scleractinian corals with large calyces - a common feature of corals not living in symbiosis with zooxanthellae - in association with planktonic organisms (indicating pelagic environments) are typical for deep-water reef deposits. 1 - Lophelia, 2 - Sponge borings, 3 - Pteropod shells, 4 Planktonic foraminifera
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1 cm
Fig. 20.14. Poorly cemented, highly porous coquina with interparticle and moldic porosity. Note different stages of preservation of the originally aragonitic bivalve shells. Some shells are completely dissolved and their outlines are only visible because of their micritic envelopes. Others remain nearly unaltered, still showing their original aragonitic skeletal microstructure. Most of the shells, however, have been transformed to calcite spar, but part of the former internal structures are still visible in form of ‘ghost structures‘ (see below). This type of transformation is termed ‘polymorphic transformation’, and the resulting spar is named ‘neomorphic spar’. The alteration here took place in the meteoric environment. Geopetal structures are represented by an infill of peloids. Pliocene: Roman quarry, Mallorca, Spain. 1- Preserved aragonitic bivalve shells, lamellae still visible, 2 - Shells, tangential cut, 3 - Completly dissolved shell (moldic porosity), in places partly refilled with cement, 4 Cementation of pore space, 5 - Cement-encrusted peloids
2 mm
Fig. 20.15. Detail from the same outcrop. The primarily laminated structure of the aragonitic bivalve shell can be recognized by the ghost-like lamination running through the freshly grown calcite crystals. Part of the shell remains empty. The former outline of the shell is indicated by a thin overgrowth of rim cement. Stabilization of peloidal grains by rim cements. Locality as above.
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0.5 cm
Fig. 20.16. Bivalve and brachiopod shells embedded in a fine-grained peloidal matrix. The sharp outlines of the bivalve shells indicate that dissolution of the aragonitic shell material occurred after cementation of the matrix. Interestingly, the bivalve shells are partly filled with fine-grained sediment. This indicates that after aragonite dissolution fine-grained material was swept into the pore network filling the space of the dissolved shells. This must have taken place very early in diagenetic history of the sediment when soft mud was still available. Mixing of fine-grained and very fine-grained sediment indicates burrowing activity. Silurian (Wenlock): Gotland, Sweden. 1 - Brachiopod shell, 2 -Bivalve shell, partly filled with micritic material after dissolution of the aragonitic shell
References Roberts, J.M., Wheeler, A., Freiwald, A., Cairns, S. (2009): Cold-Water Corals. – Cambridge Univ. Press, 334 pp., Cambridge Titschack, J., Radke, U., Freiwald, A. (2009): Dating and characterization of polymorphic transformation of aragonite to calcite in Pleistocene bivalves from Rhodes (Greece) by combined shell microstructure, stable isotope,
and electron spin resonance study. – J. Sed. Research 79, 332-346 Vennin, E., Aretz, M., Boulvain, F., Munnecke, A. (eds.) (2007): Facies from Palaeozoic Reefs and Bioaccumulations. – Mémoires du Muséum national d’Histoire naturelle, 195, 341pp.
APPENDIX This appendix contains the following items: 1. Answers to the exercises found in the text. 2. Notes for all important details found on the updated CD attached to the inside of the back cover. 3. Permission for the use of figures from other journals upon request of the editors.
1 Answers to Exercises Chapter 8 Pl. 46/7 on p. 358: Limestone classification Folk classification. Poorly sorted algal biopelsparite. Dunham classification. Poorly sorted peloid-skeletal grainstone. These names do not reflect the whole spectrum of the criteria that are important in the environmental interpretation of the sample. Criteria include the striking bimodal sorting indicative of mixing of sediment (note the infilling of peloids in the saucershaped shell), the occurrence of peripherally micritized dasyclad green algae and some aggregate grains. Other criteria not considered by the names are the low-diversity composition of the algal assemblage (all sections show just one species of Diplopora) and the patchy distribution of inter-grain cavities filled with sparite, indicating strong burrowing of the sediment. Some of these characteristics can be included within the names, e.g. sediment composition (intrabiopelsparite, if you use the increasing mode of grain syllables) or peloiddasyclad grainstone. But it will be difficult to summarize all environmentally significant features in the name of the rock sample. The rock name serves as an initial categorization, but does not reflect all the criteria that define a meaningful microfacies type. Chapter 10 Fig. 10. 62. on page 574: Which fossils are shown in the figure? A. Echinoderm fragment. Why? Diagnostic criteria are the extremely fine meshwork, the dusty appearance and parallel cleavage planes (Sect. 10.2.4.9). The skeletal grain is enveloped by a micrite rim produced by the infill of microborings. Therefore, the grain is a cortoid. Late Triassic: Gosaukamm, Austria.
B. Echinoid spine. Why? The cross section shows the characteristic symmetrical radial pattern around a hollow center (see Pl. 96). Note the same dusty appearance of the element as in A. Late Triassic: Gosaukamm, Austria. C. Serpulid tubes. Why? Characteristic features are curved, unpartitioned tubes forming a gregarious ‘colony’. The fine-grained calcitic walls appear dark in transmitted light. Differences in thickness are due to the double-layer microstructure of the wall (Sect. 10.2.4.6; Pl. 92 ). Late Triassic: Gosaukamm, Austria. D. Bryozoan. Longitudinal and cross section. Why? Typical is the fanlike pattern resulting from differences in the arrangement of tubes, and in the morphology and size of the tubes of the inner axial zone and near the surface (Sect. 10.2.3.4, Pl. 85). Late Tertiary (Miocene): Eggenburg, Austria. E. Coralline calcareous sponge. Why? The oblique section exhibits the characteristic central cavity surrounded by an outer zone consisting of irregular and differently shaped, micrite-filled canals (Sect. 10.2.3.1). Uppermost Late Carboniferous: Carnic Alps, Austria. F. Tabulate coral. Why? At first sight the section is reminiscent of a bryozoan, but the tubes are subdivided by horizontal partitions characteristic of tabulate corals and an internal differentiation of the colony (see D) is absent. Middle Devonian: Carnic Alps, Italy. G. Trilobite. Why? The section shows the characteristic ‘shepherd’s crook’ appearance of the curved carapace segment. Middle Cambrian: Öland Island, Sweden. H. Dasyclad green alga. Why? Diagnostic criteria include the segmentation of the algal body, and the regular pattern of the branches (pores) in whorls around the central axis (Sect. 10.2.1.7). Early Tertiary (Paleocene): Kammbühel, Austria. I. Radiolaria. Why? Diagnostic criteria are the very small size, the spherical shape of the test and a surface
E. Flügel, Microfacies of Carbonate Rocks, 2nd ed., DOI 10.1007/978-3-642-03796-2_BM2, © Springer-Verlag Berlin Heidelberg 2010
Appendix
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zone characterized by pores (Sect. 10.2.2.2). Jurassic: Bavarian Alps, Germany. J. Dasyclad algae. Why? The section is characterized by a regular pattern similar to that in G, caused by pores arranged in whorls. Note the differences in the size of the pores and the branching near the periphery (Sect. 10.2.1.7). Mid-Cretaceous: Subsurface, Arabia. K. Octocoral sclerites. Why? Needle-like elements are known from sponges, holothurian echinoderms and octocorals. The elements shown in the figure are octocoral sclerites, as indicated by the spine-like projections. Cretaceous: Northern Alps, Austria. L. Benthic foraminifera. Why? The lepidocyclinid test (center) displays the taxonomically important central initial chamber surrounded by densely spaced median chambers. The smaller lenticular tests are discocyclinid foraminifera exhibiting the characteristic median layer of chambers (see Pl. 74). Tertiary (Oligocene): Western Greece. M. Dasyclad algae. Why? Some dasyclads are characterized by a wide main axis and only thin walls with irregularly arranged pores (see Box 10.3 and Pl. 63). Early Cretaceous: Turkey. N. Tentaculitid shell. Why? Diagnostic features are the conical shape and the ringed surface of the shell (Sect. 10.2.4.5). Middle Devonian: Praha, Czech Republic. O. Corals. Why? The thamnasteroid colony consists of corallites whose centers lack definite walls, but are closely united by confluent septa (Sect. 10.2.3.3, Pl. 84). Jurassic: Northern Calcareous Alps. P. Ammonite shell. Why? The parallel section shows the planispirally coiled shell whose whorls increase in size (Sect. 10.2.4.4). Late Triassic: Gosaukamm, Austria. Q. Brachiopod shell. Why? The section displays the characteristic foliated lamellar structure that distinguishes the shell from bivalve shells. Another diagnostic feature are small (black) pores within the shell (Sect. 10.2.4.1). Early Permian: Australia. R. Bivalve shell or ostracod shell? Why? The section of the recrystallized, double-valved shell shows an overlap of one valve by the other, which would normally correspond to a characteristic feature of ostracod shells (duplicature; Sect. 10.2.4.7). On the other hand, the size of the section argues against this and suggests an interpretation as a small bivalve. A decision can not be made, because the diagnostic microstructure of the wall is not preserved. Uppermost Late Carboniferous: Carnic Alps, Austria. S. Larger Foraminifera. Why? Fusulinids. Because the sections do not cut through the initial chamber and
the geometrical arrangement of the following chambers, further determination is not possible. However, the whorls subdivided by septa identify the fossils as fusulinids (see Pl. 68). Early Permian: Carnic Alps, Austria. T. Planktonic foraminifera. Why? Characteristic features are coiled tests with globular chambers. The larger section is Orbulina typified by a spherical test with a perforated wall (see Pl. 73). Late Tertiary (Miocene): Apulia, Southern Italy. U. Coralline red alga. Why? Although the very fine network of corallinacean red algae can hardly be seen, an identification as coralline alga is reasonable because of the conspicuous spar-filled cavities (conceptacles) arranged in lines (see Sect. 10.2.1.2; Pl. 54). Late Cretaceous: Eastern Desert, Egypt. Chapter 11 Fig. 11.5 on p. 586: General exercise. Bottom/top structures. The sedimentary top is at the right of the photograph, in a northeast direction. This is indicated by the deposition of grains on the long flat, mollusk shells. Grains are represented by bioclasts and aggregate grains. The latter have a different size and are rich in micrite (lumps). Fossils are dasyclad green algae (distributed all over the thin section), benthic foraminifera (top right), bivalve shells (center) and gastropods (bottom right). The skeletal grains are slightly transported and admixed with aggregate grains. Poor preservation of dasyclads, bivalves and foraminifera indicate a primary aragonitic mineralogy. The foraminiferal test (top right) can be identified as Aulotortus sinuosus Weynschenk, a species indicative of an upper Late Triassic age. Environment. The occurrence of abundant dasyclads indicates a warm, very shallow, well illuminated depositional environment and a mobile and unstable substrate with redistribution of grains by waves or currents. The sample exhibits criteria of the Standard Microfacies Types SMF 17 (aggregate-grain grainstone) and SMF 18-DASY (bioclastic grainstone with abundant dasyclads). Both types are characteristic of sediments deposited in shallow shelf lagoons. Porosity. The originally high interparticle porosity is occluded by marine cements. Differences with Pl. 100 and Pl. 101. Mutual features between the sample shown in this figure and the samples shown in Pl. 100 and Pl. 101 are the depositional texture (grainstone), the occurrence of composite and aggregate grains, the mixture of these grains with skeletal grains and the occurrence of both, dasy-
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Appendix
clad algae and foraminifera. However, dasyclads and foraminifera shown in Pl. 100 and Pl. 101 differ taxonomically from those in Fig. 11.5. Further differences relate to grain size, sorting and cement types. Diagnostic microfacies criteria. The microfacies type should be defined by the specific dasyclad assemblage and the association with aggregate grains. Chapter 14 Fig. 14.32 on p. 724: Standard Microfacies Type. According to the key in Box 14.7 this is a sample which should be attributed to SMF 16-NON-LAMINATED. The thin section exhibits more or less equally sized peloids associated with a circular section of calcispheres. The latter indicate a Middle-Late Devonian or Early Carboniferous age of the sample (see Sect. 10.2.1.9: Pl. 60/1). If you look carefully you will find some smaller foraminifera (e.g. top center; center and lower right) that indicate an Early Carboniferous age. The finegrained peloidal grainstones of the Standard SMF 16NON-LAMINATED is typical of restricted inner ramp and platform interior environments.
2 Looking at the Attatched CD The attached CD contains comparison charts for visual estimation of frequency values in thin sections and a data set of references dealing with ancient and modern carbonates.
2.1 Comparison Charts Comparison charts for visual percentage estimation have been published by Bacelle and Bosellini (1965) and by Matthew et al. (1991). The Bacelle and Bosellini charts (uncompressed TIF format on the enclosed CD) can be enlarged for convenient use. The Bacelle and Bosellini comparison charts compare different grain categories and different percentage values and consider the ‘psychological distortion effect’ which results from looking at white objects on black backgrounds and black subjects on white background (see Fig. 6.10). The Matthew et al. comparison charts have been developed for estimating the frequency of temper grains in ancient ceramic sherds. They can be useful for microfacies studies as well. In contrast to the Bacelle and Bosellini charts, these charts also consider differences in grain sizes and sorting (see Fig. 6.11). Because of the high permission fee for the publication on
these charts can not be included on the CD. For a complete set of charts please see Matthew et al. in Middleton, A. and Freestone, I. (eds., 1991): Recent developments in ceramic petrology. – British Museum Occasional Paper, 81, 211-263.
2.2 List of References This book is based on a data set of references (now enlarged to more than 16,000 in March 2009) selected over the years. This data set which can be found on the CD accompagnying the book also includes all references cited in the book. • The updated list of references is presented as Word (for Mac and Windows) and Rich Text Format documents. • An RTF file containing all keywords is added. • The original data were collected as Bookends Pro data, which can easily be transferred into the last version of Bookends. For the software see Bookends 7.x for Mac (
[email protected]). • A data set in Endnote X1 (for Mac and Windows) is added. • Versions in Civati for Windows and BibTeX for Mac, Windows and Unix are included. All the references are cited in the following manner: Author – Titel – Editors – Journal – Volume – Pages – Date – Publisher – Location – Type (Journal, book, etc.) – Abstract (not used)– Keywords – Notes (not used). Practical use: Each reference includes one or several keywords designated by a code number K. You will find code numbers under ‘Further reading’ after the basic references (‘Basics’) included at the end of chapters and/or sections. Use these K-codes to produce a list of references which is relevant for your specific field of interest. For example K033 and K078 will be useful in gathering papers dealing with matrix and micrite, K200 includes a list of papers important for understanding relations between carbonate reservoirs and carbonate facies.
3 Permissions The following institutions and publishers have friendly granted permission to use figures from their journals and publications: Ägyptisches Museum Berlin for portrait on Fig. 19.6 American Association of Petroleum Geologists for
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Fig. 4.5 from Folk, R.L. (1959): Practical petrographic classification of limestones. – AAPG Bull. vol. 43, 1-38 Fig. 7.3 from Longman, M.W. (1980): Carbonate diagenetic textures from nearsurface diagenetic environments. – AAPG, Bull., vol. 64, Fig. 1 on p. 463 Figs. 8.4, and 12.1 from Folk, R.L. (1962): Spectral subdivision of limestone types. – AAPG Mem., vol. 1, 62-84. Blackwell Publishing, Oxford for Fig. 4.5b from Munnecke et al. (1997): Microspar development during early marine burial diagenesis... – Sedimentology, vol. 44, detail of Fig. 5 on p. 986 for Fig. 4.24 from Gerdes, G. et al. (1994): Structural diversity of biogenic carbonate particles in microbial mats. – Sedimentology, vol. 41, 1273- 1294 for Fig. 4.26 from Strasser, A. (1986): Ooids in Purbeck limestones (lowermost Cretaceous) of Swiss and French Jura. – Sedimentology, vol. 33, 711-727 for Fig. 6.2 from Tucker, M.E. (1981): Sedimentary petrology: an introduction. – Fig. 2.3 on p. 15 for Fig. 7.23 from Tucker, M.E. (1985): Shallow-marine carbonate facies and facies models. – In: Brenchley, P.J. (ed.): Recent developments and applied aspects. – p. 147-169; 1 Fig. slightly modified for Fig. 12.3 from Tucker, M.E. and Wright, V.P. (1990): Carbonate sedimentology. – Fig. 4.8 on p. 108 for Fig. 14.6 from Lukasik, J.J. et al. (2000): An epeiric ramp low energy, cool-water facies in a Tertiary inland sea, Murray Basin, S. Australia. – Fig. 12 on p. 876 for Fig. 15.26 from Kenter, J.A.M. (1990): Carbonate platform flanks: slope angle and sediment fabric. – Fig. on p. 788 for Figs. 15.30 and 31 fromTucker, M.E. and Wright, V.P. (1990): Carbonate sedimentology. – Fig. 4.8 on p. 108; Figs. 5.32 and 5.33 on p. 255 and p. 256 British Museum London for Fig. 6.11 details taken from Matthew, A.J., Woods, A.J., Oliver, C. (1991): Spots before your eyes: new comparison charts for visual percentage estimation in archaeological material. – In: Middleton, A., Freestone, I. (eds.): Recent developments in ceramic petrology. – British Museum Occasional Paper, vol. 81, 211-263 Cambridge University Press for Figs. on p. 554 and p. 556 from Black, R.M. (1989): The elements of palaeontology. Elsevier Publishing for Fig. 7.22A from Randazzo et al. (1984): Classification and description of dolomitic fabrics. – Sed. Geology, vol. 37, Fig. 8 on p. 181. for Fig. 4.3 from Hillis-Colinvaux, L. (1980): Ecology and taxonomy of Halimed: primary producer of coral reefs. – Advances in Marine Biology, vol. 17, Fig. 2 on p. 6 and Fig. 11 on p. 18 Friedman, G.M. Troy, for Fig. on Pl. 3 from Friedman, G.M. and Sanders, J.E. (1978): Principles of sedimentology. – Fig. 12-26 on p. 372. Geological Society of America for Fig. 7.1 C from Hardie L.A. (1996): Secular variation in seawater chemistry: an explanation for the coupled secular variation in the mineralogies of marine limestones and potash evaporites over the past 600 m.y. – Geology, vol. 24, Fig. 4 on p. 281 Institute of Paleontology, University of Erlangen for all figures published in Facies. Journal of Geology edited by the Institute of Geosciences, Chicago for Fig. 4.15 from Logan, B.W., Rezak, R. and Ginsburg, R.N. (1972): Classification and environmental sig-
Appendix
nificance of algal stromatolites. – Journal of Geology, vol. 72, Fig. 4 on p. 76 for Fig. 6.4 from Folk, R.L. and Robles, R. (1964): Carbonate sands of Isla Perez, Alacran Reef complex, Yucatan, vol. 72, Fig. 9 on p. 267 Kansas Geological Survey, The University of Kansas, Lawrence for Fig 16.1 from Fischer, A.G. (1964): The Lofer cyclothems of the Alpine Triassic. – In: Merriam, D.F. (ed.): Symposium on cyclic sedimentation vol. 169, Fig. 7 on p. 113 Manson Publishing Ltd. for Fig. 7.5 from Adams, A.E., Mackenzie, W.S. (1998): A color atlas of carbonate sediments and rocks under the microscope. – Fig. 197 on p. 156. Schlager, W., Amsterdam, for Fig. 14.2 Society for Sedimentary Geology (SEPM) for Fig. 2.11 from Hay, W.W. (1974): Studies in paleo-oceanography. – SEPM Spec. Publ., vol 20, Fig. 2 on p. 3 for Fig. 3.6 from Kidwell, S.M. et al. (1986): Conceptual framework for the analysis and classification of fossil concentrations. – Palaios, vol. 1, Fig. 4 on p. 233 for Fig. 3.7 from Kidwell, S.M. & Holland, S.M. (1991): Field descriptions of coarse bioclastic fabrics. – Palaios, vol 6, Fig. 1 on p. 427 for Fig. 4.5a from Folk, R.L. (1965): Recrystallization of ancient limestones. – SEPM Spec. Publ. vol. 13, Fig. 3 on p. 22 for Fig. 6.3 from Longiaru, S. (1987): Visual comparators for estimating the degree of sorting from plane and thin sections. – J. Sed. Petrol., vol. 57, Fig. 4 from p. 793 for Fig. 7.1D from Kiessling, W. (2002): Secular variations in the Phanerozoic reef system. – SEPM Spec. Publ., vol. 72, Fig. 27 on p. 657 for Fig. 7.13 from Demicco, R.V. and Hardie, L.A. (1994): Sedimentary structures and early diagenetic features of shallow marine carbonates. – SEPM Atlas Ser., Fig. 2 on p. 5 for Fig. 7.22B from Sibley, D.F. and Gregg, J.M. (1987): Classification of dolomite rock textures. – J. Sed. Petrol., vol 57, Fig. 2 on p. 969 for Fig. 8.1.C fromWood, R. (1993): Nutrients, predation and the history of reef-building. – Palaios, vol. 8, Fig. 1 on p. 530 for Fig. 14.5 from Gischler, E. and Lomando, A.J. (1999): Recent sedimentary facies of isolated platforms, BelizeYucatan... – J. Sed. Research, vol. A69, Fig. 3 on p. 752 for Fig. 14.26 from Achauer, C.W. (1982): Sabkha anhydrite ... In: Handford, C.R. et al. (1982): Depositional and diagenetic spectra of evaporites- a core workshop. – Fig. 7 on p. 206 for Fig. 14.31 from Spence, H.G. and Tucker, M.E. (1999): Modeling carbonate microfacies in the context of high-frequency dynamic ... – J. Sed. Research, vol. B69, Fig. 2 on p. 948 for Fig. 16.3 from Hinnov, L.A. and Goldhammer, R.K. (1991) Spectral analysis of the Middle Triassic Latemar limestone. – J. Sed. Petrol., vol. 61, Fig. 4A on p. 1177 for Fig. 16.24 from Kiessling et al. (2002): Phanerozoic reef patterns. – SEPM Spec. Publ., vol. 72, Fig. 7 on p. 634 Springer Verlag for figures published in Springer books and journals. Taylor and Francis AS (www.tandf.no/leth) for Figs. 16.20, 16.21, 16.22 from Kiessling et al. (2003): Patterns of carbonate platform sedimentation. – Lethaia, 36, p. 212-213. Thieme Verlag for Figs. 10. 14 and 10.15 from Berger, S. and Kaever, M.J. (1991): Dasycladales. – Fig. 2.1, 2.2, 2.3 and 2.6
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Index
Note: Pages in bold: Headings Pages in italics: Text of text-figures and tables Pages underlined: Text on plate descriptions Pages marked with an asterisk contain definitions used in the chapters 4 and 5.
A Aachen, Germany 410 Aalenian (Jurassic) 757, 872 abrasion 75, 90, 91, 92, 95, 97, 104, 174, 211, 232, 510, 523, 667, 696, 717, 834, 875, 897 Abu Dhabi 25, 528, 708, 745 abundance zone 570 abyssal 9, 46, 48, 454, 486, 858 Acanthocircus 484 Acanthophyllum 384 acanthurid surgeonfish 387 Acastea diaphorogona 483 accommodation space 756, 816, 817, 818, 827 accretionary buildup 524 accretionary organomicrite *83, 99 accretionary slope 48 accumulation 24, 50, 56, 60, 80, 354, 595, 606, 842 rate 211, 483, 659, 870 Acervulina 123 acervulinid foraminifera 137, 386, 476, 886 Acetabularia 430, 435, 444 acetate replica 65, 832 Acicularia 430, 435 Acicularia cf. pavantina 444 Aciculella 444 acid organic macromolecule 81, 82 Acidolithes 918 acid residue 63, 107, 262, 558, 559, 848, 850 acinetes 742 acritarchs 101 Acropora 249, 342 Acrothoracica 389 acrylate 900 actaeonellid gastropods 530, 865 actinomycetes 120 Actinostroma papillosum 500 Actinotubella 562 Adnet limestone 198, 212, 230, 796, 797; Plate 139 on p. 798 Adnet, Salzburg, Austria 198, 212, 218, 230, 343, 352, 358, 546, 554, 556, 797, 798 ADP (Aragonite-Dominated Precursors) 99, 274 Aegean Sea 914 Aeolisaccus 559, 562 aerobic 613 Africa 11, 14, 20, 41, 379, 623, 849, 892
Agadir, Morocco 728 Agaricia 601 Agathammina austroalpina 678 Agathazell, Bavaria, Germany 182, 478 Agelas cf. conifera 491 agglomerate 230 agglutinated 454, 540, 635 microbialite 374; Plate 50 on p. 375 stromatolite 374 thrombolite 372 aggradation 7, 12, 660, 756 aggrading neomorphism 94, 95, 96, 303, 322 aggregate grain 37, 38, 102, 116, 146, *163, 164, 166, 176, 248, 356, 438, 578, 582, 584, 590, 612, 626, 628, 636, 680, 700, 702, 713, 714, 717, 749, 757, 762, 785, 810, 813, 821, 848, 850, 852, 888; Plate 15 on p. 165 algal 164 Bahamian 166 formation 163, 164 significance 166 terminology 163 ahermatypic corals 41 Alacran reef, Mexico 249 Alamogordo (New Mexico) 675, 700, 842 Albacete, Alicante, Spain 284 Alberta, Canada 877, 879 Albian (Cretaceous) 150, 304, 414, 418, 468, 702, 826, 871 albite 622, 644 Alboran Block 849, 850 alcyonarians 41, 346 Alexandria, Egypt 914 algae 14, 15, 16, 20, 67, 96, 119, 120, 121, 124, 128, 132, 192, 222, 314, 342, 343, 346, 369, 376, 381, 382, 384, 385, 386, 387, 394, 395, 404, 422, 436, 455, 458, 559, 562, 563, 564, 584, 602, 610, 611, 612, 616, 625, 626, 636, 637, 638, 656, 663, 671, 684, 688, 700, 704, 706, 713, 714, 717, 719, 727, 738, 755, 762, 807, 812, 820, 830, 831, 836, 842, 848, 850, 865, 869, 873, 886, 898, 921 brown 427, 604 endolithic 119, 738 endosymbiontic 512, 616 epilithic 119 halimediform algae 88 photosymbiontic 611, 612, 873 planktonic 556, 738, 793 significance 406 algal aggregate grains 163 algal assemblage 132, 406, 438, 444, 570, 637 algal bioclasts 290 algal bioherms 20 algal blades 34, 382, 427, 428 algal clast 708 algal cup reefs 416 algal cyst 4, 436, 450, 452, 742
930
algal debris 38, 118, 414, 824 algal filaments 113, 119, 308, 344, 620 algal intraclast 167 algal laminite 364 algal mat 25, 75, 82, 122, 125, 192, 302, 704, 706, 708, 747, 749 algal meadow 432, 460 algal mound, reef 21, 385, 739, 786 algal needle 78 algal nodule 21, 414 algal peloid 113 algal pisolite 157 algal ridge 416 algal spores 170, 452, 764 Algeria 638 Alizarin Red S 53, 333, 740, 900 alkalinity 369, 504, 619, 620, 838 Allgäu, Germany 354, 464 allochemic micrite 340, 366 allochemical rock 356, 357 allochems 312, 356, 362 allochthonous 117, 181, 476, 482, 551, 653, 686, 698, 712, 717, 723, 770, 773, 775, 777, 780, 786, 793, 796, 807, 885 accumulation 465 carbonate 264, 340, 578 limestone 348, 350, 354, 366, 578 allocycle 804, 813 allodapic limestone 774, 775 allomicrite 67, *75, 76, 78, 90, 93, 268, 578, 589, 655, 716, 717, 837, 838, 842, 875, 888 Alpe Ra Stua, Cortina d’Ampezzo, Italy 487 Alpine-Mediterranean region 198, 200, 220, 221, 432 498, 532, 534, 553 Alpinophragmium 454, 835 Alpinophragmium perforatum 678 Alpinothalamia slovenica 498 Altensteig 1 well, Southern Germany 466 alteration diagenetic 520 metamorphic 336 Altopiano di Asiago, Italy 218 alveolar structure 726, 727, 818, 822 Alveolina 108, 454, 476, 478 Alveolina lepidula 478 alveolinids 230, 459, 472, 478, 481, 632; Plate 74 on p. 476; Plate 75 on p. 478 Alveolites 510, 673 Alveosepta 454, 465, 466, 481 Amathus, Cyprus 908 amino acid 82, 102 Ammobaculites 454, 466 ammonification 83, 369, 504 ammonites 4, 78, 91, 137, 214, 221, 350, 522, 533, 534, 662, 757, 759, 795, 798, 801, 822, 827, 828, 850; Plate 90 on p. 535 Ammonitico rosso 200, 212, 215, 524, 796 Ammovertella 382 amoeboid mosaic 74 amphibians 558, 738 Amphipora 500, 673, 675, 750, 752 Amphiroa sp. 444 Amphistegina 304, 478, 480, 886
Index
AmStrat-CanStrat system 60 Amtmanngalgen, Gesäuse, Austria 438 anaerobic 56, 58, 613, 615, 620, 647, 858 bacteria 51 biofacies 613 Analamanga, Madagaskar 366 ancestral red algae 405, 406, 412, 413, 420, 422, 428; Plate 56 on l. 421 Anchicodium 427 Andros Island 25, 88, 745 angle of repose 788 angularity 864, 911 Angulocellularia 409 Angulodiscus permodiscoides 464 anhedral 74, 96 Anhui, southern China 322 anhydrite 53, 66, 277, 292, 330, 622, 644, 647, 648, 662, 663, 708, 716, 720, 749, 766, 806, 881, 882, 884, 889 chicken-wire 648, 708, 749 Anisian (Triassic) 164, 358, 385, 389, 393, 516, 520, 563 ankerite 271, 292 annelid tubes 540, 542 anoxic 215, 301, 408, 613, 614, 653, 656, 659, 662, 708, 756, 892 Antarctica 42, 366, 564 Antarcticeramus rabotensis 68 Anticosti Island, Canada 919 Anthozoa, anthozoans 509 Anthracoporella 108, 430, 444 Anthracoporella spectabilis 432 Anti-Atlas, Morocco 68, 178, 186, 390, 538, 838 Antony and Cleopatra 914 AOM 81, 82 Aoujgalia 420 Apennines, Italy 150, 440, 442, 466, 474, 542, 553, 570, 571, 684, 698, 819 aphotic 44, 51, 379, 394, 396, 611, 635, 636, 637, 638, 754, 842, 873, 874 Apterinella 460 Aptian (Cretaceous) 424, 435, 427, 446, 465, 468, 470, 524, 526, 831, 860 aptychi 533, 534, 536, 797 Aptychus limestone 226, 484, 534, 536, 797 Apuseni Mts., Romania 571 Arab Formation 877, 880 Arabian Gulf 24, 80, 594, 708, 745 Arabicodium 628 aragonite 13, 20, 21, 24, 37, 41, 43, 66, 68, 70, 83, 99, 119, 120, 206, 267, 268, 270, 272, 274, 290, 292, 299, 306, 308, 320, 321, 330, 374, 402, 408, 418, 424, 427, 428, 430, 506, 515, 522, 529, 530, 533, 540, 562, 616, 617, 620, 630, 632, 637, 652, 730, 735, 753, 784, 798, 807, 821, 831, 837, 843, 844, 854, 856, 857, 860, 866, 867, 868, 874, 886, 924 botryoid 296 compensation 48, 270 dominated 99, 272, 274, 717 lysocline 88, 270 needles 78, 80, 88, 89, 94, 428, 736 aragonite production 867 aragonite-calcite interval 874
Index
aragonite/calcite ratio 616 aragonitic skeleton 420, 922 arborescent thrombolites 372 Archaebacteria 369 archaeocyath reef 378, 389, 409 archaeocyathids 372, 492, 504, 506, 507, 717, 831 918; Plate 82 on p. 505 Archaeocyathus 918 Archaeodictyomitra sixi 483 Archaeogastropoda 529 Archaeolithophyllacean red algae 412, 413, 422 Archaeolithophyllum 135, 139, 420, 422, 427 Archaeolithophyllum lamellosum 420 Archaeolithoporella 139, 241, 303, 346, 379, 428, 560, 562, 563, 564, 569, 781, 844 Archaeolithoporella hidensis 564 Archaeolithothamnium 414, 444 Archaeolithothamnium gosaviense 385 archaeology 903 archaeometry 908 Archean 739, 868 Arctic 42, 44, 496, 450 area counting 247, 254, 255, 256, 258 areal cover 831, 832 Arenigian (Ordovician) 212, 564, 644 arenite 865, 866 bioclastic hybrid 865, 866 calcilithic hybrid 865, 866 Argelita, Valencia, Spain 644 Argentina 352, 530, 541, 644 argillaceous limestone 53, 614, 773, 806 Argustrea quadrigemina 510 Argyrotopos, Epirus, Greece 476 arid 622, 642, 644, 659, 660, 661, 663, 700, 704, 708, 745, 747, 749, 752, 806, 815, 816, 818, 863, 881, 882 array calcite 297 arthropods 51, 111, 188, 542, 611, 650, 717, 726 articulate brachiopods 519 articulated coralline algae 412, 413, 414 Artinskian (Permian) 566, 872 aseismic volcanic ridge 48 ash 911 Ashgillian (Ordovician) 554, 872 Asmari limestone 884 Asselian (Permian) 140, 420, 428, 436, 460, 462, 564, 582, 871 assemblage 468, 482, 628, 674 allochthonous 482 autochthonous 672 cold-water 483 depth-sensitive 675 parautochthonous 482 zone 465, 570 Assilina 181, 182, 476, 478 association pattern 62, 396, 442, 459, 465, 570, 579 Assuan, Egypt 710 Asteriginacea 476 Asteroidea 550 Asterozoa 550, 551 astraeoid corals 512 Astrosclera 496 Astrostylopsis 502
931
ataxophragmiacean foraminifera 465, 470, 472 Atdabanian (Cambrian) 507 Atlantic 40, 42, 44, 138, 386, 506, 643 Atlas Mountains, Morocco 68, 466, 534, 700, 706, 757, 760 atoll 276, 332, 374, 623, 660, 663, 667, 878 Atractyliopsis carnica 436 Auerberg ceramic 908, 912 Auernig, Carnic Alps, Austria/Italy 546 auloporid corals 344, 510, 842 Aulotortus communis 678 Aulotortus sinuosus 678 Australia 2, 12, 14, 20, 21, 24, 25, 40, 102, 372, 376, 507, 515, 618, 628, 664, 669, 806, 863 Austria 56, 135, 140, 157, 352, 354, 409, 414, 420, 444, 450, 460, 472, 516, 520, 615, 768, 798, 812 Austrocolomia 454 Austrocolomia marschalli 678 authigenic minerals 97, 214, 637, 641, 643, 644, 800, 857; Plate 109 on p. 645 autobreccia 857, 858 autochthonous 117, 181, 341, 344, 524, 551, 552, 686, 698, 712, 713, 720, 723, 770, 771, 786, 796 growth 586 input 653 limestone 340, 346, 348, 364, 576, 624 autoclasts 167 autocycle 804, 805, 813 automicrite 67, *75, 76, 78, 82, 93, 96, 118, 290, 371, 497, 504, 563, 578, 589, 655, 716, 786, 787, 788, 830, 837, 838, 842, 874, 875 autotroph 395, 493, 626, 636, 754, 862 avalanche 772, 788 Axothrix 562 azooxanthellate 44, 512, 853 B Ba 70, 857 Ba/Ca ratio 624 Baccanella 374, 559, 562 Baccanella floriformis 566 Baccanella irregularis 563 Baccelle and Bosellini chart 257, 261 Bacinella 130, 382, 385, 386, 560, 562, 566, 568, 694, 762 Bacinella-Lithocodium association 385 back-scattered electron image 66 background sedimentation 49, 598, 774, 778, 779, 782 backreef 38, 442, 452, 459, 460, 464, 510, 664, 674, 678, 686, 688, 694, 702, 714, 755, 831, 834, 844 backstepping 756, 899 bacteria 14, 18, 75, 84, 102, 103, 117, 369, 370, 371 376, 379, 387, 388, 394, 395, 483, 562, 564, 566, 656, 727, 788, 838, 857, 858, 898 autotroph 390 chemolithoautotrophic 81, 857 chemosynthetic 623, 857 denitrifying 369, 370 Fe/Mn bacteria 55, 82, 84 heterotroph 81, 91, 184, 390 nitrifying 370 sulfate-reducing 55, 620, 893
932
bacterial aggregate 16, 371 container 84 degradation 163, 504 mud 16 precipitation 16 Bad Harzburg, Germany 150 Bad Ischl, Austria 218 Baffin Bay, Canada 84 Baffler Guild 346, 831, 832, 833 bafflers 427, 428, 496, 510, 528, 569, 663, 673, 680, 807, 831, 832 bafflestone 341, 342, 343, 344, 366, 436, 510, 528, 576, 624, 637, 663, 690, 723, 739, 833, 835, 837, 848, 908, 918; Plate 41 on p. 345 baffling 63, 89, 344, 346, 436, 563, 578, 606, 624, 747, 754, 830, 836, 838, 856 Bahama Bank 81, 112, 116, 146, 163, 164, 222, 263, 302, 493, 762 Bahama type 252, 700 Bahamas 12, 14, 24, 28, 37, 78, 84, 119, 125, 146, 186, 207, 222, 279, 288, 330, 556, 594, 657, 706, 745, 769, 821, 878; Plate 3 on page 39 Bahamian carbonate slope 50, 117 Bahamian-type aggregate grains 163 bahamite grain 116 Baid area, Oman 400 Bajocian (Jurassic) 366, 440, 510 Bakalovaella elitzae 442 balanids 42, 44, 123, 544, 546, 916, 917; Plate 93 on p. 545 ball-and-pillow structure 640 Baltic Sea 617, 669 Baltic Shield 212, 795 Balya, Turkey 426 bank 426, 428, 464, 478, 507, 593, 663, 667, 669, 698, 719, 774, 878, 880 bank margin 815 bar 660, 702, 770 Barbados 12, 166, 276, 304 Barbafera 559 Barents Sea 40, 42, 566 Bari, Puglia, Italy 468 Barmstein limestone, Jurassic 252, 779, 786, 797 barnacles 34, 41, 42, 44, 108, 270, 366, 384, 544, 546, 630, 632, 667, 853 barnamol grain association 547, 627 baroque dolomite 332 Barrandei Limestone, Devonian 764 Barremian (Cretaceous) 442, 446, 465, 468, 480 Barremian-Early Aptian 468, 470 barrier reef 623, 663 Bartolsfelde, Saxony, Germany 516 baryte 207 basalt 842, 913 base-of-slope 46, 49, 169, 775, 777, 780, 781, 782, 802; Plate 137 on p. 783 basin 26, 29, 50, 200, 313, 370, 480, 486, 497, 537, 538, 540, 551, 552, 560, 592, 612, 617, 637,641, 642, 647, 655, 656, 659, 660, 661, 662, 665, 675, 678, 708, 713, 739, 742, 755, 760, 771, 775, 780, 781, 782, 784, 785, 786, 789, 790, 792, 793, 795, 796, 797, 798, 803, 804, 815, 816, 817, 818, 821, 827, 843, 846,
Index
850, 863, 877, 878, 880, 892 carbonate-evaporite 815 deep-water 662, 794 intrashelf 827 restricted basin 29 shallow basin 29 basin analysis 73, 80, 182, 223, 725 basin margin 802 basin model 884 basin-to-shore transect 660 basinal 465, 496, 551, 594 facies 534, 642 limestones 508 Bathonian (Jurassic) 440 Bathurst model 120 bathyal 9, 46, 48, 186, 454, 457, 486, 514, 635, 858, 873 bathymetric zone 396, 635, 637 bathymetry 394, 486, 635, 659 Bathymodiolus 857 bathypelagic zone 9 Bauneia 502 Bavaria, Germany 396, 432, 468, 766, 768, 906 beach 174, 177, 243, 245, 246, 252, 253, 284, 301, 480, 523, 663, 666, 698, 806, 913, 917 barrier 24 wedge 478 beachrock 12, 24, 37, 38, 74, 93, 118, 167, 169, 197, 274, 277, 290, 301, 304, 308, 352, 542, 748 Béchar Basin, Morocco 637, 638 bed set 717 bed thickness 56, 62, 225, 803, 805, 806 bedding 53, 55, 59, 62, 64, 177, 181, 182, 185. 186, 480, 581, 607, 662, 719, 735, 746, 775, 778, 781, 788, 800, 802, 805, 884, 888, 890, 893, 904, 905 plane 14, 56, 177, 217, 220, 247, 593, 613, 739, 775, 779, 813 structure 58, 184 surface 56, 184, 789 type 59 bedload deposition 184, 787 Bedoulian (Cretaceous) 468 bedset 56 Beedeina 462 Begunjscica Mts., Slovenia 182 belemnite battlefield 536, 537 belemnites 71, 533, 534, 537; Plate 90 on p. 535 Belgium 193, 700 Belize 24, 80, 119, 249, 296, 660, 769 Belmont Fm. 38 Belton, Texas 350 bending, plastic 314 benthic 8, 455, 528, 875 algae 96, 405 community 341, 616 depth zone 9 foraminifera 42, 616, 628, 632, 635 foraminiferal-oxygen index 614 benthopelagic coupling 43 benthos 402, 558, 580 benthos/plankton ratio 757 Berchtegaden, Germany 786 Beresella herminae 436
Index
beresellid algae 435, 436, 438, 507 Bergama, Turkey 426, 432, 556 Bergisches Land, Germany 560 Berlichingen, Germany 524 Bermudas 14, 37, 107, 125, 276, 304, 308, 541, 542 Berriasian (Cretaceous) 432, 440, 446, 483, 487, 541, 544, 547, 766, 900 Berry Islands, Great Bahama Bank 164 beta calcrete 726 Betic Cordillera, Spain 186, 420, 510, 538, 796, 848, 850 Biancone limestone, Cretaceous 93, 484, 797 Bight of Abaco 88, 95, 98 bimineralic skeleton 103 Bimini Island, Bahamas 38 bimodal 61, 248, 251 bimol grain association 627, 670 binary data analysis 263 Binder Guild 380, 478, 831, 832, 833 binders 380, 416, 564, 569, 673, 678, 831, 832 binding 25, 56, 80, 184, 366, 371, 374, 376, 404, 563, 713, 754, 850, 868 bindstone 341, 342, 343, 344, 346, 349, 366, 436, 576, 624, 663, 666, 690, 700, 704, 714, 719, 723, 735, 739, 745, 747, 750, 752, 782, 806, 810, 833, 837, 848, 850; Plate 41 on p. 345 fenestral 366, 596, 706, 824 laminated 704 micritic 366 microbial 858 thrombolitic 366, 578 bio-controlled deposition 102 bio-induced formation 103 bio-retexturing 185, 186, 587 biocalcification 75 biocementstone 118, 342, 343 biochemical conditioning 381 bioclast 36, 40, 44, 60, 61, 78, 94, *100, 101, 107, 108, 163, 174, 175, 222, 257, 268, 292, 296, 304, 358, 360, 362, 366, 368, 406, 416, 438, 484, 515, 518, 523, 524, 530, 578, 596, 606, 630, 644, 649, 865, 682, 686, 694, 696, 700, 702, 717, 718, 755, 759, 772, 777, 785, 787, 795, 820, 852, 865, 921 phosphatic 208, 650 preservation 108 significance 107 sizes of 102 terminology 101 bioclastic 366 debris 524 fabric 59, 61 grains 295 sand 40 biocoenoses 450 bioconstructed limestone 18, 342, 493 biocrystals 125, 270 biodegradation 386, 852 biodiversity 262 bioerosion 4, 20, 40, 81, 88, 90, 91, 95, 97, 104, 170, 205, 211, 220, 303, 308, 379, 384, 385, 386, 387, 388, 389, 390, 393, 394, 395, 396, 410, 416, 523, 596, 623, 624, 667, 696, 716, 769, 770, 853, 854, 856, 875
933
biofabric 61, *181, 201, 480, 523, 554, 581, 590, 610, 673, 718, 888, 904; Plate 18 on p. 183 biofacies 53, 404, 480, 595, 597, 613, 635, 657, 680, 793 biofilm 6, 18, 82, 84, 120, 150, 158, 302, 306, 369, 371, 384, 606, 698, 747, 838, 898 biogenic concentration 60 biogenic encrustation 180, 207, 381, 396; Plate 51 on p. 383 biogenic graded bedding 185 biogenic silica 19 biogenic stratification 184, 211 biogenic structure 53, 59 biogenic tissue 656 biogenicity 374 bioherm 19, 20, 62, 493, 740, 869, 896 biolamination 182, 184, 370, 581, 610 biolithite 342, 348, 364 biological control 1, 135, 184, 340, 873 biologically induced 873 biomarker 73, 656, 858, 859 biomicrite 357, 358, 368 biomineralization 66, 67, 101-104, 270, 370, 873 biomold 283, 837, 900 biomorpha 101 biosparite 357, 578 biosparrudite 357 biostratigraphic marker 189, 468 biostratigraphic zonation 465, 462, 466, 470 biostrome 62, 432, 444, 502, 526, 666, 719, 826, 831, 869 biotic 10 change 605 composition 139, 263, 603, 614, 665 diversity 624, 632, 634 bioturbation 14, 20, 60, 62, 92, 96, 180, 184, *185, 186, 188, 189, 198, 201, 205, 208, 247, 260, 262, 308, 372, 416, 480, 483, 523, 526, 578, 582, 587, 591, 595, 596, 606, 610, 612, 613, 614, 615, 624, 637, 662, 669, 670, 682, 717, 738, 762, 795, 802, 820, 850, 854, 858, 859, 875, 889, 921; Plate 19 on p. 187 classification 188 deep-sea 186 fabric 186, 188 index 186, 188 biozonation 422, 440, 442, 444, 447, 448, 453, 459, 460, 465, 466, 468, 476, 478, 480, 538, 570, 571 biozone 201, 447, 459, 464, 465, 470, 480, 570 Bir Abu El Husein area, southern Egypt 308 birdseyes *190, *192, 196, 197, 371, 636, 704, 706, 747, 748, 749, 824, 889, 904; Plate 20 on p. 191 bitumen 56, 170, 613, 614 bivalve limestone 523 bivalve reef 524 bivalve shell 108, 164, 304, 352, 366, 381, 519, 524, 598, 616 bivalves 4, 18, 19, 20, 34, 40, 42, 44, 51, 60 71, 90, 137, 178, 188, 189, 214, 220, 304, 342, 346, 350, 356, 389, 393, 394, 395, 396, 401, 427, 481, 522, 523, 524, 526, 528, 580, 591, 596, 597, 598, 606, 612, 616, 618, 622, 628, 630, 632, 662, 663, 667, 684, 686, 688, 692, 696,
934
717, 719, 725, 737, 738, 739, 748, 749, 750, 777, 795, 796, 800, 809, 810, 812, 822, 824, 828, 831, 853, 854, 856, 857, 858, 859, 860, 864, 871, 873, 875, 917, 920, 922-924; Plate 87 on p. 525 boring 4, 389 burrowing 614 classification 522 ecology 523, 524 halobiid 528 inoceramid 524 microstructure 522, 524 ostreid 714 pelagic 534, 782 posidoniid 528 rudist 869 teredinid 440 unionid 132 black grains 168, 169 black lithoclasts 168 black pebble 12, 167, *168, 169, 170, 190, 252, 304, 344, 596, 648, 739, 747, 810, 820, 822, 921 Black Sea 51 black shales 559, 613, 614, 656 blackened grains 169 Blastochaetetes 502 blastoids 550, 591 Bleiberg Type 892 block faulting 730, 732, 889 Blue Holes 222 blue-green algae 21, 370, 404, 408, 410, 412, 564 Blumautal, Bavaria, Germany 566 Bogazkale, Turkey 904, 905, 908 Bohinska Bela, Slovenia 462 Bonaire, Netherlands Antilles 330 bonebed 560 bones 558, 559, 560 Bonetocardiella 800 Bonhof, Bavaria, Germany 494 borehole 38, 119, 301, 387, 592, 802, 820 Borelis melo 476 borers 4, 119, 205, 210, 381, 386, 390, 392, 523, 606 borings 66, 84, 96, 121, 126, 131, 139, 178, 185, 207, 208, 210, 212, 302, 371, 380, 381, 386, 387, 390, 394, 395, 396, 410, 414, 524, 526, 556, 580, 605, 606, 607, 608, 610, 649, 662, 712, 766, 819, 820, 831, 854, 898; Plate 52 on p. 391 algae 387 association 386 bivalves 91 demospongid 389 gastropod 393 intensity 394 phoronid 392 sipunculid 392 sponges 604 Bornetella 430, 434 Bornhardtina 750, 752 boron 619, 622 Borzaites atavus 488 Bositra 215, 528 Botomian (Cambrian) 507 botryoid 296, 300,. 306, 417, 843, 844, 858, 860
Index
botryoidal lumps 163, 164 bottom current 92, 96, 180, 210, 484, 486, 534, 581, 595, 646, 682, 798, 856 bottom flow 184, 598 bottom-level association 507 boudinage 200 Boueina hochstetteri 424 Boueina marondei 424 boulder 379, 606, 771, 773, 774, 784, 803, 847, 848, 906 Bouma Model 773, 774, 775, 778, 784 boundary plane 56 boundstone 313, 342, 344, 346, 348, 349, 364, 507, 528, 576, 608, 662, 672, 678, 690, 719, 781, 782, 784, 788, 807, 834, 858 coral/microbial 824 serpulid 542 sponge 896 Bourgogne, France 128 brachialia 520, 553, 554, 556 brachiopod limestone 520U, 522, 637 brachiopod shell beds 522 brachiopod spines 520, 537, 547 brachiopods 41, 59, 71, 74, 178, 181, 182, 196, 198, 207, 208, 220, 260, 342, 362, 404, 427, 459, 518, 519, 524, 576, 579, 580, 596, 597, 606, 614, 616, 617, 618, 628, 630, 636, 650, 654, 662, 666, 675, 682, 686, 692, 696, 717, 719, 750, 752, 754, 757, 759, 764, 768, 777, 789, 790, 812, 822, 842, 853, 857, 869, 875, 919, 920, 921, 924; Plate 86 on p. 521 articulate 519, 520, 522 endopunctate 520 impunctate 520 inarticulate 520 microstructure 519, 520 pseudopunctate 520 punctate 520 shell structure 519 brackish 21, 25, 447, 448, 450, 454, 522, 541, 542, 544, 617-619, 706, 820, 663, 826 Bradyina 460 branchstone 341, 342 Brandenberg, Tyrol, Austria 865 Brazil 44, 136 break in sedimentation 57, 205, 206, 212, 649, 652, 820 break in slope 665 breakage 188, 189, 226, 268, 277, 310, 312, 314, 322, 534, 717 breakdown 80 of carbonate grains 95 of carbonate platform 233 of shells 74 breccia 53, 58, 59, 62, 166, 168, 169, 206, 216, 221 222, 223, *228, 230, 233, 236, 243, 244, 252, 354, 452, 605, 640, 659, 662, 665, 666, 675, 678, 686, 712, 730, 732, 734, 747, 770, 772, 775, 781, 786, 787, 788, 800, 801, 810, 817, 820, 822, 834, 847, 848, 858, 880, 882, 892, 893; Plate 26 on p. 231; Plate 27 on p. 237 autoclastic 238 caliche 234, 235, 727 classification 233 collapse 622, 731, 732, 749
Index
composite 781 depositional 233, 241, 590 dessication 235 diagenetic 233, 239 dilatation 238 fault 238 fissure fill 234 forereef 234, 235, 236, 600, 771, 834 foreslope 784 formation 229, 233 internal 229, 232, 234, 238, 239 intraformational 229, 708, 750, 752 karst 730, 732 lacustrine 236 main types 234 mass-flow 232, 236, 238, 241, 242, 772 megabreccia 660, 731 microbreccia 731 monomict 229 non-depositional 233, 235 pedogenic 232 peritidal and shallow-marine 234, 235, 748, 750 polymict 229, 230, 810 pseudobreccia 234 rockfall 234 sedimentary 663 shatter 223 shear 234, 238 shrinkage 233 significance 242 slope 771, 780 solution-evaporite-collapse 14, 233, 234, 235, 236, 747 stylobreccia 234 tectonic 233, 238, 239 transgressive 750 brecciation 169, 224, 229, 230, 232, 279, 592, 640, 686, 727, 728, 735, 784, 801, 822, 857, 859 Breitenberg, Lake Wolfgang, Austria 798 Brice Canyon, Utah, USA 728 Brilon, Germany 362 British Columbia 493 Brittany 42 brittle shear zone 224 Broeckella belgica 444 brown algae 853 Brushy Canyon Formation 846 bryomol grain association 41, 44, 515, 547, 626, 627, 630, 670, 854 bryonoderm grain association 627, 630 bryooids 122 bryooncoids 122 Bryopsidales 423 bryozoa-foraminifera association 632 bryozoan boring 394 bryozoan debris 476, 516 bryozoan limestone 852, 854 bryozoan sand 515, 630 bryozoans 2, 3, 34, 41, 42, 44, 74, 90, 104, 128, 130, 207, 208, 264, 270, 284, 301, 313, 346, 358, 381, 382, 384, 385, 386, 387, 388, 389, 390, 392, 395, 404, 459, 491, 502, 515, 544, 554, 559, 569, 606, 608, 618, 625, 630, 632, 638, 650, 666, 667, 670, 686, 692, 717, 719,
935
725, 746, 754, 760, 785, 795, 800, 819, 831, 842, 844, 848, 850, 853, 856, 869, 871, 875, 886, 917-922; Plate 85 on p. 517 attached 515 bifoliate 516 carbonate production 515 celleporiform 630 cheilostome 515, 630, 854 classification 515 cryptostomid 516, 608, 918 cyclostome 515, 516, 852, 854, 917 cystoporid 516 diversity 515 ecology 515 fenestrate 108, 515, 516, 608, 675, 842, 843, 844, 919 growth form 515 membraniform 630 microstructure 516 paleoenvironmental significance 515 trepostomid 389, 516, 630, 918 Buchenstein (Livinallongo) Formation, Triassic 782 building stone 260, 314, 336, 528, 725, 895, 903, 906, 908; Plate 151 on p. 909 buildup 3, 6, 18, 42, 196, 428, 465, 480, 481, 658, 666, 688, 717, 757, 770, 830, 831, 878, 880, 882 Bükk Mountains, Hungary 322 bulk analysis 71 bulk sample 59 bulk volume 318 Buntsandstein (Triassic) 374 Burgwik, Gotland Island, Sweden 306 burial 274, 275, 289, 292, 296, 308, 310, 313, 314, 320, 330, 333, 362, 465, 484, 514, 533, 551, 552, 590, 641, 734, 863 condition 644 criteria 301 diagenesis 67, 68, 106, 268, 277, 354, 580, 642, 653, 655, 877, 890 dolomitization 278, 332 model 331 solution 230 stress 288, 313 burning behavior 896 burrow 14, 58, 59, 76, 117, 170, 185, 188, 193, 200, 302, 313, 399, 582, 606, 608, 630, 659, 666, 667, 669, 675, 692, 719, 748, 749, 750, 800, 802, 807, 824, 858, 905, 924 abundance 186 anastomosing 189 crustacean and insect 14 deformation 312 excavation 186 burrow density 682 burrowed limestones, significance 188 burrowing 55, 96, 113, 131, 177, 180, 186, 189, 192, 194, 198, 205, 208, 210, 349, 352, 366, 378, 480, 556, 578, 595, 596, 606, 610, 619, 624, 646, 680, 682, 692, 704, 712, 716, 717, 739, 747, 760, 764, 779, 782, 795, 798, 805, 820, 822, 850, 858, 875, 889, 890, 904 bypass margin 522, 629 bypass slope 48
936
Index
C C-M diagram 247, 253 Ca/Mg ratio 333, 874, 895 Cadosina 452, 450 Caicos platform, Caribbean 146 Calabar Flank, Niger Delta, Nigeria 414 Calabria, Italy 438, 922 Calamophylliopsis flabellum 512 Calcarea 492 calcarenite 14, 37, 38, 243, 246, 248, 252, 257, 592, 596, 678, 723, 746, 774, 780, 781, 800, 802, 846, 896, 898 calcareous algae 37, 71, 80, 88, 105, 140, 178, 180, 196, 208, 233, 241, 249, 252, 400, 401, 402, 404, 405, 406, 418, 447, 458, 491, 570, 571, 582, 584, 593, 604, 612, 635, 637, 666, 692, 719, 738, 746, 748, 749, 752, 760, 770, 842, 850, 864, 875, 888, 904 calcareous dolomite 328 calcareous green algae 78, 444, 624, 626, 628, 632, 634, 635, 636, 671, 680, 702, 717, calcareous plankton 49, 75, 92 calcareous red algae 42, 121, 270, 346, 413, 610, 611, 626, 671, 746 calcareous sponges 491, 492 Calcarina 480 calcidinophycean algae 78, 452 calcification 15, 82, 84, 89, 184, 369, 370, 371, 372, 376, 404, 408, 410, 412, 417, 423, 424, 426, 427, 430, 448, 493, 512, 612, 616, 620, 727, 762, 838 Calcifolium 427 calcilaminite 802 calcilutite 75, 95, 243, 352, 596, 802, 896 calcimicrobes 75, 123, 128, 135, 343, 371, 378, 382, 395, 399, 402, 404, 405, 406, 408, 410, 412, 496, 504, 507, 568, 717, 777, 833 calcimudstone 75, 349, 352, 358, 362, 366, 598, 708 calciodinoflagellates 91, 186, 406, 800 calcirudite 243, 244, 248, 257, 366, 592, 596, 774, 780, 802 calcisiltite 56, 73, *74, *95, 97, 98, 243, 244, 248, 596, 662, 680, 682, 716, 717, 718, 719, 750, 759, 760, 773, 774, 789, 802, 810, 837, 888 Calcisphaerula 452 calcispheres 78, 97, 100, 186, 436, 450, 452, 474, 483, 484, 524, 566, 653, 700, 764, 795, 800, 850; Plate 66 on p. 451 calcite 15, 16, 70, 220, 222, 233, 268, 271, 284, 339, 371, 522, 529, 559, 616, 630, 637, 735, 800, 881, 895 bladed 18 dendritic 13 fascicular-optic 297 length-fast, length slow 13 neomorphic 321, 923 calcite compensation depth 48, 50, 93, 271, 277, 801 calcite lysocline 271 calcite mosaic 274, 278, 299, 310, 346, 534 calcite pseudomorphs 14 calcite spar 73, 325, 358 calcite vein 223, 224, 226, 682; Plate 25 on p. 227 calcite-dominated precursor 99, 274, 632
calcite-dominated time 272, 717 calcite/dolomite/non-carbonate triangle 340, 895 calcitization 95, 145, 326, 333, 484, 497 calciturbidite 62, 63, 168, 248, 459, 570, 686, 688, 698, 776, 779, 780, 781, 782, 785, 788, 789, 790, 792, 794, 807; Plate 138 on p. 791 calcium phosphate 84, 519, 547 calclithite 167 calcrete 2, 11, 12, 13, 14, 206, 352, 376, 726, 727, 728, 731, 734, 806 calcspar 300 Calgary, Canadian Rocky Mountains 608 caliche 11, 12, 21, 74, 93, 97, 120, 157, 158, 162, 167, 185, 190, 192, 232, 235, 299, 352, 361, 364, 564, 663, 698, 725, 726, 728, 742, 813, 827, 828 alpha 726, 728 crust 726, 818 descriptive term 726 intraclasts 16 nodule 726, 728 pisoid 710, 813 profile 11, 726 calichification 726, 749, 806 Callianassa 111, 186, 194 Callovian (Jurassic) 483, 785, 831, 872 Calpionella alpina 487, 488 Calpionella elliptica 488 Calpionella zone 487 calpionellids 91, 97, 105, 402, 453, 483, 484, 487, 490, 570, 662, 684, 795, 797; Plate 76 on p. 485; Plate 77 on p. 489 limestone 92, 490 Calpionellites coronatus 488 Calpionellites darderi 488 Calpionellites major 488 Calpionellopsis oblonga 488 Calyptogena 857 Cambrian 82, 103, 113, 124, 134, 155, 188, 195, 306, 316, 334, 343, 372, 378, 384, 387, 389, 390, 392, 404, 407, 409, 412, 418, 430, 438, 459, 483, 492, 496, 504, 507, 519, 520, 522, 529, 533, 540, 542, 546, 547, 548, 550, 554, 559, 564, 566, 596, 597, 606, 628, 664, 692, 717, 773, 780, 795, 814, 831, 836, 869, 871, 872, 874, 875, 892, 918 Cambro-Ordovician 366, 646, 780 Campanian (Cretaceous) 68, 444, 472, 490, 528 Campbelliella striata 440 Camposauro, Apennines, Italy 474 Canada 40, 630, 860, 880, 892 Canary Island, Atlantic 111 Canberra, Australia 409, 410 Canning Basin, Australia 140, 334, 500, 510 Cantabrian Mountains, Spain 150, 436, 510 capillary pressure 282, 881 capillary system 900 Capitan limestone 222, 223, 844 Capitan Reef, Guadalupe Mts. 379, 836, 843, 844; Plate 145 on p. 845 Capitanian (Permian) 426, 780 Caprinidae, rudists 526 carbon 71, 588, 619, 654, 623 carbon isotopes 37, 71, 654
Index
carbonate allochemical 340 allochthonous 3, 49, 340, 769 bomb 70 budget 4, 369 cycle 803 factory 26, 28, 745, 756, 868, 873 fluorapatite 648 precipitation 6, 7, 16, 18, 80, 132, 370, 458, 873 Carbonate Facies Model 65 carbonate minerals 70, 268, 270, 301, 637 carbonate mud 37, 75, 193, 362, 423 carbonate platform 1, 27, 28, 30, 37, 102, 379, 480, 756, 796 carbonate production 3, 7, 8, 31, 50, 91, 118, 210, 216, 458, 480, 578, 623, 624, 626, 660, 665, 667, 737, 753, 755, 759, 769, 804, 807, 816, 821, 836, 862, 863, 864, 870, 872 carbonate ramp 28, 60, 578, 582 carbonate reservoir 877, 878, 883, 895 carbonate saturation 43, 308 carbonate-siliciclastic environment 54, 862, 863, 864 carbonate/clay ratio 198 carbonatite 801 Carboniferous 63, 114, 118, 134, 136, 154, 181, 193, 196, 296, 322, 364, 372, 378, 384, 412, 417, 420, 422, 428, 432, 436, 438, 442, 444, 448, 450, 452, 458, 459, 460, 462, 502, 506, 507, 510, 516, 519, 520, 530, 533, 541, 546, 551, 552, 560,562, 563, 564, 566, 568, 597, 605, 608, 638, 650, 661, 664, 700, 731, 750, 788, 814, 831, 836, 842, 848, 850, 852, 857, 869, 872, 874, 878, 892, 907 carbonification 560 Caribbean 11, 14, 36, 600, 602, 623, 668, 821 Carinthia, Austria 146, 334 Carixian (Jurassic) 215, 798 Carlsbad Cavern, New Mexico, U.S.A. 306, 710 Carnian (Triassic) 135, 146, 150, 212, 350, 354, 366, 382, 418, 438, 498, 556, 559, 766, 781, 782, 787, 836, 837, 844, 857, 860, 865, 873, 874, 880, 881, 889, 904 Carnic Alps, Austria /Italy 114, 344, 350, 366, 382, 385, 390, 418, 420, 428, 432, 436, 460, 462 498, 510, 550, 553, 554, 628, 692, 732, 831 Carozzi model 150 Carpathians 424, 522, 570, 571, 677 Carpathoporella 562 Carrara, Alpi Apuani, Italy 336 Carthage, northern Tunisia 304, 472, 474, 906 Caserta, Campania, Italy 468 Caspian Sea 20 Cassiduolidea 556 casts 121, 387, 390, 402, 529 cataclasis 238 cathodoluminescence (CL) 1, 64, 65, 66, 67, 68, 71, 221, 259, 267, 290, 299, 300, 301, 310, 330, 401, 483, 520, 655, 673, 732, 734, 889, 903; Plate 5 on p. 69 cave 10, 11, 698, 730, 731, 732, 819, 882; Plate 129 on p. 733 cave carbonate 13, 222, 732, 734 cave pearls 13, 157, 160 cave pisolite 158
937
cavity 62, 96, 194, 208, 220, 298, 384, 412, 647, 772, 787, 834 cay aquifer 276 Cayeuxia 371, 385, 409, 410, 762 Cayman Brac Island, British West Indies 208 CCD 455, 486, 794 CDP (Calcite-Dominated-Precursors) 99, 274 celleporiform 630, 632 cement 2, 4, 51, 61, 65, 66, 67, 68, 73, 82, 118, 120, 163, 178, 184, 193, 194, 196, 197, 205, 210, 212, 217, 218, 220, 221, 255, 259, 260, 267, 268, 270, 271, 272, 274, 276, 277, 279, 288, 289, 292, 300, 301, 304, 308, 310, 312, 314, 324, 325, 326, 340, 343, 346, 350, 352, 354, 356, 361, 362, 371, 374, 381, 385, 387, 390, 417, 428, 454, 468, 494, 504, 524, 529, 533, 534, 538, 541, 542, 550, 560, 582, 584, 589, 610, 628, 630, 643, 652, 654, 655, 656, 673, 677, 688, 690, 694, 698, 714, 728, 757, 766, 787, 788, 807, 833, 834, 835, 836, 837, 844, 857, 858, 859, 860, 865, 868, 873, 874, 880, 881, 889, 904; Plate 31 on p. 291; Plate 32 on p. 293; Plate 33 on p. 305; Plate 34 on p. 307 A and B cement 310 acicular 292, 294, 296, 438, 736, 843 anhydrite 301, 882 aragonite 14, 292, 533, 735, 857 aragonite fibrous, needle 290, 292, 308 aragonite spherulitic 308 asymmetrical 299, 300 bladed 294, 304, 837 blocky 14, 268, 284, 295, 297, 303, 306, 560, 766, 833 botryoidal 275, 294, 296, 346, 784, 837, 844, 857, 858, 860 burial 278, 300, 306, 310, 314, 320, 362, 374, 438, 628, 688, 700, 734 circumgranular 300, 438, 762 cryptocrystalline 93 displacive 784 dogtooth 3, 16, 68, 290, 294, 299, 300, 865 dolomite 301, 304, 329, 330, 740 dripstone 162, 294, 748, 837 drusy 14, 284, 294, 299 early 25, 194, 277, 297, 302, 304, 316, 780, 787, 844, 866, 896 equant 300, 301, 306 equigranular 284 fabric 274, 292, 299, 300, 673, 889 fan 299, 844 fibrous 37, 194, 195, 210, 290, 292, 294, 296, 308, 342, 731, 735, 788, 837, 858, 860 granular 3, 38, 268, 295, 300, 462, 731, 835 gravitational 180, 252, 298, 300, 304, 306, 748, 749, 766 High-Mg calcite 290, 306 interparticle 282, 301, 312, 314, 582, 649,700 intraskeletal 416 isopachous 284, 289, 292, 300, 304, 306, 334, 346, 710, 731, 760, 784, 824 isopachous fibrous 289, 296, 853 isopachous microcrystalline calcite 304 Low-Mg calcite 14, 38, 271, 308
938
marine 7, 71, 290, 292, 300, 303, 362, 835 marine phreatic 190, 304, 306, 844, 900 marine vadose 300 meniscus 14, 37, 38, 162, 276, 292, 294, 300, 304, 710, 728, 736, 748, 749, 837 meteoric 43, 284, 290, 292, 299, 304, 306, 308, 333, 582, 710, 856, 886 meteoric-phreatic 853 meteoric-vadose 292, 300, 304, 806, 818 Mg-calcite 38, 292, 308, 854, 857 micritic 37, 292, 296, 342, 362 microcrystalline 74, 93, 94, 117, 120, 295, 299, 300, 306, 343, 352, 837, 858, 860 microstalactitic 37, 300, 748 non-luminescent 734 overgrowth 296 palisade 298, 308, 731 peloidal 117, 292, 295, 306 pendant 14, 180, 276, 296, 300, 304, 306, 364 phosphate 207 pore-filling 282, 296, 356, 550 pore-lining 282, 302 radial-fibrous 68, 119, 268, 289, 298, 731, 844, 857 radiaxial 236, 289, 296, 298, 306, 548, 732, 788, 837 radiaxial fibrous 218, 221, 268, 289, 294, 297, 298, 303, 812, 833, 835 raggioni 298 reef 273, 304, 306, 843 rim 37, 117, 182, 306, 649, 682, 694, 798, 822, 886, 923 scalenohedral calcite 306 sequence 68, 306, 310, 837, 889 sinter 364 splayed-fibrous 300, 85, 860 stalactitic 731 stratigraphy 67, 68, 290 sulfate, sulfide 858 symmetrical 299, 300 synsedimentary 714, 786, 837, 842 syntaxial 306, 554, 548, 696, 760 syntaxial overgrowth 295, 298, 299, 300, 301 type 13, 68, 197, 274, 292, 295, 301, 302, 348, 723, 760, 830, 833, 837, 889 vadose 14, 38, 710, 735, 746, 757, 806, 818, 827 cement crust 15, 299, 300, 304, 306, 379, 379, 380, 836 cement crystal 71, 117, 270, 290, 296, 322, 800, 900 cement reef 836 cement-framestone 364 cement/grain ratio 314 cementation 2, 6, 8, 19, 40, 43, 50, 71, 73, 75, 78, 91, 94, 95, 106, 108, 156, 163, 164, 166, 169, 184, 185, 190, 193, 194, 198, 200, 206, 210, 267, 268, 271, 275, 276, 278, 282, 283, 288, 289, 290, 301, 302, 303, 304, 308, 310, 312, 314, 320, 334, 340, 346, 350, 372, 374, 438, 494, 497, 528, 541, 554, 559, 582, 592, 606, 610, 630, 632, 663, 714, 725, 726, 734, 738, 746, 748, 754, 769, 774, 787, 794, 795, 798, 806, 835, 837, 844, 854, 873, 878, 881, 882, 886, 889, 892, 896, 898, 905 cementstone 342, 361, 362, 710, 735, 844, 857, 858, 868 Cenomanian (Cretaceous) 444, 450, 466, 468, 472, 524,
Index
526, 713, 826, 871 Cenozoic 101, 121, 181, 270, 382, 389, 392, 393, 407, 412, 414, 424, 430, 435, 452, 456, 459, 478, 480, 486, 512, 515, 522, 524, 529, 544, 546, 564, 616, 626, 628, 630, 632, 635, 653, 654, 659, 664, 671, 702, 717, 793, 831, 852, 869, 872 cephalopod limestone 211, 533, 534, 538, 560, 796, 801 cephalopods 178, 220, 312, 533, 570, 618, 794, 795, 796, 872; Plate 90 on p. 535 ceramic 257, 903, 905, 906, 908, 910; Plate 151 on p. 909 and p. 911 ceratoporellids 506 chaetangiacean red algae 426 Chaetetes 436 chaetetid coralline sponges 130, 140, 502, 507, 762, 822; Plate 81 on p. 503 chaetetids 4, 386, 492, 502, 504, 506, 507, 766, 835 chalcedony 483, 484, 622, 643, 734 chalk 48, 65, 66, 78, 91, 92, 120, 211, 271, 272, 284, 314, 392, 504, 558, 643, 716, 800, 827, 877, 878, 880, 884 chalk reservoir 880 chamosite 637 Changsiangian (Permian) 346, 426 charcoal 908 Charly Johnson Guyot, Pacific 212, 304 charophycean algae 20, 155, 404, 406, 407, 619, 663, 728, 738, 740, 766; Plate 65 on p. 449 Charophyta 364, 442, 447, 448, 450, 452, 619 Cheilosporites 236 Cheilostomata, bryozoans 515, 630 chemical compaction 278, 318, 320 chemoautotrophic 51, 857, 858, 860 chemoherm 857, 858 chemoheterotroph 394, 611 chemolithotroph 84 chemostratigraphy 789 chemosynthesis 51, 184, 857 Chemtou Marble, Tunisia 322 Cheops pyramide 476 chert 18, 53, 65, 92, 268, 271, 276, 483, 484, 486, 497, 497, 559, 643 644, 659, 662, 732, 734, 742, 792, 797, 798, 800, 802, 903, 911 Chichaoua, Morocco 650 chicken-wire anhydrite 648, 708, 749 Chiemgau Alps, Bavaria, Germany 114 China 346, 532, 560, 756, 802 Chios Island, Greece 520 chironomid insect larvae 16 chitin 519, 547 Chitinoidella 490 Chitinoidella boneti 488 chitons 387 Chlamys islandica 44 chloralgal grain association 438, 626, 627, 628 chlorite 641 chlorocline 611, 635 chloroforam grain association 627, 628, 630 chlorophycean algae 404, 458, 563, 626 chlorosponge grain association 627 chlorozoan grain association 626, 627, 630, 634
Index
Choffatella decipiens 470 chonetid brachiopods 520 Chromophyta 404 Chrysalidina 454, 472 Chrysophyta 49 Cidaris, cidaroids 553, 554, 556, 558 Ciliata 453, 487 cipit boulder 788 Ciprocampylodon 562 circulation 18, 31, 42, 91, 512, 623, 654, 660, 662, 663, 668, 669, 671, 694, 700, 702, 719, 794, 848, 850 circumcrustation 374 725 circumgranular crack 726, 727, 728 circumgranular rim 120, 334 cirripedians 34, 90, 208, 389, 390, 392, 393, 394, 395 ascothoracican 392 Cisalveolina fallax 472 CL see cathodoluminescence Cladocora 42 Cladophora 376, 620, 738, 740 Cladophorites green algae, 735, 739 Claracrusta 135 classification of carbonate rocks 2, 250, 257, 310, 340, 346, 349, 361, 404, 591, 592, 593, 890, 895; Plate 43 on p. 351; Plate 44 on p. 353; Plate 45 on p. 355; Plate 46 on 359; Plate 47 on p. 363; Plate 48 on p. 365 Carozzi classification 243 depositional texture 348 descriptive 339, 340 Dunham classification 339, 364, 680 Embry and Klovan classification 339, 343, 364 expanded Dunham classification 348, 349 expanded Folk classification 356 Folk classification 339, 357, 358; Plate 46 on p. 359 genetic 58, 339, 340 mineralogical-chemical 340 mixed siliciclastic rocks 361, 366; Plate 49 on p. 367 Plumley classification 340, 591 Strohmenger and Wirsing classification 339, 357, 360 textural 339, 340, 895 Wright classification 339 clast *100, 166, 174, 228, 229, 236, 239, 252, 468, 605, 657, 661, 664, 665, 688, 723, 732, 772, 790, 810, 812, 847, 848 diagenetic 172 orientation 593 roundness 230 volcanic 230 clast analysis 166, 169, 229, 232, 236 clast support 229, 731, 772 clast/matrix ratio 232 clastic input 37, 44, 447, 794, 863 clasticity index 251, 257 clathrate hydrates 197 Clathrocoilona 384, 500, 673 clay 47, 54, 56, 94, 95, 96, 121, 169, 279, 301, 313, 316, 318, 352, 362, 592, 641, 642, 646, 662, 667, 670, 682, 739, 747, 764, 771, 774, 800, 802, 888, 895, 896, 898, 908, 910, 912 content 182, 313, 321, 322, 642, 794, 795, 884 minerals 54, 78, 320, 332, 619, 641, 642, 644, 896, 897
939
clay matrix 911 clay seam 94, 198, 696 claystone 18, 648, 742, 826, 827 cleavage 270, 297, 298, 301, 306, 326, 550, 551 Cleopatra VII 914 Cleopatra’s beach, Sedir Island, Turkey 913, 914 Climacammina 401 climate 6, 11, 13, 14, 31, 182, 358, 616, 728, 735, 769, 771, 803, 816, 862, 863, 881 arid 11 humid 730 semiarid 11, 725 climate change and fluctuations 13, 15, 50, 210, 604, 625, 628, 630, 641, 654, 804, 817, 870, climatic control 40, 55, 264 clinoform 781, 807 clinostratification 781 clionid sponges 40, 91, 118, 390, 922 clotted fabric 96, 111, 118, 312, 371, 395, 726, 735 clotted micrite 344 Cloudcroft, New Mexico, U.S.A. 702 cloudy inclusion 67 cluster analysis 251, 262, 263, 581, 627, 628, 677, 790, 792 cluster reef 835 clymenids 796 Clypasteroidea 556 Clypeina 176, 268, 434, 440, 444, 713, 762 Clypeina jurassica 432, 440 CO2 14, 40, 504, 756 CO2 partial pressure 333 coal 67, 659 coarse equant spar (ESC) 300 coarsening upward sequence 58, 806, 808, 813, 817 coast 8, 24, 174, 175, 452, 541, 546, 593, 594, 604, 611, 623, 653, 708, 717, 747, 863 coastal playa 132 coated grain 12, 36, 102, 121, 182, 186, 714, 723, 742, 782, 812, 828 coating 117, 587, 694 ferrugineous 798 micritic 858 microbial 726 coccoid cyanobacteria 370, 620 coccolith ooze 49, 71 coccolithophorids 49, 78, 91, 92, 93, 111, 284, 404, 406, 484, 490, 684, 793, 795, 797, 800, 872, 875 coccoliths 50, 78, 93, 490 coconut-meat 731 codiacean algae 423, 424, 562, 563, 566 Codiomorpha 562 Codium 423 Codonofusiella 462 coelobionts 559 coenocline 480 coenozone 440, 465, 570 cold seep 51, 374, 857 cold-seep carbonate 858, 490; Plate 148 on p. 861 cold-temperate 626, 634 cold-water 34, 40, 43, 44, 279, 308, 386, 515, 546, 551, 593, 610, 626, 628, 630, 803, 853, 854, 954 Coleoidea 533 collagen 559
940
colonization 89, 119, 380, 381 607, 830 color 54, 55, 195, 662, 897, 903, 905 Colospongia 835 Colospongia dubia 498 Comelico, northern Italy 460 commensalism 564, 569 community 211, 372, 382, 636, 681, 717, 770 compaction 49, 55, 60, 68, 74, 106, 117, 149, 152, 180, 182, 185, 193, 205, 218, 220, 224, 232, 268, 271, 277, 282, 288, 301, 310, 311, 312, 313, 314, 316, 320, 321, 332, 352, 362, 484, 493, 533, 606, 607, 637, 658, 708, 766, 774, 800, 881, 889, 890, 904; Plate 36 on p. 315 burial 314, 882 carbonate sand 314 chemical 301, 310, 311, 313, 314, 318, 320, 696, 890 criteria 312 differential 314 experimental 311 index 312 measurement 311, 312 mechanical 288, 301, 310, 311, 313, 314 model 311 compaction law 185, 313 comparison charts 174, 245, 246, 257, 258 compensation depth 611, 623, 794 competition 381, 384, 405, 416, 497, 625 component analysis 266, 807, 821 composite crack-seal vein 225 composite grain 37, 38, 163, 344, 584, 649, 698, 700, 702, 723, 760, 770 composite triangular diagram 628 composition 61, 78, 339, 468, 575, 578, 580, 603, 605, 673, 772, 868, 895 compositional log 60, 788, 789 compositional maturity 101, 176, 26, 7180 compressive strength 895, 904 computer models 657, 658, 724 conceptacles 414, 416, 420, 444 conceptual model 6, 657 conchiolin 529 Conchocelis 394 Conchodon 809 conchostrakes 739 concretion 200, 214, 484, 650 condensation 63, 204, 206, 210, 211, 212, 214, 215, 216,314, 536, 587, 696, 794, 796, 801, 828 condensation horizon and section 178, 204, 211 789, 816 condensation sequence 136, 203, 206, 211, 223, 559, 794 condensation surface 203, 206, *211, 216, 820 condensed grainstone 361, 480 conductivity 896 cone-in-cone 538, 541 conglomerate 53, 58, 161, 168, 216, 228, 230, 233, 239, 243, 244, 428, 592, 640, 688, 708, 719, 747, 771, 772, 773, 809, 810, 812, 847, 848, 865, 914 disorganized 773 flat-pebble 747, 749 intraformational 746, 747, 810 lag 712 significance 242
Index
Coniacian (Cretaceous) 383, 396, 827 coniatoids , coniatolite *158, 748 Conicospirillina 481 Conicospirillina basiliensis 466 Connexia 584, 844 Connexia carniapulchra 236 conodonts 71, 106, 387, 404, 559, 560, 650, 796, 848, 850; Plate 97 on p. 561 conservation of limestone 900; Plate 150 on p. 901 consistency of the substrate 610 constituent analysis 254, 258, 263, 264 constituent ranking 176, 260 constructor guild 342, 680, 831, 832, 833, 835 constructors 380, 478, 496, 673, 831, 832 contact 290 concavo-convex 311 fitted 772 grain-to-cement, grain-to grain, grain to pore 312 point, sutured, tangential 311 container organomicrite *84, 96, 374 continental margin 756, 793, 794, 795 continental rise 46, 48 continental shelves 9, 26, 46, 515, 793 continental slope 8, 9, 29 continuous sampling 63 contourite 47, 48, 486, 801, 802 conversion table 243 convex-up orientation 180, 686 convolute bedding 640, 775 cool water 2, 26, 35, 42, 44, 82, 390, 455, 515, 626, 627, 628, 630, 670, 667, 852, 853, 868; Plate 4 on p. 45 cool-temperate 31, 625, 630, 632 cool-temperate zone 34, 41, 626, 632 cool-water factory 873 cool-water coral reef Plate 147 on p. 855 Coorong, Australia 84, 308 Copenhagen, Denmark 854, 856 copepods 112 copper 379, 892 coprolites 170, 314, 824 coquina 60, 137, 178, 239, 523, 526, 548, 580, 598, 662, 696, 723, 738, 752, 759, 782, 810, 923 coral association 854 coral bank 853 coral bioherms 41 coral community high-growing 678 coral debris 366, 578, 601, 604, 605, 777, 810 coral groups 509 coral limestone, cool water 321, 577, 853, 904 coral reef 34, 36, 40, 385, 386, 396, 493, 512, 600, 602, 616, 623, 624, 625, 632, 635, 692, 771, 777, 789, 804, 833, 854, 856, 863, 864, 871 coral skeleton 617, 656, 854 coral thicket 41, 853, 856 coralalgal 37, 38, 82 coralligène 41, 416 Corallina 413 corallinacean red algae 41, 140, 284, 381, 382, 404, 412, 406, 413, 416, 417, 418, 412, 420, 444, 481, 612, 616, 617, 620, 630, 632, 635, 702, 819; Plate 54 on p. 415 coralline algal gravel 44 coralline genera identification key 413 coralline red algae 34, 41, 44, 75, 90, 97, 122, 123,
Index
137, 138, 140, 264, 284, 292, 308, 381, 382, 404, 405, 406, 413, 414, 416,417, 418, 412, 420, 444, 476, 478, 541, 542, 544, 608, 611, 612, 632, 667, 725, 748, 831, 854, 871, 886, 917 articulated, crustose 444 coralline red algal reef 44 coralline sponges 346, 492, 497, 506, 576, 676, 690, 766, 789, 790, 792, 831, 835, 844, 871 coralliths 123 corals 41, 71, 74, 180, 208, 236, 248, 249, 252, 264, 268, 271, 274, 306, 310, 342, 346, 352, 354, 377, 381, 382, 384, 385, 386, 390, 395, 402, 404, 432, 438, 458, 481, 491, 502, 507, 509, 512, 524, 566, 584, 590, 600, 602, 604, 606, 608, 612, 616, 618, 625, 628, 634, 654, 656, 666, 668, 675, 680, 688, 690, 692, 714, 755, 762, 766, 771, 772, 789, 792, 807, 809, 810, 824, 828, 831, 833, 834, 835, 848, 854, 864, 865, 871, 886, 904, 905, 908, 922 ahermatypic 769, 856 alveolitid 500 auloporid 334 azooxanthellate 512, 853, 854 deep-water corals 512, 512, 922 diversity 512, 853 hermatypic 616, 626, 628, 856, 873 preservation 512, 514 reef-building 385, 393, 616 rugose 260, 284, 362, 374, 384, 392, 500, 509, 510, 514, 636, 673, 692, 752, 764, 768 scleractinians 314, 385, 509, 512, 514, 577, 601, 616, 635, 636, 717, 754, 854; Plate 84 on p. 513 soft corals 509, 600 tabulate corals 289, 344, 384, 507, 509, 510, 514, 598, 673, 692, 717, 764, 768, 919; Plate 83 on p. 511 zooxanthellate 612, 616, 624, 625, 635, 853 Cordilites 420 core 61, 648, 661, 680, 884, 896 core analysis 816, 862, 878, 884, 885 747 Corg cornstone 11 Coroong model, dolomitization 331 correspondence analysis 263, 724, 790, 792, 843 corrosion 90, 93, 207, 211, 484, 523, 796, 859, 860, 897 corrosion surface 204, 858 cortex 89, 118, 121, 143, 156, 413, 414, 423, 424, 426, 427, 558 of ooids 143 of rhodoids 139 cortoids *100, 114, *118, 119, 120, 121, 164, 176, 182, 251, 312, 358, 390, 436, 579, 611, 624, 628, 636, 694, 702, 714, 717, 725, 749, 757, 850, 888, 889; Plate 10 on p. 115 Cosenza, Calabria, Italy 853 Couches Rouges, Cretaceous 802 counting 254 coverage 260, 602, 603, 604 coverstone 341, 342, 343, 346, 366, 494 Cow Head Peninsula, Newfoundland 564 Cozumel, Yucatan, Mexico 390, 600-602, 603, 605
941
crabs 40, 111, 175, 185, 189, 738 cracking 900 cracks 126, 207, 224, 225, 239, 726, 742, 809 Cracow, Poland 114 Crassicolaria parvula 488 Crassicollaria brevis 488 Crassicollaria intermedia 488 Crassicollaria parvula 487 craze plane 364 creeping 771, 779, 780 Cretaceous 13, 61, 90, 92, 113, 118, 122, 124, 134, 166, 181, 208, 211, 216, 306, 310, 313, 314, 350, 358, 364, 372, 379, 382, 384, 385, 390, 392, 394, 396, 407, 410, 412, 414, 416, 417, 418, 420, 424, 426, 435, 440, 442, 444, 447, 448, 450, 452, 456, 458, 459, 465, 466, 468, 470, 474, 478, 480, 486, 487, 488, 490, 492, 496, 500, 504, 506, 508, 512, 515, 524, 528, 530, 533, 534, 541, 542, 547, 554, 556, 558, 559, 562, 563, 566, 568, 569, 570, 576, 597, 606, 608, 610, 613, 616, 620, 626, 628, 630, 634, 635, 644, 646, 647, 661, 664, 682, 686, 698, 706, 710, 714, 716, 728, 740, 754, 756, 766, 773, 797, 800, 802, 803, 807, 814, 819, 827, 828, 831, 834, 852, 864, 865, 869, 871, 872, 874, 878, 880, 884, 892, 904, 905, 906, 907, 908 Cretaceous-Tertiary boundary 435, 522, 524, 650, 856 Cretacicrusta 562 Crete Island, Greece 728 Cribrogenerina 454, 460 Cribrosphaerella ehrenbergi 284 cricoconarids 537, 538, 556 crinoid accumulation 60, 551, 552, 553, 554, 591, 638, 696, 875 crinoid arm elements 556 crinoid columnals 182, 251, 346, 550, 551, 552, 764 crinoid packstone 230, 349, 554 crinoid plates 352, 554 crinoid stem, ossicles 298, 510, 550, 551, 552, 554 crinoid-brachiopod limestones 522 crinoids 4, 41, 78, 221, 306, 390, 459, 486, 522, 548, 550, 551, 552, 580, 582, 591, 596, 597, 606, 608, 638, 675, 692, 696, 754, 768, 785, 795, 798, 830, 842, 843, 848, 850, 918-921; Plate 95 on p. 555; Plate 96 on p. 557 allochthonous deposits 552 classification 550 planktonic 481, 551, 553, 556, 684 stalked 550, 551, 556 transport 552 Crinozoa 550 Croatia 18 Croococcales 408 cross bedding 14, 58, 177, 178, 182, 184, 220, 523, 581, 590, 593, 598, 647, 663, 593, 718, 719, 725, 752, 774, 806, 893, 904 cross lamination 58, 177, 299, 775 cross stratification 58, 229, 593, 595, 613, 637, 666 cross-plot 884 crossed lamellar microstructure 105 crumbly fracture 225 crust 117, 118, 136, 162, 189, 194, 197, 200, 205, 207, 210, 212, 214, 325, 334, 364, 370, 371, 373, 374, 384, 385, 395, 412, 414, 416, 417,
942
444, 497, 542, 560, 562, 563, 566, 577, 606, 620, 690, 710, 727, 731, 735, 747, 749, 757, 766, 781, 790, 824, 838, 844, 857, 886, 898, 899, 900, 920, 921 algal crust 178 biogenic 268, 379, 381, 384, 782, 827 caliche 663 cement 663 Fe-Mn 82, 757, 794, 795, 796, 798, 801 goethite 214, 215 micritic 117, 770, 858 microbial 186, 374, 690, 714, 757, 784, 837 pedogenic 728 pisolitic 710 spongiostromate 268 stromatolitic 807 thrombolitic 82 weathering 897 crustacean arthropods 18, 111, 404, 518, 542, 546, 766 crustose corallines 416 cruststone 341, 342, 366 cryptalgal 184, 190, 195, 371, 747, 838, 843 cryptic habitat 62, 82, 236, 346, 370, 373, 384, 406, 493, 496, 559, 560, 562, 566, 604, 690, 844 cryptocrystalline 75, 119 Cryptostomata, bryozoans 515 crypts 412 crystal boundary 73, 290, 324 crystal diminution 321, 548 crystal fabric 76, 324, 328, 648, 897 crystal fan 868 crystal growth 68, 117, 295, 296, 321, 899 crystal lattice 68, 271, 321 crystal mold 333 crystal morphology 16, 295, 306 crystal shape 96, 271, 324 crystal silt 59, 97, 178, 192, 299, 306, 310, 326, 582, 608, 694, 731, 762, 766, 812, 818, 837 crystal size 74, 75, 76, 78, 92, 94, 96, 97, 104, 290, 321, 324, 325, 330, 334, 361, 884, 889 crystallization texture 324 Ctenostomata, bryozoans 388, 515 cumulative curve 244, 245, 247, 251 Cuneiphycus 422 Cuneolina 472 Cuneolina ex gr. pavonia 470 cuneolinid foraminifera 465, 470 Cupressocrinites 764 current 18, 42, 80, 117, 177, 182, 184, 211, 246, 332, 339, 348, 354, 356, 400, 401, 460, 483, 491, 523, 526, 538, 553, 588, 589, 593, 595, 596, 606, 634, 713, 737, 745, 746, 750, 766, 782, 794, 795, 796, 798, 862, 864 bottom 798, 801, 802, 842 counter 801 current deposit 696, 834 current indicator 56, 63, 553, 598, 610, 775 current transport 211, 214, 480, 536, 538, 604, 665, 779, 781 current velocity 550, 832 cutting effects 400, 401 cuttings 680, 884, 885, 888, 896 Cuzzo di Lupo, Palermo, Sicily 438
Index
cyanobacteria 14, 15, 19, 20, 21, 67, 75, 78, 80, 81, 84, 88, 97, 105, 118, 119, 120, 122, 124, 125, 132, 135, 136, 138, 140, 155, 184, 194, 195, 364, 369, 370, 371, 376, 379, 382, 384, 386, 387, 390, 394, 395, 396, 402, 404, 406, 407, 408, 409, 410, 412, 436, 438, 452, 491, 493, 504, 512, 562, 563, 564, 566, 568, 569, 579, 600, 604, 611, 612, 618, 620, 624, 656, 663, 671, 696, 704, 706, 717, 735, 738, 739, 740, 748, 749, 762, 764, 850, 856, 875, 898, 920; Plate 53 on p. 411 boring 119, 612 coccoid 84, 118, 408 filamentous 119, 125, 408 nonmarine 404 nonskeletal 408 oncoids 646 planktonic 408 porostromate 766, 768, 812, 824 spongiostromate 812 cyanobacterial mat 84, 93, 111, 190, 192, 193, 601 cyanoids, cyanoliths 122, *123, 124, 125, 126, 134, 137, 140, 579, 611, 696, 807, 810, 824, 875; Plate 12 on p. 141 cyanophytes 370, 404, 410 cycle 6, 62, 658, 752, 803, 804, 805, 821, 866, 881 basinal 807 depositional 136, 642, 806, 807, 812 diagenetic 806 fifth-order 804, 817 first-order 805 fourth-order 804, 805, 817 high-frequency 804, 806 Latemar cycle 806 Lofer (Dachstein) cycle 808 Milankovitch 805 peritidal 806 platform 524, 804 second-order 805 small-scale 63, 659 third-order 804, 805, 808 cycle duration 817 cycle thickness 805 cyclic carbonates 63, 593, 803-805, 810, 814, 885; Plate 140 on p. 811 cyclic sedimentation 57, 59, 62, 63, 101, 139, 184, 635, 647, 658, 788, 807, 892 cyclicity 6, 656, 792, 803, 805, 807, 814 deep-marine 807 in reefs 807 pelagic 807 cyclocrinitids 422 Cyclogyra 713 cyclone 249, 594, 600 Cyclostomata, bryozoans 515 cyclostratigraphy 442, 803 cyclothem 63, 523, 630, 752, 803, 808 Cylindroporella 440, 442, 628 Cymopolia 430 Cymopolia elongata 444 Cymopolia inflataramosa 444 Cyprus 801, 802 cystoids 550, 554, 630; Plate 95 on p. 555
Index
943
Cystoporata, bryozoans 515 cysts 432, 442, 452, 500 Czatkovice, Crakow 158, 170, 190 Czech Republic 538, 673, 686 D δ13C 619, 624, 642, 654, 737, 858, 912 δ18O 624, 737 Dachstein, Austria 678, 810 Dachstein cycle 808, 812 Dachstein limestone (Triassic) 506, 524, 792, 789, 809 dacryoconarids 534, 537, 538, 540 Dactyloteuthis 537 Damous-el-Karita, Carthage, Tunisia 906, 908 Danian (Tertiary) 444, 831, 852, 856 Darwin Guyot, Pacific 304 dasyclad assemblages 440, 442, 444, 446 dasyclad green algae 108, 114, 164, 176, 184, 230, 236, 241, 260, 268, 325, 368, 390, 401, 405, 406, 422, 430, 432, 434, 435, 436, 438, 440, 444, 447, 459, 465, 466, 507, 524, 564, 568, 570, 582, 584, 591, 612, 628, 635, 638, 663, 678, 692, 694, 702, 704, 713, 714, 719, 742, 746, 748, 750, 752, 768, 777, 779, 781, 782, 809, 810, 827, 828, 844, 848, 850; Plate 59 on p. 432 Cenozoic 444 definition 430 determination, classification 434, 435, 438 diagnostic criteria 430, 431, 432 Early Cretaceous 442; Plate 63 on p. 443 Jurassic 440; Plate 62 on p. 441 Late Cretaceous Plate 64 on p. 444 Late Paleozoic Plate 60 on p. 437 Paleocene Plate 64 on p. 444 significance 442 Triassic 438, 442; Plate 61 on p. 439 dasyclad spores 436 dating of clasts 169 Dead Sea 20, 81, 540 Debki well, Poland 314 Debnik, Poland 418, 500, 554 debris 14, 76, 96, 135, 172, 302, 304, 346, 470, 534, 538, 548, 597, 598, 600, 604, 692, 696, 716, 764, 769, 795, 798, 801, 809, 810, 820, 827, 835, 843, 872, 877, 878, 888, 918, 919 debris flow 19, 47, 48, 49, 50, 118, 184, 230, 298, 464, 660, 662, 686, 688, 696, 698, 717, 723, 759, 769, 771, 772, 773, 780, 781, 784, 785, 788, 796, 802, 822, 834, 847, 880, 888 debrite 770, 772 decapod crabs 132, 544, 547 decarboxylation 279 Deckerella 108 declivity 769, 788 decompaction 311, 313 dedolomite 238, 333, 334; Plate 39 on p. 327 dedolomitization 55, 66, 180, 279, 205, 325, 326, 333, 334, 819, 889, 896; Plate 40 on p. 335 deep burial 78, 224, 271, 275, 277, 299, 301, 332, 333 deep marine 3, 8, 26, 31, 47, 48, 56, 58, 90, 207, 274, 768, 186, 406, 453, 455, 480, 486, 491, 492, 493, 496, 504, 509, 532, 541, 544, 548, 550, 554, 455, 558, 566, 588, 606, 626, 643, 653,
654, 696, 757, 769, 773, 774, 793, 800, 802, 804, 822, 857, 866, 874, 922; Plate 139 on p. 798 deep shelf, deep shelf margin 91, 544, 662, 682, 692, 696 deep-sea trench 9, 46 deep-water 635, 844 canyon 682 limestone 63 stromatolite 212 deepening 63, 576, 641, 807, 817 deepening-upward 804, 809, 817, 819, 820, 822 definition of MFT (microfacies types) 575, 576, 578, 582; Plate 100 on 583; Plate 101 on p. 585 deformation 311, 312, 314, 320, 322, 358, 490, 640, 893, 913 dehydration 301, 311, 622 Delaware Basin 843, 844, 846 Delphi, Greece 474 delta 660, 862, 864 demosponges 113, 389, 492, 493, 494, 496, 506, 507, 856 dendrolites 372, 373, 378, 395, 408 Dendrophyllia 853 Dendrophyllia candelabrum 854 denitrification 369 deposit feeder 185, 558, 606, 607, 610, 624 deposition 80, 341, 361, 364, 530, 551, 587 mechanical 893 out-of-suspension 795 depositional break 642, 648, 655 condition 8, 9, 54, 61, 182, 200, 578, 582, 589 criteria 268 fabric 578, 581, 786, 889 history 582 microfacies 267, 268, 582 pattern 242, 615, 648, 659, 661, 868 rates 642, 805 texture 50, 61, 232, 260, 288, 339, 340, 361, 578, 584, 591, 610, 756, 898 depositional angle 769 depositional constraints 803 depositional environment reconstruction 848 depositional models 657, 884 depositional sequence 805, 817, 884 depositional system 657, 658, 816 dereption 204 desert basin 647 desiccation 157, 192, 193, 218, 288, 312, 354, 364, 647, 706, 710, 728, 746, 749, 750, 806, 809, 812, 818 desiccation crack 708, 735, 746, 750, 752, 806 Destroyer Guild 831 destruction 66, 90, 102, 175, 241, 299, 386, 389, 390, 392, 662, 90 desulfurication 504 detrital minerals 251, 257 detritus 167, 168, 358, 371, 547, 662, 754 feeder 532, 554 Devonian 68, 113, 134, 166, 195, 196, 222, 289, 296, 362, 370, 372, 378, 379, 384, 396, 409, 410, 412, 420, 424, 426, 428, 438, 448, 450, 452, 458, 459, 491, 496, 500, 510, 519, 520, 522, 530, 533, 534, 537, 538, 544, 546, 547, 548,
944
554, 559, 560, 563, 564, 566, 575, 597, 638, 640, 661, 664, 672, 673, 674, 682, 684, 716, 732, 738, 750, 752, 764, 768, 773, 788, 794, 796, 814, 821, 836, 840, 872, 878, 880 Devonshire Bay, Bermuda 38 dewatering 193, 236, 311, 352, 640 Diadema type 556 Diademnoides 562 diagenesis 1, 6, 8, 13, 16, 18, 43, 47, 54, 55, 57, 58, 59, 61, 67, 68, 78, 90, 91, 94, 95, 97, 99, 106, 107, 127, 149, 175, 185, 189, 194, 197, 205, 206, 224, 226, 254, 268, 271, 272, 275, 277, 278, 279, 282, 299, 300, 302, 303, 304, 308, 310, 313, 321, 325, 332, 334, 340, 346, 362, 372, 390, 402, 405, 465, 497, 514, 533, 564, 582, 604, 607, 641, 643, 644, 636, 648, 652, 653, 654, 658, 710, 753, 794, 805, 818, 837, 854, 856, 866, 867874, 877, 881, 882, 888, 889, 892, 893, 896 burial 275, 276, 277, 301, 306, 322, 324, 881 early 12, 278 facies-selective 107, 766 freshwater 4, 121, 276, 528, 879 late 68, 278 marine 276, 277, 819, 881 marine phreatic 277 marine-vadose 252, 837 meteoric 276, 364, 753, 881, 882 meteoric phreatic 438 meteoric-vadose 205, 762, 806, 837 unconformity-related 279 vadose 306, 766, 824 diagenesis of fossils 67 diagenetic alteration 58, 68, 71, 104, 118, 127, 262, 310, 368, 399, 405, 520, 617, 641, 654, 655, 735, 798 diagenetic environment 67, 75, 97, 267, 271, 274, 275, 276, 277, 290, 292, 296, 299, 306, 308, 310, 334, 484, 710 diagenetic fabric 316, 656 diagenetic grainstone 354 diagenetic history 67, 74, 206, 224, 308, 582, 836, 878, 895 diagenetic micrite 94 diagenetic microfacies 267, 268, 582; Plate 28 on p. 269 diagenetic overprint 71, 198, 595, 648, 889, 890 diagenetic packstone 480 diagenetic pathway 107, 216, 515, 615, 626, 820, Plate 35 on p. 309 diagenetic potential 43, 267, 308, 854 diagenetic process 1, 55, 60, 169, 175, 177, 188, 200, 267, 272, 308, 310, 336, 350, 362, 484, 579, 617, 644, 660, 877 diagenetic rock type 97, 361; Plate 47 on p. 363 diagenetic sequence 308, 890 diagenetic sieve 106, 107 diagenetic solution 232, 619, 653 diagenetic stage, succession 310, 325, 646 diagenetic taxa 399 diagenetic texture 267, 310, 582, 889 diagram, aragonite-calcite-terrigenous material 867 diastem 168, 169, 204 diatoms 18, 19, 20, 49, 125, 369, 404, 458, 611, 624 Dibunophyllum 420
Index
Diceras 528 Dichothrix 735 Dictyoconus algerianus 470 Dicyclina 465, 470 Didemnum 562 Didymi Mts., Greece 344 differential energy index field 592 digital image 254, 257, 282 dike 58, 217, 640, 731 dilution cycle 866 dimension stone 903 Dinant, Belgium 450, 460 Dinantian (Carboniferous) 732, 842 dinoflagellates 101, 450, 452, 458, 483, 512 dip angle 576, 770, 788, 798 Diplopora 268, 430, 435, 444 Diplopora annulata 438 Diplopora phanerospora 432 disaerobic 613, 836 disarticulation 105, 523, 546 Discocyclina 476, 478 discocyclinids 481 disconformity 204, 208, 809, 818 discontinuity 4, 59, 177, 203, 204, 212, 214, 215, 216, 262, 642, 653, 756, 757, 796, 805, 904; Plate 22 on p. 209 glossary 204 discontinuity surface 62, 130, *203, 204, 208, 211, 212, 605, 646, 795, 818: Plate 22 on p. 209 Discosiphonella lercarensis 498 discriminant function analysis 247 disintegration 75, 92, 424, 426, 897, 899 disjectoporids 835 dismicrite 192, 357, 358, 759, 762 disphotic zone 26 displacive crystallization 728 dissolution 13, 14, 18, 21, 40, 47, 49, 50, 80, 90, 91, 92, 93, 94, 104, 106, 119, 120, 149, 184, 185, 189, 195, 200, 206, 208, 210, 211, 212, 218, 221, 222, 224, 232, 262, 268, 270, 271, 272, 276, 278, 279, 283, 292, 298, 299, 302, 304, 306, 308, 310, 316, 318, 321, 332, 333, 350, 389, 390, 414, 428, 494, 497, 504, 510, 514, 529, 533, 534, 548, 622, 630, 643, 652, 661, 663, 686, 728, 732, 734, 753, 766, 769, 772, 794, 795, 801, 853, 865, 866, 867, 881, 882, 886, 888, 890, 897, 898, 905, 924 level 49, 277 meteoric 364, 730 seam 185, 268, 316, 318 dissolution clast 801 dissolution vug 886 dissolution/precipitation reprecipitation 271, 272, 330, 924 distal tempestites 595 distal turbidite 57 Distichoplax biserialis 444 distribution coefficient 654 distribution pattern 37, 371, 406, 438, 442, 446, 450, 478, 480, 483, 630, 636, 657, 671, 672, 677, 760, 789, 875 disturbed micrite 192 Ditrupa 381, 632 Divaca, Slovenia 448, 478 diversity 62, 107, 216, 260, 377, 385, 406, 442, 464,
Index
480, 483, 510, 515, 523, 526, 576, 587, 612, 613, 614, 615, 616, 618, 619, 632, 663, 702, 725, 748, 749, 752, 757, 821, 843, 848, 856, 871 Djebel Assiz, northern Tunisia 182 Djebel (Jbel) Bou Dahar, High Atlas, Morocco 759 Djebel Tebaga, southern Tunisia 424, 428, 462 dogtooth, see cement dolocretes 11 dolomicrite 708, 720, 746, 748 dolomicrostone 328 dolomite 12, 13, 53, 66, 68, 70, 71, 169, 190, 253, 259, 267, 268, 271, 277, 278, 289, 292, 308, 316, 322, 324, 325, 328, 329, 330, 334, 339, 349, 622, 641, 644, 647, 648, 652, 662, 663, 675, 708, 716, 719, 746, 749, 752, 805, 826, 827, 828, 877, 881, 882, 883, 884, 892, 895, 896, 903, 904, 907; Plate 39 on p. 327 baroque 68 burial 332 cement 276, 329 classification 328 crystal mold 180 crystal size texture 328 description and terminology 328 fabric 153, 326, 328, 329 ferroan 333 hydrothermal 883 mosaic 304, 326 primary 330 replacement 330, 356 sucrosic 883 zonar 326 dolomite clast 813 dolomite crystals 70, 303, 334, 644 dolomite problem 325, 328 dolomite reservoir 882 dolomite rhombohedron 68, 326, 333, 334 dolomite texture 330 Dolomites (Italy) 426, 498, 584, 781, 782, 788, 808, 813 dolomitization 55, 66, 68, 70, 106, 128, 152, 169, 172, 186, 189, 190, 267, 268, 271, 272, 278, 279, 282, 283, 288, 313, 314, 316, 320, 325, 328, 329, 330, 331, 582, 604, 644, 748, 766, 877, 878, 879, 881, 882, 883, 886, 888, 889, 896; Plate 40 on p. 335 burial 334 by seawater 331 models 325, 330, 331, 332, 334 replacement 304, 326, 334 selective 121, 189, 334 dolomold 205, 334 dolosparstone 328 dolostone 66, 328, 339, 340, 669, 813, 865, 868, 895 Donax 524 Donets Basin, Ukraine 462 Donezella, donezellid algae 76 436 Dorag model 332 downhole mixing 885 drainage inbibition capillary pressure 878 dripstone 11, 13 dropstone 172, 625, 630, 860 drowned platform 131, 134, 218, 221, 223, 553, 624, 793, 794
945
drowning 4, 210, 220, 478, 625, 630, 660, 664, 694, 756, 757, 759, 760, 789, 816; Plate 133 on p. 761 drowning unconformity 756, 817 drusy mosaic 300 dry-peel technique 65 dual point counting 255 Dubai, UAE 528 Duino, Trieste, Italy 526 dune 14, 56, 186, 246, 666, 667, 698, 708, 717, 725 Dunham classification 61, 244, 348, 350, 361, 366, 578, 591, 668, 723, 837, 904, 905 durability 903 duricrusts 11 Dvinella 436 Dweller Guild 502, 606, 673, 681, 831, 832 Dybowskiella 382 dynamic microfacies type 723 dysaerobic 56, 537, 613, 614, 647 dysoxic 613, 614, 892 dysphotic 394, 396, 611, 612, 636, 637, 638, 873, 874 E Earlandia 454, 460 earthquake 49, 59, 194, 218, 230, 239, 640, 688 eccentricity 804, 805 echinoderm fragments 114, 251, 301, 306, 313, 428, 538, 548, 550, 551, 556, 630 echinoderms 3, 71, 104, 106, 137, 182, 186, 195, 260, 270, 290, 298, 306, 334, 350, 352, 358, 366, 400, 401, 404, 472, 518, 546, 548, 550, 554, 556, 578, 579, 580, 581, 582, 584, 597, 598, 618, 630, 644, 662, 666, 670, 675, 682, 686, 688, 692, 696, 717, 719, 757, 759, 760, 766, 777, 782, 789, 790, 795, 798, 800, 812, 820, 822, 824, 828, 834, 850, 853, 854, 865, 869, 875, 908, 910, 920, 922 classification 550 in thin sections 550 echinofor grain association 627 echinoids 34, 40, 41, 44, 90, 188, 189, 298, 387, 427, 480, 548, 550, 551, 554, 556, 580, 582, 606, 632, 650, 667, 725, 784, 800, 812, 917; Plate 95 on p. 555; Plate 96 on p. 557 classification 554 irregular echinoids 556 regular echinoids 554, 556 spines 444, 553, 554, 556, 558, 917 ecologic reef 830, 831 ecologic succession 210, 405, 589, 624 ecologic zonation 264, 500, 600 ecosystem 384, 600, 604, 610, 623, 754 edgewise conglomerate 229, 428 Edwards Formation, Cretaceous, Texas 330, 350 efforescence 899 Efluegelia 128, 147, 420, 422, 524, 644, 827 Egypt 111, 713, 886, 899 Eh condition 55, 154, 379, 653 EI log, energy index, limestone classification 592 Eifel, Germany 19, 448 Eifelian (Devonian) 424, 500, 510, 538, 750, 752 El Haouria, Cap Bon, Tunisia 284, 725
946
Eldfisk Field, North Sea 800 Elianella 420 Ellef Ringness Island, Canadian Arctic 860 Ellipsactinia 432, 481, 502 Emanuella 752 emergence 210, 220, 728, 813, 818, 819 emersion 169, 204, 206, 364, 809, 817, 818, 827, 886, 892 empty bucket 667, 874 Emsian (Devonian) 195, 306, 334, 382, 390, 507, 510, 520, 534, 538, 554, 556, 598, , 762, 840, 842 encrinites 551, 552, 696, 849 encrustation, biogenic 131, 163, 170, 208, 212, 230, 235, 268, 302, 306, 310, 346, 372, 380, 381, 382, 384, 395, 414, 464, 507, 512, 523, 560, 566, 580, 596, 605, 606, 607, 608, 636, 662, 673, 694, 696, 712, 738, 746, 748, 762, 764, 772, 785, 844, 857, 905, 908 encrusters 37, 42, 59, 90, 164, 205, 211, 222, 342, 362, 369, 371, 374, 379, 380, 381, 382, 384, 385, 395, 412, 414, 481, 515, 530, 559, 564, 569, 606, 624, 663, 673, 717, 762, 772, 836, 837 endo-upwelling 194, 332, 623 endobenthos 186, 380, 610 endolithic organisms 40, 88, 90, 91, 386, 396, 533, 636, 737 endostromatolites 131, 220, 372 endosymbionts 51, 456, 563, 568, 610, 612, 628, 632, 635 Endothyranopsis 454, 459 Endothyranopsis crassa 460 energy environment, classification 591 energy level 153, 176, 248, 292, 352, 587, 588, 589, 590, 591, 592, 593, 600, 665, 674 enfacial junction 290, 322, 324 England 330, 541, 766 Enigma 562 enterolithic 647, 648, 708 Entomozoacea, ostracods 544, 546 envelope formation 120, 923 enveloping 590 environment, depositional 7, 10, 24, 44, 102, 105, 107, 117, 203, 249, 330, 360, 370, 379, 386, 388, 404, 453, 465, 480, 492, 512, 520, 522, 551, 575, 576, 578, 584, 591, 658, 668, 672, 676, 680, 692, 698, 712, 754, 806, 821, 826, 827, 868, 878 environmental change 57, 127, 203, 723, 756, 818, 826 environmental constraints 262, 447, 453, 587, 617, 624, 753 environmental control 4, 63, 261, 361, 384, 405, 480, 575, 616, 648, 658, 680, 724, 816 Eocene 91, 181, 240, 366, 417, 444, 447, 476, 478, 480, 512, 530, 533, 692, 694, 708, 728, 886, 899 eolian carbonates, eolianite 10, 11, 12, 14, 37, 38, 120, 172, 186, 197, 290, 304, 592, 641, 663, 717 Eothrix 554 Eovelebitella 430, 434, 436, 850 Eovelebitella occitanica 432 epeiric 28, 47, 647, 648, 658, 669, 795 epeiric platform 27, 658, 660, 669, 681 epeiric ramp 660 epibenthos 34, 81, 88, 90, 380, 382, 523, 606, 667, 836
Index
epicontinental 26, 92, 553, 554, 669 Epidaurus, Argolis, Greece 218, 801 epifauna 596, 610, 662, 692 epifluorescence 67, 73, 428 epigenetic 892 Epimastopora 428, 436, 628, 781 epipelagic zone 9 epiphytes 128, 382, 480, 437, 746, 748, 853 Epiphyton 409, 410, 504, 507 epiphytozoans 382 episclerophytes and episclerozoans 381 episkeletobionts and episkeletozoans 382 Epistacheoides 422, 638 epizoans 128, 382, 515, 853 epizoophytes 381 epizoozoans 382 EPS see extracellular polymeric substances equatorial zone 626, 630, 862 Erfoud, Anti-Atlas, Morocco 306, 334, 510, 520, 534, 538, 554, 556, 598 Erlangen PaleoReef Database 393 Ernstbrunn, Lower Austria 218, 466 erosion 24, 48, 58, 90, 113, 166, 169, 170, 184, 194, 200, 203, 206, 221, 222, 233, 236, 241, 248, 354, 386, 464, 578, 596, 606, 641, 642, 659, 708, 746, 756, 759, 769, 770, 772, 775, 777, 779, 781, 784, 798, 806, 816, 818, 821, 862, 863, 865, 893 base 523, 596, 598 mechanical 90, 92 rate 302, 370, 658 submarine 24, 605, 786 erosion surface 203, 595, 827 erosional break 812 erosive contact 595, 596, 777 escarpment 771 Esfandiar, Iran 785 estimating methods 245, 258 estuarine 24, 487, 588, 617, 619 etching 57, 64, 65, 113 Ethelia 417 eubacteria 369 eucaryonts 18, 82, 404, 408, 830, 838 Eugonophyllum 420, 427, 428 Eugonophyllum mulderi 428 euhedral 74, 96, 644, 732 eulittoral 740 euphotic 44, 611, 612, 636, 662, 663, 666, 712, 754, 755 Eureka, Mamouth Mts., USA 650 euryhaline 618, 750 eustatic 210, 730, 803 eutrophic 49, 394, 623, 624, 670, 836, 864 euxinic 560, 613, 646 evaporation 11, 14, 19, 290, 302, 304, 330, 647, 816 evaporation model 330, 331, 748 evaporite 53, 65, 193, 238, 267, 301, 316, 330, 332, 333, 619, 622, 642, 644, 647, 648, 659, 681, 700, 708, 710, 719, 720, 747, 749, 816, 819, 827, 843, 879, 880, 882 minerals 279, 333, 622, 647, 648 evenness 619 event deposits 184, 59, 63, 597 eventstone 340, 597
Index
947
Everglades, Florida 13, 408 evolution 462, 464, 490, 658, 868, 869, 789, excentrique 731 exfoliation 897 exokarst 13 exposure 4, 288, 746, 769, 788, 818 exposure index 704 exposure surface 203, 205, 216, 643, 806, 813, 817, 820 extensional movement 220, 224, 226, 889 extinction 225, 289, 297, 298, 301, 378, 435, 548, 559, 615, 623, 871 extinction cross 538 extracellular polymeric substances (EPS) 82, 369, 371 extraclast 164, 166, *167, 170, 172, 176, 230, 239, 248, 320, 358, 360, 366, 579, 590, 591, 596, 678, 714, 739, 757, 760, 774, 779, 786, 800, 806, 820; Plate 16 on p. 171 Exuma, Bahamas 38, 167 Exvotarisella 638
F φ skewness 245 φ sorting standard deviation 245 φ value 243, 244, 251 fabric 60, 66, 67, 73, 74, 76, 177, 239, 301, 316, 322, 324, 334, 377, 395, 578, 590, 607, 613, 673, 681, 717, 800, 837, 884, 888, 890, 893, 905 depositional 177, 349, 350, 368, 648, 717 diagenetic 97, 177, 352, 648 facies analysis 6, 28, 447, 491, 641, 654, 655, 878, 885 number 593, 807 shift 264, 660 zone 480 facies belt 6, 28, 447, 491, 647 657, 659, 660, 661, 662, 663, 671, 680, 713, 716, 718, 821, 826, 827, 843 facies control 895 facies model 1, 6, 657, 674, 803, 893, 896 facies type 668, 673, 815, 895, 896 Facies Zone (FZ) after Wilson 657, 661, 662, 663, 664, 671, 672, 676, 678, 682, 684, 686, 688, 692, 692, 694, 696, 698, 700, 702, 704, 706, 708, 712, 713, 714, 718, 762, 764,768, 822, 824, 848, 893 factor analysis 247, 262, 263, 266, 790, 792 factory benthic carbonate 873 cool-water 873 mud-mound 873, 874 tropical 873, 874 fair-weather wave base 196, 460, 523, 595, 635, 636, 637, 638, 659, 662, 663, 664, 665, 666, 692, 694, 712, 842 Fakse, Denmark 854, 856 Faksephyllia 854 Famennian (Devonian) 378, 392, 418, 459, 518, 534, 544, 732, 831, 840, 872 Faroe Island Shelf 44 Fasciella 420, 638, 848, 850 Fatima, Portugal 140 fault 218, 224, 736, 756, 768, 794, 857, 882, 884 Favosites styriacus 510
favositids 764 Favreina 314, 440, 544, 547, 766, 824; Plate 93 on p. 545 Fe 55, 68, 70, 207, 271, 301, 330, 332, 333, 623, 624, 653, 736, 747, 857, 892, 897 Fe-ooids 154, 214 Fe-oxides, hyroxides 198, 333, 465, 649, 710, 820 Fe/Mn crust, nodule 4, 82, 212, 214, 215, 822 fecal pellets 40, 49, 92, 93, 110, *111, 112, 113, 116, 117, 118, 248, 249, 302, 544, 547, 606, 646, 650, 692, 700, 717, 726, 749, 764, 766, 858, 860, 875 Feichtenstein, Salzburg, Austria 566 feldspar 66, 362, 622, 642, 643, 644, 911 Fenestella 108 fenestrae 14, 197, 344, 364, 371, 706, 747, 762, 850 fenestral fabric 58, 126, 128, 176, *190, *192, 197, 201, 358, 374, 578, 636, 681, 704, 747, 749, 750, 806, 809, 810, 824, 889, 904; Plate 20 on p. 191 fenestral oncoid 440, 694 fenestral pore 193, 197, 312, 344, 371, 374, 746, 747, 748 Ferafra Oasis, Western Desert, Egypt 181 ferricretes 11 ferroan calcite, dolomite 271, 274 fertilizer 896 Feuerkogel, Aussee, Austria 212, 556 fibrous cement see cement field work 53, 251, 655 filamentous cyanobacteria and algae 21, 408, 412 filaments 89, 97, 125, 182, 184, 190, 218, 220, 230, 350, 372, 373, 413, 414, 418, 423, 424, 426, 428, 484, 486, 524, 526, 528, 554, 562, 684, 735, 747, 757, 759, 760, 782, 790, 795, 796, 822 filter feeder 118, 520, 550, 593, 857 filtration fan 550 fining-upward sequence 58, 90, 659, 884 firing temperature 911 firmground 166, 186, 189, 204, 208, 210, 212, 606, 608, 610, 854 fish 40, 386, 387, 560, 614, 692, 738, 739 bones, scale teeth 559, 560, 649, 650, 712, 737, 850; Plate 97 on p. 561 fissure 4, 59, *216, 217, 220, 221, 222, 238, 310, 372, 560, 605, 730, 731, 732, 798, 800, 819, 889, 899 filling 217, 220, 221, 222, 223, 271 synsedimentary 218, 801 tectonic 218 terminology 224 Fistulipora 516 fitted grainstone 361 fitting 232, 318, 320, 710, 731 flame structures 640, 732 Flaserkalk 538, 775 flash flood 641, 863 flat-pebble 169, 184, 230, 235, 239, 596, 708, 747 floating rhombs 334 floatstone 214, 341, 349, 350, 354, 356, 362, 366, 507, 523, 589, 597, 605, 608, 632, 663, 670, 675, 686, 688, 692, 696, 706, 708, 719, 723, 737, 739, 747, 759, 781, 785, 786, 810, 824, 826, 827, 834, 837, 848, 852, 853, 856, 858,
948
886, 898, 921; Plate 45 on p. 355 flooding 205, 206, 597,647, 675, 732, 757, 759, 760, 788, 804, 806, 817, 820 flooding surface 3, 12, 136, 138, 523, 672, 806, 808, 816, 817, 819, 820, 821 Florida 13, 24, 25, 28, 41, 78, 80, 84, 88, 89, 119, 186, 249, 279, 288, 330, 364, 424, 594, 632, 657, 660, 664, 769, 836, 863 flow rate 220, 282, 292 flowstone 11, 731 fluid and gas expulsion 857, 858 fluid inclusion 1, 66, 67, 225, 277, 278, 301, 306, 326, 332 fluid seepage 857 fluorapatite 649 fluorescence microscopy 64, 65, 66, 67, 82, 739 flute mark 56, 593, 773, 774 fluvial carbonates 18, 23 fluvial transport 174, 175, 641, 710 fluxoturbidite 775, 786 flysch 538 fold 336, 462, 640, 780, 79 foliated microstructure 105 Folk classification 246, 348, 356, 357, 360, 361, 366, 368, 578, 591, 905 for-algalith 123 foraminifera 4, 20, 37, 38, 41, 44, 49, 50, 71, 78, 90, 97, 108, 128, 130, 132, 137, 164, 176, 184, 208, 214, 215, 220, 233, 241, 249, 252, 264, 268, 290, 304, 348, 350, 352, 358, 370, 374, 381, 382, 384, 385, 386, 387, 390, 395, 399, 400, 401, 402, 406, 422, 428, 437, 438, 442, 444, 447, 452, 453, 454, 455, 456, 457, 458, 459, 460, 464, 465, 468, 470, 472, 474, 476, 478, 480, 481, 482, 484, 491, 492, 542, 551, 559, 564, 566, 568, 569, 571, 584, 593, 612, 616, 618, 624, 626, 628, 632, 633, 636, 638, 644, 662, 663, 666, 667, 668, 671, 672, 676, 677, 678, 680, 684, 686, 690, 692, 694, 700, 702, 704, 706, 710, 713, 714, 717, 746, 760, 762, 766, 768, 771, 774, 777, 782, 785, 90, 795, 796, 798, 800, 802, 824, 826, 842, 848, 850, 853, 860, 870, 873, 875, 880, 881, 904, 908, 910, 912, 922 agglutinated 106, 212, 382, 454, 455, 456, 458, 459, 460, 465, 468, 478, 481, 616, 678, 719 benthic 2, 270, 440, 450, 453, 455, 459, 460, 465, 470, 472, 478, 480, 486, 510, 546, 570, 632, 671, 676, 678, 694, 700, 702, 713, 719, 748, 749, 760, 762, 781, 798, 800, 810, 812, 820, 822, 827, 828, 837, 848, 850, 859, 864, 865, 869, 870, 886, 917, 920, 922 boring 563 Carboniferous: Plate 67 on p. 461 classification 454, 464 Cretaceous: Plate 70 on p. 468; Plate 71 on p. 470; Plate 72 on p. 473; Plate 73 on p. 474 diagnostic criteria 453 distribution 465, 481 encrusting 124, 135, 137, 140, 427, 460, 480, 696, 835, 856, 886 Jurassic Plate 69 on p. 466 Larger foraminifera 399, 400, 402, 453, 458, 459, 468, 472, 478, 480, 628, 635, 670, 671, 692,
Index
717, 719, 754, 755, 762, 766 microstructure 75, 106, 382, 453, 454, 455, 452, 454, 459, 460, 465, 481 orbitoidacean 466, 472; Plate 72 on p. 473 orbitoidacean rotaliinid 466, 632 orbitolinid 468; Plate 70 on p. 468 pelagic 450, 452, 797, 819 Permian Plate 67 on p. 461 planktonic 76, 91, 296, 453, 455, 456, 457, 465, 466, 474, 478, 480, 481, 616, 618, 624, 654, 672, 795, 797, 827, 828, 853, 872; Plate 73 on p. 474 protoglobigerinid 466 sessile 214, 215, 381, 386, 444, 542, 690, 850 systematic categories 455 taxonomic determination 457 Tertiary Plate 73 on p. 474; Plate 74 on p. 476; Plate 75 on p. 478 Triassic Plate 111 on p. 679 foraminifera-bryozoa association 632 foraminiferal abundance analysis 256 foraminiferal assemblage 456, 465, 480, 570, 635, 677 foraminiferal limestone 181, 400 foraminiferal ooze 49 foraminiferal zone 460 foramol grain association 43, 480, 626, 627, 630, 632, 634, 670, 771 forereef 218, 289, 297, 412, 500, 506, 534, 552, 560, 600, 664, 668, 673, 674, 686, 714, 792, 837, 846, 888 foreset 773, 787 foreslope 497, 659, 688, 696, 712, 781, 834, 836, 846 formation test 890 Forni Avoltri, Carnia, Italy 182, 382, 428, 553, 564 Forschia prisca 460 fossil assemblage 145, 592, 671, 675 fossilization 389, 404 Fossillagerstätten 220 fossils in thin-sections 108 fractal distribution pattern 189 fractionation 616, 617, 654 fracture 14, 192, 217, 220, 222, 224, 225, 226, 280, 288, 336, 554, 726, 728, 857, 858, 880, 882, 883, 884, 889 fractured reservoir 884 fracturing 182, 224, 238, 268, 279, 308, 314, 318, 528, 686, 787, 859, 877, 879, 880, 881, 882, 884, 886, 889, 890 fragmentation 93, 96, 104, 105, 200, 402, 427, 507, 523, 591, 834, 857 framboidal pyrite 644, 647 framebuilders 37, 59, 63, 168, 412, 416, 563, 564, 566, 569, 624,663, 664, 832, 838, 854, 856, 873 framestone 341, 342, 343, 344, 346, 364, 366, 510, 576, 624, 663, 690, 719, 723, 762, 766, 833, 837, 856, 905; Plate 42 on p. 347 framework 4, 42, 288, 342, 343, 344, 346, 416, 512, 690, 733, 754, 762, 770, 807, 817, 830, 872 France 3, 13, 41, 128, 140, 195, 289, 374, 396, 450, 459, 500, 510, 512, 520, 538, 570, 644, 673, 692, 764, 766, 796, 802, 833 Franciscan Complex, California 800 Frankenwald, Germany 538, 560, 684 Frasnian (Devonian) 193, 500, 831, 870, 871, 872
Index
949
Frasnian/Famennian extinction 379 freezing 67, 68, 897, 898, 899 freezing damage 900 French Reef, Florida Bay, U.S.A. 190, 424 frequency 7, 257, 258, 260, 358, 406, 459, 581, 586, 592, 603, 673, 677, 807, 847, 864 frequency analysis 2, 243, 261, 263, 480, 540, 792 frequency curve 244, 247 frequency distribution 263, 480, 540 frequency in thin sections 405 frequency index 256, 257 frequency measurement 256, 262 frequency of clasts 229, 232 freshwater 6, 11, 13, 14, 15, 16, 19, 25, 95, 217, 334, 364, 369, 374, 404, 408, 422, 477, 522, 544, 617, 618, 619, 642, 646, 652, 663, 732, 734, 737, 809 freshwater algae 122, 447 freshwater lake 19 freshwater lens 274, 276 Friedman diagram 246, 251 Frondicularia 466 frost weathering 896, 897, 900 Frutexites 409, 412 Fucus-Ascophyllum belt 916 Fuerteventura, Canary Island, Spain 157, 290, 292, 542 fungi 119, 120, 208, 369, 386, 387, 388, 390, 394, 395, 396, 564, 612, 727, 788, 898 Fürstenstand, Graz, Austria 764 Fusanella 559, 562, 567 fusulinid foraminifera 114, 170, 182, 241, 314, 322, 366, 399,400, 454, 458, 459, 462, 464, 467, 582, 584, 677, 668, 694; Plate 68 on p. 463 fusulinid zonation 808 FZ see Facies Zones G Gaissau, Hallein, Austria 358, 598, 678 Galala Basin, Egypt 728 Galeanella 454, 678 Galeanella panticae 678 gamma-ray-neutron log 884 Garnberg, Künzelsau, Germany 598 Gartnerkofel, Carnic Alps, Austria 530 Garwoodia 409, 410 gas bubbles 16, 192, 193, 735 gas clathrate hydrates 197, 230, 858 gas migration 889, 890 Gastrochaeniidae 389 gastropods 2, 18, 19, 20, 38, 61, 90, 108, 111, 112, 176, 178, 182, 214, 221, 241, 255, 268, 334, 350, 352, 364, 366, 378, 387, 390, 393, 427, 428, 522, 529, 530, 554, 582, 584, 598, 606, 614, 618, 622, 630, 663, 667, 686, 688, 692, 696, 698, 704, 706, 714, 719, 725, 738, 739, 740, 747, 748, 749, 750, 752, 757, 766, 777, 809, 810, 812, 822, 824, 844, 850, 857, 865, 886, 917, 919, 921, 922; Plate 89 on p. 531 actaeonellid 865 boring 393 microstructure 529 pelagic 49, 793 preservation 529, 530
Gebel Maghara, Sinai, Egypt 470 Gemeridella 559, 562 geoarchaeology 903 geochemical analysis 1, 40, 53, 63, 64, 80, 92, 225, 254, 289, 337, 370, 619, 624, 641, 652, 653, 655, 764, 804, 807, 877, 906 geochemistry 6, 70, 614, 635, 652, 725, 731, 805, 858 geopetal 96, 177, 178, 180, 201, 310, 344, 358, 392, 401, 647, 675, 714, 717, 731, 787, 865, 892, 893, 919, 922; Plate 17 on p. 179 criterion 177 fabric 126, 139, *177, 180, 229, 312, 352, 389, 482, 772, 782, 801, 819, 835, reoriented 178, 180, 181 geothermal gradient 314, 623 Getaia 562 Gettysburg Seamount, Atlantic Ocean 917 Ghawar field, Saudi Arabia 877, 880 ghost structure 322, 643 Ghyben-Herzberg model 276 Gilbert hurricane 602 Ginsburg model 804 Girvanella 113, 122, 128, 135, 137, 371, 372, 384, 385, 409, 410, 764 Girvanella ducii 410 Givetian (Devonian) 410, 458, 500, 510, 519, 538, 796, 870, 871, 872 Gizeh, Cairo, Egypt 10, 11, 22, 474, 864, 868, 899, 907 glacial-marine deposit, glacial carbonates 11, 14, 42, glaciation 230, 628, 630, 868, 740 glacigene sediment 868 glacio-eustasy 804, 808 glaebule 726 glass 905, 911 glauconite 131, 205, 207, 208, 210, 212, 214, 257, 548, 624, 637, 643, 644, 646, 712, 757, 795, 800, 854 authigenic 646 detrital 646 pellets 646 glauconitization 211 glendonite 14, 860 gliding plane 322 global change 6, 760 Globigerina 455, 474, 819, 898 globigerinid foraminifera 454, 455, 456, 474, 481, 819, 898, 908 Globigerinoides 474 Globochaete 218, 221, 556, 559, 566, 790, 798 globoids 13 Globorotalia 474, 481 Globotruncana 450, 471, 474, 684 Globotruncana ex gr. arca 474 Globotruncana stuartiformis 474 globstone 342 Glomospira 678 Glomospirella 678 Goddess Iuno 910 goethite 54, 207, 210, 214 Golden Lane field, Mexico 880 Gondwana 630 goniatites, cephalopods 534, 794, 796 Goosenecks, San Juan, Utah 284, 358
950
gorgonians 41, 346 Gosaukamm, Austria 164, 268, 374, 566, 678, 690, 694, 706, 789, 790, 792, 835 Gotland Island, Sweden 181, 394, 544, 919, 920, 924 graded bedding 54, 158, 595, 596, 647, 773, 775 gradient analyses 636 grading 177, 178, 229, 232, 239, 523, 673, 680, 774, 776, 777 inverse 177, 773, 922 reverse 177, 773, 775 grading curve 776 grain 4, 14, 62, 73, 109, 110, 111, 247, 263, 302, 308, 312, 321, 340, 341, 348, 349, 350, 352, 356, 358, 360, 361, 371, 374, 382, 579, 580, 587, 591, 593, 606, 652, 680, 747, 724, 864 association 44, 625, 626, 627, 628, 630, 632, 634 contact 74 diminution 81, 94, 106 distortion 277, 314 lacustrine 364 morphometric criteria 590 non-skeletal grain *100, 628 point contact 350 skeletal grain *100, 101 significance 101 telescoping 312 grain analysis 723, 788 grain association analysis 2, 4, 44, 272, 625, 626, 627, 628, 630, 632, 670, 680, 723, 753, 792, 819, 848, 852, 869, 874 grain boundary crack 225 grain bulk percentage 356 grain category 257, 258, 259, 261, 348, 360, 366, 578, 580 grain composition, compositional log 6, 264, 581, 606, 668, 712, 713, 723, 769, 781, 788, 789, 790, 800, 827, 898 grain contact 292, 311, 312, 314, 316, 318 grain crushing, deformation 311, 314 grain density 350 grain diameter 590 grain diversity 361 grain flow 47, 48, 49, 184, 771, 773, 781 grain orientation 62, 63, 177, 181, 182, 201, 590, 773 grain packing 350 grain penetration 362 grain shape 94, 173, 248, 313 grain size 54, 56, 62, 74, 96, 97, 103, 173, 174, 175, 176, 178, 189, 230, 243, 244, 245, 246, 247, 248, 249, 250, 251, 254, 257, 272, 277, 313, 314, 320, 357, 366, 370, 515, 578, 587, 590, 592, 595, 605, 606, 607, 610, 642, 666, 668, 746, 753, 754, 755, 773, 774, 776, 778, 779, 781, 786, 802, 806, 807, 820, 863, 879, 895, 896, 897, 912 avarage grain size 52, 244, 252 cumulative curve 247 data, conversion 247 distribution 65, 198, 243, 244, 245, 247, 248, 696 homogenization 97 measurement 62, 596 parameter 243, 245, 250, 251, 590, 779, 786 scale 243, 244 significance 253
Index
studies 250, 251 grain support 94, 189, 311, 320, 348, 349, 350, 352, 354, 366, 579, 642, 694, 696, 706, 786, 788, 813, 818, 865,884, 897, 904, ; Plate 43 on p. 351 grain type 61, 100, 102, 177, 216, 258, 259, 340, 357, 358, 360, 361, 578, 579, 658, 677, 680, 714, 717, 723, 724, 760, 769, 826, 874, 890, 897, 905 grain volume 175, 312 grain-bulk 255 grain-micrite ratio 593 grain-solid 255 grain/matrix ratio, grain/micrite ratio 258, 593 grainification 354, 726 grainstone 2, 14, 24, 37, 166, 180, 188, 193, 214, 224, 236, 312, 313, 320, 348, 349, 350, 352, 354, 358, 362, 364, 436, 438, 444, 450, 452, 459, 464, 465, 480, 482, 542, 548, 551, 576, 579, 589, 593, 595, 596, 597, 598, 608, 627, 634, 637, 638, 653, 662, 663, 664, 665, 666, 668, 672, 675, 676, 680, 686, 688, 694, 696, 698, 700, 702, 712, 713, 714, 718, 719, 725, 728, 739, 740, 745, 746, 747, 752, 757, 759, 760, 762, 766, 777, 781, 782, 786, 788, 789, 790, 806, 807, 809, 812, 813, 817, 821, 824, 826, 827, 843, 844, 848, 849, 850, 858, 865, 868, 884, 886, 890, 896, 898, 899, 900, 920; Plate 45 on p. 355 condensed 354, 362 eolian 14, 37 fitted 354 Granada, Spain 848 Grand Banks, Canadian Shelf (DSDP leg 43) 468 Grand Cayman 769 granite 248, 911, 913 granular calcite 218, 221 granular phosphorites 648 granular texture 646, 650 grapestone 36, 37, 38, 102, *163, 164, 166, 356, 676, 694, 700 graphic kurtosis 245 graphic log 60 graphic mean 245 graphic skewness 245 graptolites 404, 559, 560, 684, 686, 795; Plate 97 on p. 561 grass 663, 836 gravel 167, 243, 416, 667, 754, 786, 788 gravitational infill 206 gravity 49, 230, 772 gravity deposit 769 gravity-displaced blocks 864 gravity flow 47, 48, 49, 50, 177, 178, 229, 230, 659, 770, 771, 773, 781, 789, 834 gravity mass transport 218, 769, 780 Graz, Austria 61, 424, 510, 764, 768; Plate 135 on p. 765 grazing 91, 138, 370, 378, 384, 386, 387, 390, 392, 393, 416, 619, 624, 704, 747, 757, 759, 838 Great Australian Bight 852 Great Bahama Bank 3, 25, 38, 80, 88, 111, 164, 581, 607, 668, 878 Great Barrier Reef 2, 82, 144, 146, 249, 496, 660, 863 Great Oolite, Bathonian 145 Great Salt Lake, Utah 20, 21, 146, 147, 150
Index
951
Greece 140, 465, 526, 570, 677, 822, 824 green algae 20, 95, 97, 104, 136, 140, 249, 252, 274, 376, 390, 394, 396, 404, 405, 407, 408, 409, 414, 422, 426, 427, 447, 448, 494, 504, 563, 564, 568, 592, 604, 611, 612, 620, 624, 626, 628, 632, 634, 638, 672, 675, 713, 717, 738, 739, 740, 750, 754, 828, 844 planktonic 422, 452 Green Lake, New York State 19 greenhouse 272 Greenland 40 Grimming, Styria, Austria 566 Griotte (Devonian) 796 groove mark 56, 593, 774 GroßhöfleinBurgenland, Austria 414 Grötlinbo, Gotland, Sweden 552 Grotto Beach, San Salvador Island, Bahamas 38, 164 groundmass 73, 228, 232, 239, 261, 348, 356, 361, 366, 368 groundwater 12, 332, 588, 644, 710, 734, 746 growth bands 410, 524, 589, 590, 864 growth cavity 58, 59, 192, 371, 678, 782, 784, 785, 844 growth form 59, 62, 68, 220, 260, 290, 346, 348, 373, 377, 379, 381, 395, 402, 410, 413, 423, 424, 428, 432, 460, 515, 590, 591, 600, 605, 606, 630, 635, 636, 673, 674, 723, 807, 816, 831, 837 growth pattern 67, 136, 140, 180, 590, 591, 616, 619, 764 growth potential 658, 756, 816, 873 Gschöllkopf, Sonnwendgebirge, Austria 678 Guadalupe Mountains, Texas and New Mexico, U.S.A. 843, 844, 846 Guadalupian (Permian) 303, 462, 871 guild 62, 624, 673, 807, 831, 832 guild concept 807, 837 Gulf Coast U.S. 74, 879, 880 Gulf of Aqaba 111, 863 Gulf of Batabano, Cuba 163 Gulf of Suez, Egypt 146, 147, 374, 708, 863, 881 Gurumugl, Mount Hagen, Papua New Guinea 554 gutter casts 596, 598 guyot 48, 648, 756, 794 gymnocodiacean algae 97, 184, 405, 422, 424, 426, 427, 702, 713; Plate 57 on p. 424 Gymnocodium 424, 427 Gymnocodium bellerophontis 426 Gymnolaemata, bryozoans 515 gypscretes 11 Gypsina 123 gypsum 20, 53, 66, 277, 292, 330, 333, 622, 644, 647, 648, 663, 708, 716, 720, 749, 828, 900; Plate 109 on p. 645 gyrogonite morphology, characean algae 448, 450 Gyroporella 236, 430, 584 Gzhelian (Carbniferous) 420, 428, 460
H hadal 9, 46, 48 Haddonia 454, H. heissigi 478
hadopelagic zone 10 Haidach, Tyrol, Austria 608 Halicoryne 452 Halimeda 38, 78, 88, 103, 104, 248, 249, 288, 356, 406, 423, 424, 426, 428, 604, 668, 886 Halimeda model 78, 81, 89, 702, 875 Halimeda nana 444 Halimeda ‘reefs', thickets 424 Halimeda tuna 89 halite 20, 292, 647, 663, 716, 749 Hallein, Salzburg, Austria 484, 786, 797 Hallock and Schlager model 625 Hallopora 918 Hallstatt facies 789, 790, 792, 801 Hallstatt limestone 526, 792, 796 Halobia 220, 790 Halysis 564 Hamar Laghdad, Erfoud, Morocco 548, 840 Hamelin Pool, Shark Bay, Australia 376 Hammatostroma albertense 500 Handler Range, Antarctica 306 hard bottom 554, 603, 605, 606, 607, 608 hard substrate 379, 380, 381, 382, 384, 605, 608, 646 hardground 4, 40, 50, 63, 74, 96, 97, 131, 136, 166, 170, 185, 189, 198, 203, 204, 205, *206, 207, 208, 210, 211, 212, 221, 235, 262, 300, 302, 379, 382, 387, 390, 392, 528, 534, 602, 603, 606, 610, 650, 653, 714, 757, 769, 794, 795, 796, 798, 800, 801, 806, 809, 817, 819, 820, 827, 856, 892, 917, 921; Plate 23 on p. 213 hardground clast 198, 546, 798 Hardie-Stanley model 273 hardness 896, 897 hardpan 726 Harz, Germany 114, 197, 316, 374 Hattusa, Hittite reign, Turkey 310, 904, 905 Haurania 465 Haute Savoie, France 480 Hauterivian (Cretaceous) 872 Hawasina Complex, Oman 382, 690 Hedstroemia 409 Heezen Guyot, northwestern Pacific 212 helictite 13, 731 heliolitid corals 509, 510, 764, 918; Plate 83 on p. 511 hematite 54, 55, 114, 548, 710 Hemigordius 454, 460 hemipelagic 537, 538, 653, 666, 807, 866 hemipelagic limestone, sediment 47, 48, 54, 62185 hemipelagite 802 Hemmaberg, Carinthia, Austria 907 herbivorous 111, 140, 387, 405, 607 Hergla, Hammamet, Tunisia 322 hermacline 611 Heterastridium 4, 508 heterocorals 510, 514, 848; Plate 83 on p. 511 heterogeneity of facies criteria 885 Heterohelix 474 heterolithic bedding 182 heteropods 49, 91, 532 Heteroporella zankli 438 Heterostegina 632, 886 heterotroph 84, 369, 385, 386, 493,624, 626, 636, 754, 844, 854, 873, heterozoan grain association 626, 630, 632, 771
952
Hettangian (Jurassic) 353, 872 hexacorals 512 hexactinellids 492, 493, 494, 496, 576 Hexaphyllia mirabilis 510 hiatus concretion 204, 210, 212 high energy environment 24, 28, 37, 146, 250, 339, 352, 354, 444 476, 526, 588, 589, 590, 591, 598, 634, 657, 680, 694, 713, 805, 812, 896 high-latitude 630, 863 High-Mg calcite 20, 43, 68, 70, 119, 121, 137, 144, 154, 200, 267, 268, 270, 271, 274, 276, 290, 292, 297, 304, 308, 321, 334, 374, 390, 404, 408, 412, 416, 444, 497, 498, 500, 533, 548, 562, 580, 617, 620, 630, 632, 753, 854, 874 high-temperature burial zone 67 highstand sea level 130, 216, 554, 630, 660, 661, 664, 675, 688, 713, 759, 789, 790, 815, 816, 817, 818, 820, 821, 822, 827, 828, 838 highstand shedding 784, 785, 807, 821 highstand systems tract 523, 817, 818, 820, 826, 827 Higumastra 483 hinterland 642, 772, 827, 862, 863 Hippurites colliciatus 526 Hippuritid rudists 526 Hittite reign 310, 905 HMC, see High-Mg calcite Hobuggen, Gotland Island, Sweden 530 Hochschwab, Styria, Austria 444 Hofolding-1 well, Bavaria, Germany 470 Hoher Göll, Bavarian Alps, Germany 502, 534 Holocene 12, 14, 38, 42, 50, 95, 155, 169, 288, 296, 304, 371, 519, 529, 533, 602, 657, 664, 745, 762, 844, 906 Holopella gracilior 532 holothurian sclerites 554, 556, 558 holothurians 104, 106, 550, 556, 558; Plate 96 on p. 557 Holy Cross Mountains, Poland 450, 520, 632, 730, 750, 752 Homalozoa 550 Honshu, Japan 802 Hormosinacea 470 horse tail structure 185, 316, 318 host rock 221, 222, 643, 644, 877, 892 hot vents 623 Howchenia 460 HST see highstand Huda Juzna, Slovenia 498 Hudson Bay 669 humid 19, 206, 642, 659, 660, 661, 663, 710, 734, 745, 748, 749, 752, 815, 816, 863, 881, 898 hummocky 56, 594, 595, 596, 597, 780 Hungary 190, 808 hurricane 177, 249, 593, 594, 600, 603, 604 Hurricane Gilbert 594, 600, 601, 603 hyaline microstructure 304, 454, 500 hyalithids 507 hyalosponge 626, 627 Hydra Island, Greece 218, 498 Hydractiniidae 508 hydraulic competence 754, 755 hydraulic fracturing 224, 226 hydraulic sorting 459, 774 Hydrobia 740 hydrocarbons 51, 54, 99, 264, 283, 284,
Index
318, 326, 374, 453, 500, 656, 657, 756, 768, 770, 858, 872, 879, 880, 882, 883, 900 hydrocarbon exploration 264, 878 hydrocarbon formation 857 hydrocarbon inclusion 67, 857 hydrocarbon migration 656, 884 hydrocarbon production 284, 500, 877 hydrocarbon reservoir 145, 152, 156, 162, 182, 225, 242, 292, 311, 346, 428, 436, 605, 843, 877, 886 hydrocarbon seep 42 hydrocorallines 346 hydrodynamic condition, process 117, 182, 203, 247, 248, 249, 250, 263, 348, 350, 478, 480, 497, 576, 579, 582, 588, 589, 590, 642, 667, 668, 753, 759, 773, 807, 818, 914 hydrogen index 614 hydroids 604 hydrostatic pressure 49, 277, 301, 635, 637 hydrothermal 50, 51, 220, 326, 332, 643, 654, 842 hydrothermal dolomitization 331, 884 hydrothermal fluid 51, 842 hydrothermal vent 842, 857 hydrozoans 42, 44, 105, 386, 422, 432, 491, 496, 500, 502, 506, 507, 508, 509, 562, 566, 604, 853, 854, 856 Hyella 394 Hyella gigas 390 hyolithids 537, 540,692 hypersaline 21, 80, 163, 270, 299, 330, 408, 442, 455, 541, 617, 618, 619, 622, 643, 644, 647, 648, 656, 663, 850 hypidiotopic 74 hypolimnion 18 hypothallus 414, 417, 420, 422, 444
I Ibach Total Impregnation 900 Iberg, Harz, Germany 374 ice 47, 626, 630, 654 icehouse 272 ichnofabric 59, 60, 185, 212, 607, 610 ichnofabric index 188 ichnofossils 14, 394, 396, 717, 805, 895 iden 239, 316 iden-supported fabric 198, 318 idiomorphic 644 Idrija, Slovenia 460 ikaite 14, 625 illite 277, 279, 641, 642, 644, 732 illite crystallinity 337 image analysis 66, 185, 188, 244, 259, 261, 283, 376, 395 imbrication 181, 182, 184, 211, 480, 590, 718, 781 Immouzzer, High Atlas, Morocco 648 impact 228, 740 impregnation 65, 170, 212, 465, 547, 649, 900 impunctate brachiopods 520 in-place mineralization 81 in-situ brecciation 157 inarticulate brachiopods 519
Index
inclusion 67, 74, 306, 735, 908 index fossil 444, 466, 476, 490, 524, 533, 538 index of clasticity 251 India 478 Indian Ocean, Indo-Pacific 36, 506, 613, 623 Indonesia 142, 153, 414, 490, 620, 886 industrial use 895, 896 infauna 185, 186, 188, 189, 662, 663, 692 infilling 325, 464, 482, 497, 579, 702, 706 Ingram’s scale 56 inland sabkha 20 inner shelf, see shelf inoceramid bivalves 68 Inoceramus 90, 524, 800 inozoan sponges 4, 346, 492, 498, 502, 506, 835; Plate 79 on p. 498 insect larvae 18 insoluble residue 12, 70, 316, 318, 320, 336, 362, 642, 780, 805 inter-supratidal 330, 606 interbedding 210, 553, 746, 800 intercrystalline porosity see porosity interglacial 14, 16, 38, 789, 821 intergranular pore see porosity intermediate disturbance hypothesis 600 intermound 63, 854 internal micrite 190, 578 internal sediment 55, 118, 178, 192, 194, 197, 221, 223, 270, 299, 334, 364, 559, 690 internode 447, 448, 552, 854 interparticle pore see porosity interparticle void 480 intertidal 9, 25, 185, 186, 197, 330, 344, 354, 408, 455, 492, 523, 526, 541, 554, 569, 588, 611, 617, 626, 635, 636, 647, 659, 666, 669, 704, 706, 708, 710, 745, 746, 747, 748, 749, 750, 752, 806, 812, 818, 824, 850, 875, 878 intertidal member 75, 806 intertidal zone 9, 290, 514 interwell interpretation 885 intra-platform basin 47 intra-shelf basin 34, 47 intraclast 15, 16, 20, 37, 38, 40, 41, 71, 76, 113, *166, 167, 168, 169, 170, 176, 184, 248, 249, 251, 257, 356, 357, 358, 360, 362, 366, 368, 450, 544, 579, 590, 591, 596, 597, 640, 649, 650, 666, 682, 692, 694, 716, 717, 718, 719, 723, 738, 746, 747, 750, 766, 777, 795, 798, 809, 812, 820, 843, 852, 919, 921; Plate 16 on p. 171 intraformational pebbles 523 intragranular crack 225 intraplatform 676, 678 intraskeletal void 180 intrasparite 357, 762 intrasparrudite 358 inundites 597 inverse grading 37, 54, 161, 364 inversion 268, 271, 321 Involutina 822 Involutina liassica 222, 798 involutinid foraminifera 454, 464, 465, 466, 598, 677, 678, 679, 798 Iowa, USA 68
953
ipsonite 374 Iran 508, 532, 694, 784, 884, 886 Iraq 880, 886 Ireland 893 iron see Fe crust 207, 214 minerals 153, 646 oolite see Fe sulfide 84, 103, 112, 484, 613, 647 ironstone 154 Irregularia, echinoids 554 Isakov Guyot 212 Isastraea 512 Ischyrospongia 492 Isfahan, Iran 494 Isla Perez, Alacran reef, Yucatan 71, 88, 225, 249, 332, 520, 587, 612, 635, 654, 801, 805, 808 isoclyne 442 isotherm 31, 458, 626, 634 isotope 71, 290, 299, 889 isotope analysis 70, 71, 88, 225,332, 337, 520, 582, 612, 654, 655, 801, 805, 808, 912 isotope equilibrium 83, 619, 654 isotope pattern 587, 654 Israel 20, 802 Issinella 638 Italy 150, 164, 182, 230, 236, 389, 418, 460, 498, 512, 520, 530, 542, 554, 694, 702, 780, 781, 786, 787, 788, 797, 819, 849 Iuenna, Carinthia, Austria 906 Iuno 907 Iuvavum/Salzburg, Austria 428, 908 Ivanovia 427, 428
J Jablonskya andrusovi 498 Jaccard similarity index 263, 570 Jamaica 276, 769 James Ross Island 68 Japan 560 Java Sea 669 Jebel Bou Dahar, High Atlas, Morocco 526, 760 Jodotella 692 Joulters Cays 38 Jura Mountains, Switzerland 155 Jurassic 21, 43, 90, 92, 93, 114, 118, 124, 133, 134, 136, 140, 154, 155, 166, 168, 176, 186, 190, 211, 212, 214, 215, 222, 223, 230, 252, 306, 326, 334, 358, 362, 370, 372, 373, 378, 379, 380, 384, 385, 386, 388, 390, 393, 395, 396, 407, 409, 410, 412, 414, 416, 420, 424, 432, 435, 440, 442, 444, 448, 450, 452, 458, 459, 465, 466, 474, 480, 481, 484, 487, 490, 493, 496, 500, 502, 504, 506, 512, 514, 515, 519, 520, 524, 526, 528, 530, 532, 533, 534, 536, 537, I, 542, 544, 546, 547, 553, 554, 556, 558, 559, 562, 563, 566, 568, 569, 570, 576, 597, 613, 644, 661, 664, 696, 682, 678, 684, 694 704, 706, 713, 716, 717, 727, 757, 760, 762, 766, 768, 776, 784, 785, 794, 796, 798, 797, 800, 814, 821, 822, 824, 864, 865, 869, 871, 872, 874, 875, 878, 897, 906, 907 Jurassic scleractinian corals Plate 84 on p. 513
954
Index
K K-feldspar 644 Kabylies, Algeria 850 kalkar 11 kalyptrae 372 kamenitza 13, 730 Kammbühl, Ternitz, Austria 692 Kansas-cycle 808 kaolinite 292, 641, 642, 710 Kapfelberg, Bavaria, Germany 178, 410, 777 Karawanken Mountains, Austria 358, 460, 536 Karlovy Vary (Karlsbad), Czech Republic 736 Karlsbad Sprudelstein 736 Karpinskya 448 karst 11, 13, 14, 55, 120, 125, 203, 205, 232, 310, 371, 390, 661, 663, 730, 732, 806, 819, 877, 878, 896, subsurface 13, 730 karst reservoir 882 karstification 14, 169, 194, 205, 218, 220, 223, 283, 333, 604, 730, 731, 732, 734, 749, 816, 819, 881, 882 Kasaba, southern Turkey 476 Kasimovian (Carboniferous) 344, 366, 510 katagraphs 135 Kazaghstan 877 keep-up system 818 kelp 34, 41, 42, 44, 90, 404, 667, 853, 916 Kendelbachgraben, Austria 464, 798 Kess-Kess mound, Morocco 836, 838, 840; Plate 143 on p. 839 keystone vug 193, 889 Kielce, Holy Cross Mts., Poland 520 Kimmeridgian (Jurassic) 284, 358, 390, 395, 432, 440, 448, 466, 494, 502, 528, 553, 556, 566, 648, 690, 863, 872 Kinsau well, Bavaria, Germany 306, 362, 542, 544, 547, 556, 766; Plate 136 on p. 767 Kircaova, Trabzon, Turkey 470 Kirkuk (Iraq) 880 Klonk, Czech Republic 686 Kohout convection model, dolomitization 332 Koivaella 559, 562, 563, 564, 568 Komia 358, 420 422 Koninckopora 638 Korfu Greece 334, 440, 524, 566, 682, 684, 704 Koskinobullina 385, 562 Kössen Basin (Northern Calcareous Alps) 226, 354, 598, 798 Krahstein, Styria, Austria 140, 502 Kraiburg, Bavaria, Germany 484, 682, 906 Kressenberg, Bavaria, Germany 366 Krumbein and Sloss chart 174, 236, 686, 688 Kungurian basin , Russia 379 Kurnubia 465, 481, 762 Kurnubia palastinensis 466 kurtosis 245, 250, 252 Kuwait 664 L La Paz field, Venezuela 884 Labechiid stromatopores 918
Lacazina limestone, Cretaceous 634 Lacepede Shelf, Australia 515 Lacrymorphus 562 lacustrine 2, 11, 13, 14, 18, 19, 20, 57, 58, 75, 118, 119, 120, 186, 236, 257, 333, 358, 361, 442, 448, 544, 558, 559, 564, 698, 727, 728, 731, 735, 737, 738, 740, 742, 774, 778, 875, 878; Plate 130 on p. 741; Plate 131 on p. 743 carbonates 18, 197, 364, 407 Lacustrine Microfacies Types (LMF) 737, 738, 739, 740, 742, Ladinella 560, 562 Ladinian (Triassic) 157, 182, 380, 385, 438, 553, 560, 566, 688, 710, 781, 782, 784, 872 Ladinosphaera 562 lag deposit 42, 44, 102, 108, 170, 523, 536, 595, 596, 650, 696, 708, 712, 721, 800, 808, 824 lagenid foraminifera 798 lagoon 24, 25, 57, 75, 78, 84, 99, 137, 186, 243, 302, 358, 370, 396, 423, 438, 442, 458, 459, 476, 486, 510, 530, 541, 563, 588, 592, 600, 606, 622, 644, 648, 659, 660, 663, 664, 665, 666, 668, 669, 681, 686, 692, 694, 696, 698, 700, 702, 704, 706, 708, 710, 712, 713, 714, 719, 745, 750, 755, 757, 759, 764, 766, 768, 779, 804, 813, 824, 827, 828, 834, 849, 865, 879, 882, 886, 893 brackish-water 740 coastal 24, 34 evaporitic 704 hypersaline 80, 844 protected 24, 442 restricted 452, 547, 618, 704, 827 lagoonal carbonate 91, 119, 168, 440, 406, 412, 424, 436, 438, 440, 450, 456, 465, 541, 544, 582, 605, 653, 826 lagoonal-reef system 653 Laguna Madre (Mexico) 147 Laguna Mormona, U.S.A. 84 lake 13, 15, 18, 19, 20, 57, 80, 120, 131, 361, 408, 447, 450, 487, 593, 646, 698, 728, 737, 738 brackish 19, 738 dimictic 19 ephemeral 20, 744 freshwater 11, 19, 408, 737, 738 hydrologically closed and open 18, 19 hypersaline 132, 169 meromictic 18, 19 monomictic 19 oligotrophic 740 pelagial zone 19 perennial 20 profundal zone 19 saltwater 2, 11, 16, 19, 84, 735, 737, 738 soda 16, 20, 131, 408, 738, 740 sublittoral zone 19 Lake Balaton, Hungary 20 Lake Chad 153 Lake Clifton 20 Lake Constance, Germany 19, 125, 140 Lake Tanganyika 20 Lake Thetis 20 Lake Van, Turkey 84 Lake Wolaya, Carnic Alps, Austria 382
Index
Lake Zürich and Lake Biel, Switzerland 19 Lamellata wähneri 220 Lamellitubus 559, 562 lamina couplet 56 laminae 14, 19, 56, 57, 132, 137, 185, 207, 377, 500, 597, 598, 615 laminar beds 647 laminar rhizolites 12 Laminaria 916 laminaset 56 laminated loferite 810 lamination 37, 56, 97, 136, 137, 139, 177, 182, 184, 185, 232, 288, 312, 370, 371, 372, 374, 376, 377, 395, 523, 553, 589, 596, 597, 598, 607, 613, 619, 635, 636, 647, 648, 680, 681, 682, 686, 704, 708, 712, 713, 717, 718, 720, 731, 734, 738, 742, 746, 749, 752, 773, 774, 775, 776, 777, 795, 802, 809, 858, 868, 888, 904, 905; Plate 18 on p. 183 laminite 20, 56, 239, 659, 686, 750, 752, 806, 809 laminoid-fenestral fabrics 190, 193 Lancicula 424 Lapeirausia 526 larger foraminifera 181, 182, 612, 616, 628, 632, 635, 886, 888 larvae 384, 593, 384, 533, 593, 606, 616, 684, 832, Lasiotrochus minor 460 Lasiotrochus tatoensis 460 Latemar cycle 808, 812, 813 lateritic 154, 710 Latistraea foulassensis 512 latitude 31, 455, 480, 483, 587, 625, 630, 632, 634, 821, 863, 681, 867 Lattengebirge, Northern Calcareous Alps layer 56, 58 layering 578, 890 leaching 12, 107, 121, 192, 284, 289, 316, 352, 528, 544, 819, 877, 881, 886, 888, 900 lead 892 lead-zinc deposits 379, 730 Lee Stocking Island, Bahamas 376 Leg 133, ODP 2 leiolites 372, 373, 395 Lenticulina 466 Lepidocyclina 208 Lepidocyclina (Eulepidina) eodilatata 476 Lepidocyclinidae 476, 702 Lepidorbitoides 472 Lepidorbitoides socialis 632 Lesser Himalaya, India 343 Letmathe, Sauerland, Germany 384, 458, 500, 519 lettucestone 342 Leugeo hexacubica 483 LF-A fabric 190, 193 LF-B fabric 190, 193 Liassic 4, 215, 226, 393, 440, 444, 466, 496, 524, 560, 661, 808 Libya 182, 414, 478 lichens 192, 898 light 7, 42, 51, 54, 138, 180, 379, 384, 386, 396, 405, 408, 416, 456, 458, 496, 506, 512, 569, 587, 611, 612, 623, 626, 628, 635, 636, 638, 671, 677, 754, 755, 862, 863, 864 lime mud 74, 75, 78, 313, 663, 820
955
limestone 14, 53, 340, 342, 356, 584, 647, 796, 798, 857 limestone classification 253, 380, 578, 580, 833 limestone resource see resource limestone-marl alternation, sequence 211, 215, 757, 803, 808, 862, 866, 867 limestone/dolomite cycle 619 limonite 54, 55, 207 line-counting 247, 254, 255, 256, 261 Lingula 520 Litanaia graecensis 424 lithification 37, 74, 84, 117, 128, 137, 169, 170, 194, 200, 205, 206, 208, 210, 216, 221, 222, 267, 272, 302, 303, 382, 494, 497, 587, 596, 606, 608, 619, 620, 622, 781, 787, 795, 796, 821, 854, 865, 921 Lithiotis 524, 526 Lithistid sponges 492, 496 lithoclast 20, 36, 126, 166, 167, 168, 169, 170, 178, 208, 210, 212, 214, 220, 221, 228, 292, 325, 354, 362, 366, 416, 452, 486, 570, 579, 596, 597, 649, 650, 686, 688, 694, 708, 712, 714, 738, 739, 749, 750, 757, 759, 760, 771, 772, 774, 775, 777, 779, 781, 782, 786, 790, 796, 822, 824, 852, 858, 865, 888; Plate 16 on p. 171 lithoclastic 136, 169, 366, 606 Lithocodium 385, 386, 560, 562, 566, 762 Lithocodium/Bacinella association 385, 563 lithographic limestone 75, 352, 358 lithoherm 41, 769, 856 lithology 53, 54, 60, 61, 62, 205, 484, 652, 805, 815, 847, 878, 884, 903, 904 lithophagid bivalves 390 Lithophyllum 414, 444 lithorelic 236 lithoskel 102, 108 Lithothamnion 413, 414 Lithothamnion glaciale 44, 917 lithotype 63, 885, 896 Little Bahama Bank 50, 786, 856 littoral 9, 19, 389, 684, 745 lituolinid foraminifera 176, 481, 465, 466, 766, 777 Lituosepta 465, 678 Lizard Island, Great Barrier Reef 82, 83, 117 Lizard Island model 82 LMF see Lacustrine Microfacies Type load mark 56 Lochkovian (Silurian) 530, 538, 686 Lofer cycle 524, 704, 789, 806, 808, 809, 810; Plate 140 on p. 811 Loferbachgraben, Austria 598, 678 Loferer Steinberge, Austria 808, 810 loferite 192, 196, 337, 358, 747, 809, 810 algal mat 809, 810 pellet 812 pellet and homogeneous 809 loftusian foraminifera 470 log analysis 659, 884, 885 Logan Pass, Glacier Park, Montana 170, 182 Lombardia 554 Longing Gap, Graham Land, Antarctic Peninsula 483, 484 looseground 189, 542, 606, 608, 610 Lophelia 44, 386, 388, 512
956
Index
Lophelia pertusa 44, 922 Lorenzberg, Kempten, Germany 908 Lorenziella hungarica 488 Lorenziella plicata 488 Lorenziella ruggierii 488 Loser, Styria, Austria 714 lost platform 166, 169, 236, 264, 661, 847; Plate 146 on p. 851 Louisiana U.S.A. 51 low-diversity 630, 663, 740 low-energy environment 4, 28, 37, 76, 207, 339, 352, 357, 361, 396, 464, 523, 588, 591, 594, 598, 669, 680, 694, 698, 805, 806 low-grade metamorphic process 70 Low-Mg calcite (LMC) 11, 18, 19, 20, 43, 68, 70, 71, 74, 83, 144, 149, 210, 267, 270, 271, 284, 292, 296, 299, 308, 321, 326, 390, 404, 405, 408, 412, 418, 514, 519, 520, 522, 526, 529, 533, 534, 619, 684, 753, 854, 867, 874 Lower Saxonia, Germany 390, 410, 448 lowstand sea level (LST) 216, 630, 660, 664, 688, 710, 730, 731, 759, 785, 789, 790, 815, 816, 817, 819, 820, 821, 822, 824, 826, 827, 828 lowstand shedding 821, 822 lowstand systems tract 816, 820, 826 lowstand wedge 688 lucinid bivalves 51, 857, 859, 860 Ludlowian (Silurian) 181 Ludwag, Bavaria, Germany 493, 690 lumachelles 60, 523, 598 luminescence 67, 68, 271 lump *163, 700 Lunz, Austria 409 Lupina, Carpathians, Slovakia 444 Lusitanian Basin, Portugal 168 lysocline 48, 50, 93, 277, 455, 798
M Maastrichtian (Cretaceous) 444, 472, 474, 524, 526, 632, 856, 908 Machel-Hunter model 673 Mackenzie Mts., Canada 868 Mackenziecyathus 918 macro-algae 44, 90, 96 macro-oncoid 123 macroborers 137, 208, 369, 385, 386, 387, 388, 389, 390, 392, 393, 394, 395, 396, 416, 418, 610, 837 macroid 122, *123, *137, 138, 140, 142, 480, 544, 636, 886 macromolecule 81, 83, 102, 104, 369 macrophytes 15, 18, 132, 447, 448, 727, 728, 898 Macroporella 325, 430, 742 macroporosity 65, 911 Macrotubus 559, 562 Madagascar 560, 879 Madonie Mountains, Sicily, Italy 512 Maeandrostia 498 maerl 41, 42, 44, 138, 264, 416 Magdalensberg, Carinthia, Austria 905 magnesian calcite 206, 267, 784 magnesian limestone 328
magnesium 106, 405, 617, 653 Mahengetang Island, Indonesia 153 Mahlknecht Cliff, Dolomites, Italy 688, 784 Majolica, Cretaceous 93 making microfacies types 584 Malaga, Spain 848 Malangsgrunnen Bank 44 Malambaya, Philippines 877, 879, 880 Mallorca, Spain 923 Mammoth Hot Springs, Yellowstone 15, 160 mangal 606 manganese 67, 68, 70, 205, 271, 278, 330, 379, 652, 653, 892, 897 manganese nodule 131, 214, 215, 796 mangrove 170, 606, 656, 668, 748 Marbella Formation, Carboniferous 848, 850 marble 54, 254, 267, 271, 301, 322, 324, 334, 336, 386, 903, 904, 905, 911, 912; Plate 38 on p. 323 descriptive criteria 336, 337 dolomitic 334 fabric 336 fossils in 336 twinning 336 marble inclusion 908 margin, bypass, erosional 660 marginal-marine 6, 8, 24, 26, 169, 374, 377, 526, 540, 647 Marginotruncata 474 Mariazell, Styria, Austria 444 marine environment 11, 14, 6, 26, 275, 299, 301, 304, 386, 404, 408, 412, 450, 455, 544, 551, 642, 658, 875, 881 snow 49, 92 terminology 8 marine-phreatic 68, 275, I296, 297, 304 marine-vadose 206, 274, 275, 276, 296, 301, 304, 663, 710 Marinella 124, 130, 385, 414, 766 Mark Antony 914 marker bed 135, 136, 566 Markhov chain analysis 264 marl 18, 53, 169, 313, 362, 551, 596, 598, 642, 647, 653, 662, 665, 666, 688, 717, 719, 742, 759, 764, 766, 775, 789, 792, 794, 800, 802, 805, 810, 821, 827, 828, 857, 866 marlstone 614 Marmolada, Dolomite Mountains, Italy 164 Mars 16 marsh 13, 364, 659, 663, 748 mass accumulation 59, 424, 540, 547, 551 mass extinction 346, 610, 640, 870, 871 mass flow 184, 221, 234, 340, 660, 688, 698, 771, 772, 779, 821, 847 breccia 229, 230, 233 deposit 196, 229, 257, 597 mass-wasting 710 Mastigocoleus 394 Mastigocoleus testarum 390 mat 184, 387, 647, 714, 838 matrix 59, 61, *73, 76, 96, 104, 113, 118, 135, 169, 172, 177, 198, 241, 259, 260, 272, 274, 348, 349, 350, 352, 358, 361, 362, 366, 460, 523, 530, 548, 576, 577, 579, 580, 589, 591, 592, 596, 641, 652, 655, 670, 673, 773, 790, 795,
Index
796, 798, 826, 833, 834, 837, 848, 863, 864, 874, 888, 889, 904, 905 burrowed 682 calcisiltite 666 composition 232, 837 definition 73 dolomite 197, 289, 332 fine-grained 73, 80, 96 inhomogeneous 96 laminated 96 microcrystalline 80, 356 organic 102, 103 peloidal 96 siltitic 96 type 236, 576, 578, 580, 590, 680, 888 matrix porosity 882 matrix-support 61, 229, 350, 523, 530, 772 Matthew et al. charts 257, see CD maturity 175, 260, 712, 718 maturity level 587 Maurach, Tyrol, Austria 520 maximum flooding surface 818, 820, 821, 826, 828 mean 250, 779, 790, 872 deviation 255 diagram 252 size (Mz) 245 measurement error 247 mechanical compaction 58 mechanical concentration 211 mechanical destruction 462, 580 mechanical property 224 median 245, 253 Mediterranean 11, 12, 14, 20, 40, 41, 42, 80, 89, 138, 140, 170, 207, 215, 465, 484, 524, 620, 626, 796, 797, 881, 836, 914 megablock 772, 780 megabreccia 178, 228, 229, 230, 354, 688, 772, 780, 781 megacement 298 megacyclothem 694 Megalodon 524, 809 Megalodon limestone, Triassic 523 megalodontids 523, 524, 810 Meischner Sequence, Model 773, 774, 775, 778 melobesioid coralline algae 414 meltwater 14, 42 Member A, Lofer Cycle 809, 810 Member B, Lofer Cycle 809, 810, 812 Member C, Lofer Cycle 809, 810, 812 Menorca Island, Spain 185, 538, 544, 546, 682, 850 mercury injection capillary pressure analysis 282, 890 Meseta, Morocco 500 meso-oncoid 123 mesoclot 371, 395 Mesogastropoda 529 mesohaline 836 mesolimnion 18 mesopelagic zone 9 Mesophyllum 414 mesopore 516 Mesoproterozoic 868 Mesorbitolina 468 mesotrophic 49, 379, 623, 624, 670 Mesozoic 4, 7, 49, 95, 98, 101, 113, 121, 200,
957
220, 274, 297, 376, 378, 379, 382, 389, 392, 393, 406, 408, 410, 412, 418, 430, 444, 447, 450, 452, 480, 486, 491, 504, 506, 507, 508, 515, 522, 524, 529, 533, 534, 544, 546, 547, 550, 551, 556, 558, 566, 568, 570, 591, 616, 628, 635, 653, 654, 659, 676, 702, 712, 717, 718, 719, 793, 795, 796, 831, 838, 869, 881 Messopotamella 562 metabolic 75, 103, 379, 512, 616, 623, 875 metal accumulation 892 metamorphic carbonate 267, 322, 334 metamorphic spar 322 metamorphism 54, 194, 279, 334, 336, 337, 538 metasomatic alteration 484 metasparite 325 metazoans 102, 393, 491, 612, 659, 831, 868, 918 meteoric 13, 14, 95, 106, 108, 205, 222, 271, 274, 275, 276, 277, 299, 301, 304, 310, 323, 333, 368, 647, 655, 710, 818, 881, 923 meteoric cap 817 meteoric overprint 766, 807 meteoric-marine mixing zone model, dolomitization 332 meteoric-phreatic 218, 220, 276, 296, 300, 734, 762 meteoric-vadose 38, 276, 296, 301, 304, 310, 314, 346, 663, 710 methane 51, 374, 842, 857, 858, 859, 860 methane seep 859 methanogenesis 369, 858 methylmetacrylate 899, 900 Mexico 12, 14, 24, 28, 31, 51, 252, 490, 528, 600, 602, 624, 800, 880 Meyendorffina 465 Mfamosing limestone, Cretaceous 414 MFT criteria 581 Mg-cage 95 Mg-calcite 24, 43, 66, 70, 74, 80, 83, 94, 117, 149, 268, 270, 274, 289, 292, 296, 306, 371, 502, 506, 507, 514, 540, 548, 620, 653, 837 Mg/Ca ratio 13, 270, 292, 299, 330, 412, 619, 874 mica 732, 911 Michigan, U.S.A. 880, 884 micrite 6, 16, 18, 65, 66, 68, 71, 73, *74, 75, 92, 94, 95, 96, 97, 164, 178, 194, 282, 311, 339, 344, 346, 349, 352, 354, 356, 357, 358, 361, 362, 364, 366, 368, 369, 371, 372, 374, 379, 395, 428, 510, 534, 553, 560, 589, 624, 643, 652, 653, 654, 655, 673, 674, 677, 680, 692, 694, 698, 704, 708, 716, 718, 726, 727, 735, 738, 746, 750, 757, 759, 773, 775, 777, 780, 790, 797, 802, 805, 807, 810, 822, 830, 835, 838, 854, 859, 863, 868, 873, 874, 880, 888; Plate 6 on p. 77; Plate 7 on p. 79 accretionary organomicrite 82, 83 allochthonous 75, 88 allomicrite 75, 81 aphanitic 84 automicrite 75, 81, 83, 93, 99, 116, 370 bioclastic micrite 90 cement 74 classification of 75 clasts 76, 170, 362, 448, 478, 596 clotted 372, 726, 787 container organomicrite 82, 84 crystal size of 75
958
curtain 95 definition 76 dense 372 diagenetic 92 formation 78, 80, 92, 875 grains 344 homogeneous 76 inorganic 80 internal micrite 76, 96, 99, 579 matrix 93, 298, 301, 361, 362 microbial 7, 76, 84, 341, 364, 700 minimicrite 78, 84 nannomicrite 81, 91 organomicrite 81, 82, 83 orthomicrite 75 pelagic 99 peloidal 117, 118, 178, 838 primary 75 rim 120, 207, 725 seafloor 76, 96, 578, 579 secondary 75 texture 371, 578 micrite/grain ratio 589 micrite clast 678, 700, 708, 714 micrite envelope 6, 37, 68, 100, 114, 118, 119, 120, 121, 304, 308, 312, 371, 390, 533, 680, 694, 714, 781 micritization 37, 38, 81, 94, 104, 106, 111, 116, 118, 119, 120, 144, 150, 163, 252, 271, 289, 325, 350, 410, 694, 698 micro-encrusters 372, 384, 385, 386, 624, 635, 636, 714, 837 micro-fault 640 micro-framework 690, 830, 844 micro-oncoid 123, 739, 757 microalgae 370 microbes 6, 82, 90, 99, 113, 121, 127, 164, 184, 214, 268, 343, 352, 369, 370, 371, 372, 373, 376, 377, 378, 379, 381, 382, 384, 385, 390, 504, 520, 541, 559, 624, 625, 636, 663, 690, 694, 706, 713, 714, 717, 771, 785, 788, 830, 831, 832, 833, 837, 838, 842, 860, 871, 893, 898 microbial 6, 97, 194, 440, 496, 648, 659, 833, 838, 857, 869; Plate 8 on p. 87 microbial activity, contribution 6, 12, 15, 56, 80, 116, 207, 274, 306, 344, 370, 371, 374, 376, 377, 378, 380, 440, 576, 596, 835, 842, 858, 874, 875 microbial carbonate 207, 340, 352, 370, 371, 373, 376, 378, 379, 395, 541, 578, 605 classification 371, 833 describing 370 terminology 369 microbial coating 12, 236, 289 microbial crust 4, 207, 218, 220, 325, 346, 374, 377, 379, 493, 498, 541, 576, 824 microbial decay 90, 612 microbial grains 163 microbial lamination 131 microbial mat 25, 55, 184, 188, 197, 205, 316, 369, 370, 377, 408, 410, 560, 618, 619, 624, 704, 746, 747, 749, 750, 759, 810, 868 microbial micrite 78
Index
microbial peloid 114 microbial precipitation 578, 873 microbial reef 136, 378, 393, 520, 523 microbial thrombolite 372 microbialite 6, 16, 75, 76, 82, 96, 150, 182, 342, 364, 371, 372, 373, 374, 377, 379, 380, 385, 395, 396, 578, 589, 620, 785, 844 classification 373 laminated 374 macrostructure, megastructure, microstructure 373 occurrence 377 microbolite 371 microborers 37, 38, 90, 104, 114, 119, 120, 145, 163, 369, 386, 387, 390, 393, 394, 395, 396, 416, 418, 611, 612, 624, 635, 636, 694, 762, 875 microboring 38, 90, 91, 114, 119, 120, 121, 149, 166, 208, 211, 306, 388, 524, 554, 610, 757, 837, 888 microbreccia 228, 230, 314, 686, 714, 719, 720, 731, 732, 824, 834, 905 microcavity 84, 117 microchannel 74 Microcodiaceae 564 Microcodium 12, 14, 426, 559, 562, 726, 727, 728 Microconchus 530 microcoprolite 547 microcoquina 538 microcrack 12, 67, 132, 224, 225, 226, 303, 310, 640, 682, 727, 883, 898, 900 microcrystalline calcite 38, 73, 119, 358, 592 microdolomite 274, 297 microendoliths 104, 119 microenvironment 84, 261, 369, 377 microfabric 67, 96, 99, 106, 117, 370, 372, 373, 379, 568, 740 clotted micropeloidal 84 peloidal 82 porostromate 373 pseudonodular 12 spongiostromate 124 thrombolitic 784 microfacies 1, 198, 575, 595, 652, 718, 735, 752, 798, 803, 822, 824, 833, 842, 890, 912 depositional 268, 805 diagenetic 268, 805 practial use 877 reef studies 833 small-scale 648 microfacies analysis 1, 7, 63, 66, 203, 243, 480, 641, 759, 786, 824, 847, 849, 874, 896, 905 microfacies boundary 850 microfacies concept 1 microfacies criteria 4, 6, 380, 643, 655, 828, 868, 878, 896 microfacies data 66, 71, 73, 251, 671, 680, 725, 802, 884, 885, 895 microfacies distribution 827 microfacies diversity 827 microfacies log 61 microfacies monographs 402 microfacies relay 724 microfacies sampling 63, 648 microfacies temporal changes 874
Index
microfacies types 2, 575, 576, 578, 581, 582, 584, 652, 657, 681, 719, 732, 824, 827, 828, 834, 848, 890, 893, 906, 908 allochthonous carbonates 576 carbonate ramp 716 distribution 826 dynamic 818 lacustrine 737, 738, 744 pictures 571 time-specific 575 microfacies variability 63, 235, 575 microfacies zone 221 microfenestrae 74, 312 microfissure filling 226 microfold 648 microfossils 14, 50, 63, 66, 128, 233, 299, 312, 370, 372, 376, 399, 408, 452, 484, 490, 570, 638, 672, 688, 716, 731, 774, 779, 796, 837, 848, 865, 908 in ooids 145, 146 planktonic 481, 680, 684, 808 microfracture 67, 224, 225, 230, 530, 714; Plate 25 on p. 227 microfracture index 898 microgranular 105, 454 microid 123 microkarst 205, 742 microlamination 614, 739 Microlaminoides 562 microlithoclasts 168 micromold 74 micropaleontological sample 64 micropeloid 684, 796 microporosity 65, 66, 74, 78, 278, 688, 736, 800, 880, 881, 888, 898, 911 microproblematica 4, 124, 233, 382, 384, 385, 386, 399, 404, 444, 450, 559, 564, 566, 589, 688, 762, 837, 908 discussion 560 generic list 559, 562 Paleozoic microproblematica Plate 98 on p. 565 planktonic 566 Mesozoic microproblematica Plate 99 on p. 567 microquartz 484 microrelief 205, 212 microsampling 71, 652, 654 microscleres 494 microsequence 65 microspar 73, *74, 76, *94, 95, 96, 97, 127, 182, 321, 322, 325, 371, 538, 643, 653, 710, 760, 858, 888 microsparite 12, 76, 94, 708, 726, 728 microsparitization 97 microsparstone 361, 362 microstromatolite 4, 134, 136, 212, 214, 620 microstructural types, key 522 microstructure 67, 104, 105, 106, 120, 125, 174, 371, 372, 373, 377, 395, 401, 402, 453, 500, 509, 514, 519, 522, 524, 526, 528, 529, 530, 534, 537, 538, 540, 548, 550, 655 belemnite 534 bivalves 524 brachiopods 520 crossed 518
959
crossed-lamellar 518, 522, 529, 540 foliated 518, 522, 524 homogeneous 518, 522 laminar 518 nacreous 518, 529 normal prismatic 529 of shells 518 ostracods 544 prismatic 518, 522, 530, 544 shells 518 spherulitic 518 microstylolite 239, 314, 316, 318, 362 microtaphofacies 106, 108 Microtubus 4, 559, 562 Microtubus communis 566 microvug 74 migration pathway 884, 889 Milankovitch cycle, spectrum 194, 660, 804, 805, 807, 813, 817 miliolid foraminifera 14, 164, 304, 442, 454, 455, 456, 458, 460, 465, 466, 470, 472, 476, 478, 481, 566, 616, 618, 628, 663, 668, 672, 677, 678, 702, 713, 719, 725, 820, 826, 828 886 miliolite 14 milleporid hydrozoans 508 Miloszewo well, Poland 314, 704 mineral deposit 730, 734, 770, 892, 895 mineralization 203, 205, 206, 207, 226, 320, 370, 376, 857, 877, 892, 893 mineralogy 53, 64, 99, 104, 267, 272, 277, 314, 484, 591, 641, 652, 653, 769, 831, 864, 866, 874, 888, 890, 900 nonskeletal 874 precursor 867, 874 skeletal 313, 874, 875 temporal changes 50, 874 Miniacina 381, 476, 917 minimicrite 75, 78, 97 Minnesota Lakes 19 Miocene (Tertiary) 41, 140, 266, 372, 373, 381, 394, 411, 414, 418, 423, 444, 448, 474, 476, 478, 480, 512, 516, 524, 530, 541, 542, 544, 560, 564, 612, 630, 632, 708, 728, 731, 735, 739, 740, 754, 756, 870, 871, 878, 881, 886 Miogypsina 476 Mirifusus minor 483 Mishrif Formation, Middle East 316, 374, 414, 468, 472, 523, 630, 740, 881, 883 Mississippi Valley Type mineralization 332, 379, 892 Mississippian 146, 155, 362, 379, 496, 516, 529, 550, 554, 608, 638, 661, 675, 717, 871 Missouri, U.S.A. 892 Missourian (Carboniferous) 190 Mistelgau, Franconian Alb, Germany 537 Mitcheldeania 409 mixed carbonate-siliciclastic sequence 362, 368, 597, 862, 865, 866, mixed-layer minerals 277, 613, 641 mixing zone 217, 222, 274, 277, 300, 308, 643, 748 mixohaline 541 mixotroph 754 Mizzia 236, 430, 432, 436, 444, 584, 844 Mizzia velebitana 432 modal analysis 66, 175, 262, 263, 578, 580
960
modal diameter (Mo) 245 model 884 3D model 878, 882, 884, 885, 890 beach-barrier island-lagoon model 24 beach-strandplain model 24 carbonate ramp 657, 660, 664, 675 computer 657, 658 depositional 658, 681 dynamic 657 facies 660 rimmed carbonate platform 721, 723 tidal flat island 669 Wilson 657, 660, 661 Wilson modified 660 X-Y-Z model 657 modeling 658, 659, 670, 815, 866, 877, 878 Mokkatam Group, Tertiary 899 molar-tooth calcite 868 Molasse zone 240, 264, 316 mold 105, 106, 116, 270, 274, 310, 329, 402, 430, 497, 514, 523, 529, 530, 533, 534, 748, 795, 854 mold peloid 113 moldic porosity 880, 886 molechfor grain association 627 Molignon, Dolomites, Italy 498 mollusk coquina 42 mollusk shells 37, 60, 304, 308, 346, 366, 390, 428, 524, 578, 617, 619, 630, 900 mollusks 41, 44, 51, 59, 71, 74, 97, 111, 249, 252, 271, 274, 290, 350, 381, 382, 400, 404, 518, 533, 537, 540, 546, 556, 559, 593, 596, 597, 607, 614, 616, 617, 618, 626, 628, 636, 654, 663, 666, 667, 668, 670, 672, 684, 692, 694, 714, 717, 742, 748, 754, 760, 771, 782, 795, 798, 820, 828, 835, 842, 858, 865, 869, 875, 912 Moltkia isis 854 Monastery St.Anthony, Egypt 713 Mono Lake, Death Valley, California 20 Montagne de Tauch, Département Aude, France 468 Montagne Noire, France 195 Montastraea 601 Monte Bolca, Verona, Italy 530, 692 Monte Camposauro, Apennines, Italy 150 Monte Generoso, Tessin, Switzerland 534 Monte Kumeta, Sicily, Italy 694 Monterey Bay, California 51 moonmilk 13 Mooshuben, Styria, Austria 352, 444 Moravian Karst, Czech Republic 673 Mormorion, France 644 Morocco 395, 500, 650, 784, 836 Morozovella 474 morphometric data 173, 174, 176, 175 mortar 906, 908 mosaic 74, 362, 903, 905, 907, 908; Plate 151 on p. 909 mosaic material 682, 905 Moscovian (Carboniferous) 422, 436, 460, 462, 566 moss 14, 16, 19 Mössbauer spectroscopy 911, 912 mottled fabrics 186, 188, 364 mound 21, 41, 344, 381, 428, 444, 496, 507, 510, 526, 566, 608, 630, 637, 638, 665, 666, 686,
Index
696, 717, 754, 769, 830, 831, 833, 834, 836, 838, 842, 844, 852, 854, 856, 857, 860, 874, 877, 879, 892 algal 700 bryozoan 852, 856 cold-water 856 coral 854 crinoidal 551 seep 858 mound type 834 Mouse Mts., Alberta, Canada 516 MSC triangular diagram 834 Mt. Agnello, Southern Alps, Italy 230 mucilaginous sheath 84, 112, 369, 404, 408, 606 mud 7, 24, 25, 50, 58, 80, 88, 89, 94, 95, 99, 113, 118, 147, 340, 342, 349, 350, 352, 354, 426, 442, 480, 494, 532, 605, 606, 610, 630, 644, 647, 653, 659, 663, 666, 676, 713, 714, 773, 774, 788, 790, 795, 798, 817, 852, 879 aggregate 113, 302 aragonite lime mud 80, 95, coccolith mud 91 Globigerina mud 91 hemipelagic 662 lagoonal 94, 95 Mg-calcite mud 80, 91, 95 mud budget 80, 95, 98, 210 mud peloids 113 mud precursor 99 pteropod mud 91 siliceous 483 terrigenous 659 mud accumulation 770 mud bank 656 mud clasts 356 mud compaction 194 mud crack 58, 177, 647, 746, 747, 750 mud facies 111, 764 mud flat 20, 669, 770 mud grains 113 mud intraclast 749 mud mound 6, 26, 75, 76, 82, 84, 116, 196, 197, 370, 379, 380, 422, 428, 436, 459, 460, 496, 548, 552, 598, 624, 662, 665, 666, 675, 696, 700, 716, 717, 719, 770, 771, 788, 830, 833, 834, 836, 837, 838, 843, 868, 870, 871, 880, 885, 892; Plate 143 on p. 839; Plate 144 on p. 841 mud mound factory 873 mud peloid 113, 806 mud polygon 746 mud production 80, 90,95, 465, 842 mud-support 344, 348, 349, 350, 352, 366, 884, 897, 904; Plate 43 on p. 351 mud/grain ratio 76, 96 mudstone, lime mudstone 25, 56, 74, 75, 190, 196, 214, 312, 313, 320, 348, 349, 350, 352, 358, 362, 364, 452, 484, 534, 593, 595, 596, 597, 614, 622, 637, 640, 641, 644, 659, 662, 663, 665, 666, 675, 676, 684, 686, 694, 706, 708, 717, 719, 737, 739, 742, 744, 745, 747, 748, 750, 752, 759, 764, 766, 777, 786, 789, 796, 801, 806, 808, 810, 824, 826, 834, 837, 840, 841, 848, 850, 865, 868, 882, 884, 898, 921; Plate 44 on p. 353
Index
961
argillaceous 366 fenestral 732, 762, 810, 826, 827 laminated 810 pelagic 800 Muleshoe Mound, Sacramento Mountains, New Mexico, 362, 696, 836, 842; Plate 144 on p. 841 Multimurinus 510 Multithecopora 344 Multithecopora syrinx (344 multivariate analysis 247, 250, 254, 262, 263, 264, 581, 652, 672 multivariate discriminations of MFT 580 multivariate studies 263 mummies 121 Munsell Color System 54 Muranella 559, 562 Murcia region, Southern Spain 534 Murgabian (Permian) 178, 564 Murray Basin, Australia 669 Muschelkalk, Triassic 46, 76, 116, 390, 392, 394, 524, 598, 640 muscovite 732 Mya truncata 44 mycorrhizae 727 Myodocopida, ostracods 544 Myogypsina 886 Myriophyllia rastellina 512 mytilid bivalves 389, 859
N Na 330, 619, 653 nacreous shell structure 105 Nagelfluh 242 Namurian (Carboniferous) 638 nanno-organisms 78, 91, 93 nannoagorite 75, *99 nannobacteria 371 nannoconids 91, 490, 684, 797 nannocrystals 120 nannofor grain association 627 nannofossils 49, 66, 76, 97, 684, 796, 797, 800, 801, 802, 875, 880 nannomicrite *74, 78, 91 nannoplankton 40, 76, 91, 490, 534, 798 nasselarian radiolarians 482, 483, 484, 797 nassellarian/spumellarian ratio 486 Nassfeld, Carnic Alps, Austria 420, 450, 460 nautilid cephalopods 533, 534, 544, 794, 795, 796, 842; Plate 90 on p. 535 Nayband Formation, Triassic 395, 508 near-coast 19, 456, 496, 540, 541, 559, 560, 572, 579, 604, 641, 670, 742, 862, 863, 879 nektobenthonic 538 nekton 47, 662, 666, 762 nektoplanktonic 528 Neoanchicodium 428, 582, 584 Neogastropoda 529 Neogene 122, 802, 852, 869, 872 Neoiraqia convexa 468 Neomeris (Larvaria) sp. 444 Neomeris cretacea 442 neomorphic alteration 253, 268, 308, 325
neomorphic process 73, 321, 652 neomorphic replacement 298 neomorphic spar 322, 923 neomorphism 68, 94, 267, 271, 290, 292, 303, 308, 321, 322, 324, 325, 350, 364, 514, 642 aggrading 321 coalescive 321 crystal size 328 degrading 321 Neoproterozoic 143, 188, 362, 372, 868 Neoschwagerina 226 neoschwagerinid fusulinids 462, 464 Neoteutloporella 430, 434, 444 Neoteutloporella socialis 432 Nephrolepidina praemarginata 702 neptunian dike 211, *212, 217, 218, 220, 221, 222, 223, 298, 522, 553, 794, 796, 801, 849; Plate 24 on p. 219 neptunian sill 222 neritic 10, 653, 757, 873 Nesza quarry, Budapest, Hungary 190 Netherlands Antilles 125 Neuropora 502, 785 Neutron Activation Analysis (NAA) 71, 911, 912 neutron log, neutron-density log 884, 885 Nevada , U.S.A. 197 New Mexico, U.S.A. 222, 560, 836, 843, 844 New Zealand 40, 308, 515, 628, 863 Newfoundland, Canada 42, 857 Nezzazata 472 Nicaragua 821, 863 Nigeria 414 Niphates digitalis 491 nitrate fixation 624 nitrogen 189, 588, 623, 624 nodosariid foraminifera 798 nodular 177, *198, 200, 201, 202, 647, 648, 716,727, 728, 738, 796, 857, 801,857, 889, 904; Plate 21 on p. 199 nodular limestone 200, 794, 795, nodule 198, 214, 373, 409, 417, 562, 622, 726, 728, 795, 798, 889 formation, growth 200 nomogram 247 non-autotrophic bacteria 386 non-carbonate 54, 58, 63, 96, 198, 339, 591, 653, 895, 896 non-deposition 56, 60, 203, 210, 302, 624, 698, 712, 795, 820 non-fabric selective pores 280, 283 non-laminated microbialite 374 non-marine 8, 10, 11, 14, 15, 361, 364, 450, 680, 716, 725, 738 non-marine, carbonate classification 361; Plate 48 on p. 365 non-skeletal 122,194, 868, 873 non-tropical 4, 30, 31, 40, 267, 515, 625, 628, 632, 830 Nördlingen Ries impact crater 364, 731 Norian (Triassic) 4, 178, 182, 190, 218, 236, 239, 262, 344, 358, 374, 380, 382, 393, 424, 438, 501, 508, 534, 556, 559, 564, 566, 577, 615, 677, 690, 694, 702, 704, 706, 714, 790, 789, 810, 812, 835, 836, 871
962
Index
normal grading 54, 177 normal prismatic shell structure 105 North Africa 478, 507, 533 North America 125, 507, 795, 800, 803, 808, 842, 859 North Atlantic 394 North Sea 51, 99, 524, 669, 800, 877, 880, 884 Northern Calcareous Alps 252, 385, 394, 497, 514, 522, 524, 530, 542, 556, 557, 570, 614, 677, 682, 686, 688, 698, 700, 708, 760, 786, 789, 792, 797, 798, 810, 866, 906 Norwegian shelf 42, 138, 917 Nötsch, Austria 460, 516 Nova Scotia 42 Nowakia 537, 538 Nubecularia, nubeculariid foraminifera 134 382 nuclear magnetic resonance 885 nucleus 137, 482 of ooids 143 of rhodoids 139 nuculanid bivalves 859 Nuia 366, 372, 426, 559, 564, 566 number-frequency measurement 256 Nummufallotia cretacea 632 Nummulites 181, 182, 455, 476, 478 Nummulites deserti 181 Nummulites fraasi 181 nummulitid foraminifera 182, 366, 476, 481; Plate 74 on p. 476 nutrients, feeding, food 6, 18, 40, 43, 47, 82, 111, 117, 125, 180, 185, 186, 196, 210, 341, 348. 379, 380, 384, 385, 386, 405, 408, 428, 450, 455, 458, 484, 487, 492, 496, 512, 520, 559, 570, 587, 588, 591, 593, 605, 607, 610, 613, 623, 624, 625, 634, 636, 637, 677, 753, 754, 755, 756, 757, 769, 779, 798, 816, 842, 856, 863, 864, 873, 874 O Oberalm Formation, Tithonian 93, 252, 484, 786, 797 Obersalzberg, Germany 239 Obir Mt., Austria 135, 354 obliquity 804, 805 ocean atmosphere, chemistry, circulation 71, 602, 612, 635, 654, 760, 869 ocean floor 10, 75, 93, 793, 800 oceanic atoll 27 oceanic basement 794, 800 oceanic ridge 46, 654 octocorals 104, 509, 601, 604, 853, 854, 856 Oculina 853 offshore 19, 28, 665, 666, 821, 880, 879 oil and gas reservoir 6, 320, 528, 880 oil shale 20 Öland Island, Sweden 186, 212, 516, 548, 795 Olang Dolomites, Southern Alps, Italy 164, 389, 520 Old Red Continent 738, 796 Oligocene (Tertiary) 2, 394, 414, 476, 483, 702, 766, 872, 880, 886 oligophotic 755 oligotrophic 49, 51, 379, 447, 458, 506, 623, 624 Olympic Peninsula, U.S.A. 859 OM (organic matter) 270
Oman 182, 249, 382, 399, 418, 465, 483, 514, 523, 528, 563, 564, 566, 678, 677 omission 204, 205, 207, 208, 210, 212, 214 Omphalocyclus 908 oncoid 4, 6, 21, 36, *100, 111, 114, *121, 123, 126, 127, 128, 130, 131, 132, 135, 136, 137, 157, 158, 163, 164, 177, 178, 186, 212, 248, 251, 257, 306, 316, 354, 356, 358, 360, 373, 377, 382, 384, 408, 409, 420, 436, 440, 480, 530, 579, 589, 624, 636, 649, 650, 655, 678, 680, 694, 696, 712, 714, 717, 719, 738, 739, 742, 749, 762, 766, 777, 806, 807, 812, 813, 820, 821, 848, 866, 886, 888, 919; Plate 11 on p. 129 agglutinated 706 algal 875 Archaeolithoporella 135 Bacinella 140 brackish-water 136, 140 classification 123, 127 composite 123, 124, 126 cortex of 126, 130 deeper-water 133 ferromanganese, ferruginous 131, 636 field criteria 126 fluvial 125 foraminiferal 124, 126, 128, 130 freshwater 21, 125, 131, 132, 136, 619, 739, 740, 742 geometric types 130 Girvanella 122, 124, 135, 136 goethite 214 growth 126, 128 high-energy 589 lacustrine 21, 125, 132, 140, 740 lamination 125, 126, 127, 130, 717 Late Jurassic oncoids 133 marine 125, 132 micrite 124, 140, 589, 717 micro-oncoid 742 microbial 380, 875 non-carbonate 130, 131 non-laminated 130 Nubecularia 128 nucleus 126, 130 paraporostromate 122 pelagic 134 phosphatic 131, 636 porostromate 122, 124, 126, 127, 132, 134, 135, 136, 140, 696, 714, 824, 875 red algal 612 reef oncoids 134 reworked 131 saline 131 shape 126 significance 135, 136 size 126, 127, 137 spongiostromate 21, 122, 124, 125, 126, 127, 132, 136, 706, 875 terrestrial 125 two-stage 140 types 126 zoogenic 124, 757 zoophytogenic 12
Index
oncolites 123, 132, 134, 135, 136, 137, 366 Oncolithi 122 onlap 759, 816 Oocardium 376 oogonium, charophycean algae 19, 448, 450, 739 ooid 4, 6, 14, 20, 21, 36, 38, 41, 68, 71, *100, 111, 116, 121, *142, 143, 144, 145, 146, 147, 148, 149, 150, 152, 153, 154, 156, 158, 161, 163, 164, 174, 176, 248, 251, 257, 270, 284, 314, 316, 328, 334, 356, 357, 358, 360, 362, 366, 382, 452, 459, 470, 579, 589, 591, 606, 616, 626, 628, 655, 663, 665, 666, 680, 681, 694, 696, 698, 712, 713, 714, 717, 718, 719, 738, 740, 749, 757, 760, 762, 766, 774, 785, 789, 806, 812, 820, 821, 822, 848, 850, 852, 863, 865, 866, 868, 873, 874, 875, 888, 890, 900, 913, 914; Plate 13 on p. 151 allochthonous 156 aragonite 145 asymmetrical ooid 147, *148 autochthonous 156 benthic 152 blackened 712 broken 148, 589, 698 calcite 144, 147 cerebroid ooid *148, 150 compound (complex) 143 concentric *144, 589, 698, 717 cortex 153, 698, 717 deep-water 147 deformed 148 depositional setting 155, 156 diagenesis 149, 157 differences saltwater/freshwater 21, 125, 156 dissolved 314 distorted ooid 145, *148, 589, 619 dolomite 150, 153 eccentric 148, 589 ferruginous 208, 828 fitted 152 formation 142, 143, 147, 149, 153, 155, 827 giant 143 Great Salt Lake 153 half-moon 145, 148, 178, 619, 622 high-energy 357 High-Mg calcite 144, 152 hypersaline 147 iron 143, 153, 154, 712, 892 Low-Mg calcite 144, 152, 154 marine 147, 154 micrite ooid *143, *144, 150, 589, 698, 717, 762 microbial 150, 152, 371 microfabric 143, 144, 145, 156 modern 145, 146, 147 non-carbonate 153 non-marine 146, 147, 154 normal 38, *143, 153 nucleus/cortex ratio 149 pedogenic 364 pelagic 147, 796 phosphatic 153, 650 porosity 149 quiet-water 146, 147, 591, 648
963
radial ooid 143, *144, 150, 589, 680, 698, 717, 738, 822, 850 radial-fibrous 143, 144, 145, 150, 589, 619 regenerated 148 secondary radial structure 144 shoal 146, 358, 607, 848, 850 shrunken 148 significance 155, 156 siliceous 153 size 589 spiny 145, *148 superficial 38, *142, 143, 150, 153, 284 tangential 38, 143, 144, 145, 150, 589, 850 terminology 142 types 144, 145, 149 oolite 142, 146, 607, 698, 723, 725, 762, 849, 900 oomicrite 357, 358 oomold 149, 150, 152, 283, 766, 888, 900 oosparite 357, 358 ooze 47, 49, 50, 75, 272, 474, 662, 769, 793, 800 opal, opaline 277, 483, 496, 504 open-marine 436, 444, 450, 457, 508, 524, 526, 533, 537, 538, 540, 548, 551, 553, 556, 559, 562, 608, 672, 692, 759, 764, 817, 822, 826, 827 open-space fabric 6, 59, 63, *190, 197, 636, 680, 842, 905 Operculina 476 ophicalcite 800 ophiolite 800, 801 ophiurids 352, 534, 550, 558 Ophthalmidium 678 Ophthalmidium carinatum 678 Ophthalmidium triadicum 678 ophthalmocline 611 Opisthobranchia, gastropods 529 orbital forcing 792, 804, 805 orbitoid foraminifera 466, 472, 630, 632 Orbitoides 455, 472 Orbitoides apiculata 632 Orbitolina 454 Orbitolina (Mesorbitolina) aperta 468 Orbitolina (Orbitolina) duranddelgai 468 orbitolinid foraminifera 458, 465, 468, 470, 472, 686, 719 Orbitolites 108 Orbitopsella 465 Orbulina 474, 819 ordination analysis 677, 723 Ordovician 134, 192, 197, 211, 372, 378, 384, 390, 392, 416, 422, 424, 426, 438, 459, 496, 500, 507, 510, 515, 519, 520, 537, 540, 546, 550, 558, 563, 564, 566, 596, 597, 628, 630, 664, 684, 717, 795, 802, 814, 831, 842, 852, 869, 871, 872, 879, 884, 918, 919 ore 877, 911 deposit 6, 156, 206, 379, 732, 892, 893 formation 226 minerals 484, 893 organic acids 120, 170 organic carbon 198, 614, 623, 624, 641, 648, 656, 789, 808 organic carbon/sulfur ratio 612, 614 organic detritus 483, 607
964
organic film 163, 320 organic material 56, 57, 73, 168, 193, 302, 370, 623, 646 organic matter 19, 54, 67, 82, 84, 94, 97, 101, 112, 117, 118, 120, 121, 149, 169, 170, 189, 198, 200, 248, 268, 270, 279, 295, 301, 308, 313, 337, 369, 504, 541, 605, 606, 607, 610, 613, 623, 624, 641, 646, 649, 656, 662, 706, 737, 739, 764, 805, 820, 837, 858, 866, 884, 892, 893, 897, 899 organic petrography 67, 656 organic reef 576, 830, 831 organic tissue 370, 371, 533 organism-sediment interaction 606 organoclast *101 organomicrite *83, 374 organomineralization 81, 369, 873, 875 organosedimentary deposit 372 orientation 181, 523, 596, 779 of clasts 232 of fossils 182, 184, 593 of samples 62, 63 of shells 180 pattern 182, 590 origination rate 435, 447 Orthida, brachiopods 519 Orthoceras 795 Orthoceras limestone 211, 212, 534 orthochemical rock 356 orthocone cephalopods 182 orthoconglomerate 239, 240 orthomicrite *75 orthosparite 324, 325 Ortonella 409, 410 Osage County, Oklahoma, U.S.A. 420, 694 osagid grains 121, 124 oscillatoriacean cyanobacteria 16, 18, 409 Oslip, Burgenland, Austria 414 Osteocrinus rectus 556 Osterhorn Mountains, Austria 798 Ostermünchen, Bavaria, Germany 382 Osterwald, Lower Saxony, Germany 714 ostracod mudstone 349 ostracod wackestone 452, 544 ostracods 4, 18, 19, 56, 74, 105, 106, 113, 155, 181, 186, 214, 220, 230, 240, 260, 301, 364, 404, 442, 448, 452, 518, 534, 538, 542, 544, 546, 598, 611, 612, 614, 616, 618, 619, 622, 628, 630, 635, 663, 675, 700, 704, 706, 717, 719, 731, 738, 739, 740, 742, 748, 749, 750, 752, 782, 785, 794, 795, 796, 798, 800, 809, 810, 820, 822, 824, 826, 827, 828, 875, 918, 919, 921; Plate 93 on p. 545 entomozoid pelagic ostracods 4, 544, 546 microstructure 544 ostracum 524, 526, 529 ostreid bivalves 452, 524 Ostreobium 394 Ota, Portugal 566 Otway shelf 40 outcrop 61, 62, 184, 592, 593, 604, 648, 803, 816, 820, 832, 879, 882, 906 outcrop analog study 884, 878; Plate 149 on p. 887 outcrop modeling 885
Index
outer ramp 29, 596 outer shelf 40, 427, 436, 526 Ovalveolina 472 overburden 272, 277, 301, 313, 314, 647 overgrowth 74, 78, 268, 295, 800 overlap 544, 546 overload 320, 779 overpacking 312 overpressure 230 overprint 361, 744, 866, 878, 890 oversaturation 292, 313, 914 oversteepening 49, 218, 224, 226, 228, 230, 756 overwash sediment 177 ovoids 650 Oxfordian (Jurassic) 366, 395, 432, 534, 698, 700, 714, 785, 871 oxic 612, 613 oxidation 200, 205, 271, 304, 379, 623, 662, 798, 858 aerobic 842 anaerobic 859, 860 oxidation-reduction balance 54, 897 oxide 379, 497, 757 oxygen 18, 117, 186, 189, 384, 450, 455, 464, 537, 587, 612, 613, 619, 636, 637, 662, 800, 856, 873 concentration 455 consumption 620 oxygen isotope palaeothermometry 616, 654 oxygen isotopes 616, 654, 912 oxygen level 455, 610, 796 oxygen maximum (SOM) 612 oxygen-minimum zone 8, 455, 497, 613 oxygen-restricted biofacies (ORB) 613, 615 oxygenation 189, 544, 587, 636, 656, 798 oysters 42, 207, 493, 524, 666, 827, 828, 856
P Pachycanaliculata barrandei 510 Pachyphloia/Langella 460 Pachystylostroma 918 Pacific 44, 51, 88,478, 486, 515, 594, 613, 623, 800, 859 packing 61, 152, 177, 178, 181, 311, 312, 314, 361, 480, 523, 554, 587, 610, 714, 772 packing density 800, 831 packing index 152, 312, 779 packstone 25, 38, 56, 89, 117, 188, 214, 230, 236, 312, 313, 314, 320, 348, 349, 350, 352, 358, 362, 364, 366, 416, 427, 432, 438, 444, 459, 464, 465, 480, 482, 484, 486, 523, 524, 526, 530, 538, 548, 551, 570, 579, 589, 595, 596, 597, 627, 630, 632, 634, 637, 638, 640, 659, 662, 663, 666, 668, 670, 672, 675, 676, 682, 686, 694, 696, 698, 700, 702, 704, 706, 710, 714, 716, 718, 719, 723, 737, 739, 740, 746, 747, 748, 750, 752, 757, 759, 760, 762, 764, 766, 777, 781, 782, 784, 786, 788, 789, 790, 795, 796, 798, 806, 807, 808, 809, 810, 812, 813, 822, 826, 827, 828, 834, 835, 837, 842, 843, 844, 848, 852, 856, 858, 884, 896, 899, 905, 908, 919, 921; Plate 44 on p. 353 diagenetic 362, 465
Index
fenestral 704 Padurea Craiului, Apuseni Mountains, Romania 468, 470 Pakistan 462 Palaeoaplysina 560, 562, 564, 566 palaeoberesellids 750 Palaeodasycladus 434, 440, 444 Palaeodasycladus mediterraneus 440 Palaeodictyoconus arabicus 468 Palaeonubecularia 382 Palaeosiphonium 562 Palaeosiphonocladales 436, 750 palaeotextulariid foraminifera 582, 584 Palaeothrix 562 paleo-escarpment, reconstruction 849 paleo-fire 169 paleo-oxygenation 612, 613, 614, 615 paleo-water depth 136, 142, 200, 382, 384, 390, 396, 480, 591, 634, 635, 636, 637, 638, 653, 671, 818, 837, 874, 881, 888; Plate 108 on p. 639 paleobathymetry 51, 91, 416, 460, 515, 635, 638, 658, 822 paleocaliche 13, 726, 727, 728, 822; Plate 128 on p. 729 paleoceanography 25, 38, 156, 406, 456, 483, 653 Paleocene (Tertiary) 170, 308, 352, 417, 418, 442, 444, 447, 478, 556, 692, 727, 728, 782 paleoclimate 2, 16, 43, 44, 101, 156, 185, 406, 450, 483, 524, 580, 587, 597, 604, 612, 615, 625, 626, 628, 632, 654, 734, 853, 866 paleocommunity 589, 624, 756 paleocurrent 61, 169, 428, 540, 593, 596, 842, 856 paleoenvironment 7, 8, 10, 54, 58, 59, 60, 71, 80, 99, 107, 117, 175, 242, 350, 361, 379, 400, 406, 416, 465, 466, 470, 478, 480, 491, 520, 542, 587, 652, 656, 657, 875, 890 Paleogene (Tertiary) 393, 394, 435, 452, 474, 802 paleogeography 604, 847, 869 paleohydrology 737 paleokarst 2, 14, 56, 158, 206, 217, 218, 220, 279, 663, 710, 728, 730, 732, 757, 818, 892; Plate 129 on p. 733 criteria 731 significance 734 paleolatitude 4, 43, 44, 101, 450, 480, 566, 625, 626, 628, 630, 632 paleoproductivity 654 Paleoproterozoic 868 PaleoReef database 870, 871 paleosalinity 450, 452, 455, 618, 619, 653 paleosol, paleosoil 10, 12, 13, 18, 21, 37, 55, 56, 167, 205, 564, 725, 728, 730, 734, 742, 806, 808, 818, 819, 827; Plate 128 on p. 729 paleostress 223, 226, 336 Paleotethys 436 paleowind direction 14, 597, 604 Paleozoic 7, 43, 49, 55, 84, 95, 98, 188, 139, 196, 200, 211, 270, 274, I, 301, 322, 372, 373, 376, 378, 388, 390, 392, 393, 401, 406, 408, 410, 412, 416, 418, 420, 422, 426, 428, 430, 435, 444, 450, 452, 453, 458, 459, 460, 480, 490, 506, 507, 515, 510, 519, 520, 522, 533, 534, 538, 540, 546, 547, 550, 552, 554, 558, 564, 570, 584, 588, 591, 613, 616, 626, 628, 630, 636, 654, 664, 665, 669, 692, 712, 716, 719, 727, 728, 764, 768, 780, 793, 794, 795,
965
803, 808, 838, 848, 853, 869, 871, 872, 874, 875, 878 Paleozoic coral groups Plate 83 on p. 511 Palermo, Sicily, Italy 781 palisade calcite 622 Palorbitolina lenticularis 468 paludal 354 palustrine carbonates 10, 11, 13, 18, 23, 364, 728, 737, 742, 782; Plate 48 on p. 365 palynofacies 101, 656 palynomorphs 101, 820, 859 Pamplona, Spain 208 Pamukkale near Denizli, Turkey 18 Panormidella 562 Panthalassa 524, 560, 678 Papua New Guinea 677 parabreccia 640 Parachaetetes 420, 692 Parachaetetes asvapatii 418 Parachaetetes johnsoni 418 paraconformity 204 paraconglomerate 240 Paraeolisaccus 559, 562 Paradox Basin, U.S.A. 879 Parafusulina 400 paragenetic sequence 290 paralic margin 656 parallel bedding, lamination 184, 595 Parallelopora 500 parameter diagram 779 parasequence 805, 807, 808, 816, 817, 818, 821, 824, 826 Parastriatopora 510 Parathurammina dagamarae 458 parautochthonous 465, 482, 551, 552, 698, 712, 717 Paris Basin, France 3, 128, 708, 879 particle flux 92, 189 particulate organic matter, POM 49 Pass Lueg, Salzburg, Austria 812 Passega diagram 246, 247, 251, 252, 253 patch reef 121, 414, 576, 603, 663, 666, 676, 696, 717, 719, 762, 766, 770, 777, 782, 918 pavement mosaic 74 pavimentum punicum 906, 908 Pb 67, 857 Pb-Zn deposit 892, 893 PDB standard 654 pea-stone 736 peat 663 pebble 248, 719, 773, 775, 803, 847, 906 Pedata/Pötschen Beds, Triassic 789, 790, 792 pedogenesis 12, 205, 364 pedogenic carbonates 2, 10, 11, 12, 13, 14, 21, 75, 169, 205, 206, 218, 220, 354, 368, 564, 661, 663, 364, 710, 725, 726, 735, 742, 746, 749, 757, 793, 806, 809, 818 pedogenic overprint 364, 810, 818 pedotubule 727, 728 peels 1, 64, 65, 247 pelagic bivalves 218, 221 pelagic carbonates 4, 8, 47, 48, 50, 221, 238, 267, 407, 450, 456, 483, 537, 538, 553, 653, 769, 774, 775, 789, 793, 794, 795, 800, 803, 805, 807, 814, 866, 872; Plate 139 on p. 798
966
pelagic ooze 46, 311, 314, 771 pelagic platform 658, 794 pelagic ridge 201 pelagic sedimentation 54, 205, 220, 756, 776, 868 pelagite 802 Peleponnesus, Greece 140 pellet mud 37, 111 pelletoid 113, 114, *116 pellets 95, 110, *111, 116, 188, 356, 650 Pelmatozoa 550 pelmicrite 357, 368, 530, 662, 692, 700, 706, 720, 759, 760, 810 peloid 4, 6, 12, 16, 18, 20, 36, 37, 38, 40, 41, 56, 71, 84, 97, *100, 102, *110, 111, 112, 113, 114, 116, 117, 118, 146, 163, 164, 168, 174, 178, 182, 190, 212, 222, 248, 251, 303, 304, 312, 314, 316, 321, 325, 349, 350, 356, 360, 362, 368, 371, 372, 374, 382, 395, 450, 452, 544, 547,578, 579, 598, 626, 628, 648, 649, 650, 663, 666, 668, 680, 682, 686, 690, 694, 700, 702, 704, 710, 712, 713, 714, 716, 717, 718, 719, 723, 726, 727738, 746, 749, 759, 760, 762, 766, 770, 774, 777, 782, 784, 787, 790, 810, 812, 820, 821, 824, 842, 843, 844, 848, 852,858, 888. 922, 924; Plate 10 on p. 115 algal 113, 117, 118, 875 bahamite 111, 114, 116, 119, 875 benthic 114, 116 bioerosional 113, 875 freshwater 117 glauconitized 208 lithic 113 microbial 114, 116, 316, 371, 504, 619, 700, 875 mud 114, 116, 117, 717, 875 precipitated 114, 117, 118, 875 significance of 117 types 113 peloidal 84, 96, 118, 128, 344, 366, 371, 735 peloidal fabric 184, 195 peloidal question 111 Peloponnesus, Greece 190, 316, 822, 824 pelsparite 84, 692, 700, 720, 759, 762, 850 Peneckiella 374 Penicillus 104 Pennsylvanian, Carboniferous 95, 108, 284, 358, 379, 420, 428, 480, 529, 630, 661, 668, 700, 728, 803, 808, 831, 871 Pentacrinus 551 pentamerid brachiopods 519, 520 pericontinental sea 26 periostracum 519, 522, 529 periplatform 48, 50, 88, 99, 407, 659, 662, 769, 771, 780, 788, 789 perithallus 414, 417, 420, 422, 444 peritidal 24, 25, 58, 192, 354, 428, 530, 582, 666, 704, 719, 745, 746, 749, 750, 752, 788, 806, 818, 822, 826; Plate 132 on p. 751 permeability 3, 12, 14, 18, 61, 104, 152, 189, 197, 206, 208, 225, 243, 277, 278, 282, 283, 284, 288, 292, 298, 302, 304, 306, 313, 314, 320, 340, 430, 877, 878, 879, 880, 881, 882, 883, 885, 886, 888, 889, 890, 895, 896, 897, 898, 900 Permian 63, 88, 118, 134, 135, 158, 166, 181, 222, 236,
Index
296, 303, 306, 314, 322, 354, 379, 382, 384, 389, 390, 392, 394, 396, 399, 400, 412, 417, 418, 422, 423, 424, 426, 427, 428, 432, 436, 438, 442, 444, 448, 452, 454, 458, 459, 460, 462, 464, 465, 480, 496, 498, 506, 510, 516, 519, 520, 547, 550, 551, 553, 556, 558, 560, 562, 563, 564, 566, 568, 584, 597, 628, 630, 640, 690, 694, 702, 710, 728, 737, 780, 781, 803, 814, 831, 836, 844, 852, 869, 871, 872, 878 Permian Basin 843, 879 Permian Reef Complex 157, 161, 162, 330, 460, 506, 560, 836, 843, 846 Permian-Triassic boundary 298, 462 Permian-Triassic extinction 346, 379, 389, 520 Permocalculus 424, 426, 427 Persian Gulf 3, 12, 25, 28, 84, 92, 111, 119, 146, 158, 160, 162, 163, 207, 249, 330, 618, 647, 657, 879 petrogenetogram 310 petrographic fossil diversity (PFD) 260 Petrographic Image Analyis 282 petroleum seep 857 petroleum source rock 67 petrophysical data 267, 885, 888, 890 Peyssonelia 417, 444, 620 peyssoneliacean red algae 138, 264, 296, 381, 412, 413, 417, 420, 427, 620 pH 18, 97, 292, 379, 405, 504, 864 phaeodarian radiolarians 482 Phaeophila engleri 390 Phaeophyta 404 Phanerozoic 3, 4, 6, 14, 16, 43, 44, 50, 67, 121, 267, 328, 346, 369, 377, 378, 379, 392, 398, 401, 412, 424, 491, 529, 540, 548, 550, 597, 636, 647, 648, 653, 655, 657, 660, 664, 665, 669, 739, 770, 831, 832, 836, 857, 866, 868, 870, 871, 872, 874, 878 Phanerozoic reefs 378 pharetronid sponges 492 phenolic compounds 916 phi value 243 Philippeville, Belgium 833 Philippines 142, 886 Phillipsastraea 833 Pholadidae, bivalves 389 phoronids 208, 389, 392 phosclast 650 phosooid 650 phosphate 102, 205, 208, 379, 533, 648, 649, 650, 712, 757 phosphatic 71, 208, 402, 558, 560, 624, 648, 650, 800 phosphatization, phosphogenesis 211, 212, 214, 304, 624, 649, 650; Plate 110 on p. 651 phosphorite 18, 53, 214, 624, 637, 641, 854 granular 648 orthochemical/microbial 650 reworking 648 phosphorus 588, 623, 747 photic 379, 396, 635, 637, 638, 754, 842, 843, 873, 874 photic zone 8, 26, 40, 84, 196, 436, 456, 483, 487, 611, 626, 636, 638, 647, 656, 798, 833, 852, 917
Index
photoautotroph 394, 611 photoluminescence 66 photomosaic 821 photosymbionts 456, 483, 612, 626 photosynthesis 15, 19, 104, 125, 184, 302, 369, 404, 405, 458, 512, 610, 611, 623 phototrophic 8, 16, 84, 127, 369, 385, 612 phototrophic zone 210 photozoan grain association 626, 628 phreatic 13, 16, 67, 205, 217, 276, 300, 726 Phylactolaemata, bryozoans 515 phylloid algae 104, 178, 342, 382, 405, 406, 417, 420, 422, 423, 427, 428, 566, 737; Plate 58 on p. 429 phylloid algal mound 297, 428, 686 phylogenetic 492, 459 physical property 896 phytoclast 49, 101 phytoherms 18 phytokarst 386 phytoplankton 20, 26, 483, 610, 613, 624 Pieninia 562 Pierre Chatel, Grenoble, France 512 Pietra di Salomone, Sicily, Italy 460, 780, 781 Pietraroia, Matese Mts., Italy 819 pinnacle reef 878, 881 Pine Point, ore deposit, Canada 893 piso-oncoid 123 pisoids 4, 13, 15, *100, 111, 121, 144, 146, *157, 158, 160, 161, 162, 190, 206, 248, 579, 590, 619, 650, 698, 710, 717, 726, 728, 735, 736, 748, 749, 806, 813, 819, 888; Plate 14 on p. 159 caliche 157, 158, 160, 161, 162, 710, 726, 742, 818 cave 160 depositional settings 162 fluvial 160, 161 formation 20, 160, 161 hypersaline 161 lacustrine 161 marine-vadose 158 non-carbonate 162, 710 pedogenic 710, 726 phreatic vadose 710 primary marine 161 significance 162 siliceous 158, 749 vadose 158, 160, 161 vadose-marine 158, 159 pisolite 21, 121, 161, 162, 681,698, 710, 842 pit connection 413 Pithonella 450, 452, 474 pits and grooves 207 placoderms 560; Plate 97 on p. 561 Placopsilina 212 Placunopsis 60 Placunopsis ostracina 524 planar lamination, stratification 57 Planiinvolutina carinata 678 plankton 6, 42, 43, 47, 49, 50, 80, 88, 91, 92, 96, 111, 493, 538, 540, 544, 559, 616, 624, 635, 662, 666, 667, 716, 763, 769, 774, 779, 798, 854, 872, 875 planktonic algae 49, 78, 91, 406, 450, 452, 790 planktonic foraminifera 41, 91, 102, 616 Planorbulina 444
967
plant debris 739 plants 14, 15, 16, 102, 169, 380, 381, 382, 390, 447, 593, 605, 607, 623, 656, 735, 739, 748, 827, 828, 911 Plassen Formation, Jurassic 252, 768, 786 plasticity 832, 912 plasticlasts *167 plateau 662, 756, 794 platestone 342, 356 platform 3, 4, 6, 7, 25, 28, 31, 37, 41, 75, 136, 205, 212, 216, 220, 221, 222, 257, 264, 276, 301, 313, 330, 332, 354, 369, 377, 378, 379, 380, 384, 396, 401, 402, 404, 418, 422, 423, 424, 432, 436, 442, 446, 447, 452, 455, 458, 459, 460, 464, 465, 466, 468, 470, 472, 476, 478, 480, 481, 500, 502, 507, 515, 522, 523, 524, 526, 528, 533, 534, 552, 554, 562, 563, 564, 566, 569, 570, 582, 584, 588, 594, 595, 604, 614, 623, 624, 625, 628, 630, 642, 643, 644, 646, 653, 655, 656, 657, 658, 659, 660, 661, 662, 663, 664, 669, 671, 672, 673, 676, 677, 678, 680, 681, 688, 694, 696, 698, 700, 704, 706, 712, 713, 714, 716, 728, 734, 748, 750, 752, 753, 754, 755, 757, 759, 766, 770, 772, 774, 780, 781, 782, 786, 788, 789, 790, 792, 794, 803, 804, 805, 806, 810, 814, 815, 816, 817, 818, 819, 821, 822, 824, 827, 834, 847, 848, 856, 862, 864, 869, 870, 874, 875, 878, 879, 880, 882, 886, 892; Plate 134 on p. 763 attached 28, 663, 698, 700, 768, 850 Bahamian-type 714, 760, 762, 849 drowned, drowning 27, 379, 428,497, 624, 658, 662, 756, 757, 760, 796, 800, 818, 824, 885 edge 671, 681, 694, 696 emergence 756, 771 evaporitic 663 geometry 658 growth and demise 660, 760, 783 intrabasinal 786 isolated, non-attached 27, 37, 660, 667, 668, 698, 700, 704, 760, 794, 815, 863 lost platforms 803, 847, 850 non-rimmed 660, 661, 681, 753, 770 open platform 570, 663, 671, 694, 696, 762, 768, 834 pelagic 756, 794 Phanerozoic 869 Proterozoic 875 restricted 570, 663, 696, 702, 716, 834 rimmed platform 356, 657, 660, 661, 662, 664, 665, 667, 671, 681, 700, 716, 718, 730, 753, 770, 774, 820, 821 submergence 756, 794, 795 platform and ramp, case study 759 platform carbonates 3, 63, 78, 135, 142, 301, 334, 384, 428, 447, 459, 754, 803, 868, 870, 875, 896 platform distribution 870 platform interior 412, 588, 663, 664, 669, 678, 688, 694, 698, 700, 702, 704, 706, 708, 712, 713, 757, 762, 779, 784, 789, 790, 810, 813, 820, 822, 824, 827, 848, 877, 892 platform margin 29, 37, 216, 217, 220, 223, 332, 481, 512, 526, 528, 560, 566, 570, 594, 604, 660,
968
661, 662, 663, 667, 673, 688, 698, 753, 756, 757, 759, 762, 768, 769, 770, 774, 781, 784, 785, 786, 788, 789, 820, 834, 835, 846, 848, 878, 879, 892 platform sand 780 platform sequence 570 platform size 870 platform top 28, 660, 757, 759, 820 platform-basin transect 769, 672, 787, 788, 790; Plate 138 on p. 791 playa 11, 20, 193, 647 Pleistocene 12, 14, 16, 18, 36, 38, 40, 94, 95, 102, 107, 222, 276, 290, 423, 602, 725, 728, 804, 821, 906 pleurocapsalean cyanobacteria 620 Plexoramea 562, 688 Pliensbachian (Jurassic) 230, 524, 737 Pliocene (Tertiary) 444 Plitvica, Croatia 18 Plöcken Pass, Austria/Italy 732 Plumley et al. classification 591, 593 pockmark structure 857 Podobursa 484 Pohlkotte, Lower Saxony, Germany 414 poikiloaerobic 613 poikilotopic 301 point counting 99, 247, 254, 255, 256, 257, 258, 261, 263, 266, 282, 394, 395, 602, 627, 628, 677, 694, 752, 779, 789, 847 point grid 255 Point-, line-, ribbon- and area-counting 912 point-count group 790 Poland 190, 314, 541, 546, 704, 752, 776, 892 polar 3, 19, 31, 42, 496, 509, 615, 623, 626 polar ice 805 pollen and spore 67, 101 pollution 125, 900 polychaete worms, annelids 111, 208, 389, 392, 510, 540 polycystines, radiolarians 482 Polygnathus ex gr. P. glabra 560 Polygonella 562 polymerization 899 polymorphic transformation 923 polymuds 842 polyphyletic 496, 498, 506 polyplacophorid chitons 111 polysaccharide 404 Polystrata 417 POM (Polymeric Organic Matter) 49 Pomar Model 753, 754, 755 pond 442, 663, 668, 698 Pontides, Northern Turkey 526 porcelaneous calcareous microstructure 454 porcelanite 483 Pordoi Pass, Dolomites, Italy 782 pore 3, 279, 282, 288, 292, 302, 340, 352, 371, 544, 589, 610, 882, 886, 889, 910 pore fluid 13, 91, 224, 276, 292, 320, 643, 800 pore geometry 65, 278, 282, 292, 340, 881, 890, 896, 898 pore neck 900 pore pressure 313, 314 pore size distribution 881, 889, 895, 898
Index
pore space 68, 272, 282, 292, 297, 301, 325, 334, 588, 643, 714, 889, 890 pore system 837, 881, 882, 885 pore throat 278, 282, 881, 890, 895, 896, 899 pore type 282, 283, 288, 878, 881, 884, 889, 890 pore water 67, 70, 186, 200, 228, 272, 276, 277, 288, 289, 295, 301, 313, 332, 484, 640, 641, 643 Porferitubus 562 Porifera 491, 506 Porites, poritid coral 140 601 poroperm value 661 porosity 1, 2, 3, 12, 16, 18, 61, 68, 95, 102, 104, 150, 163, 182, 189, 197, 206, 208, 225, 243, 251, 254, 258, 260, 261, 267, 272, 276, 277, 278, 279, 282, 283, 288, 289, 301, 310, 311, 312, 313, 314, 316, 334, 340, 390, 428, 430, 438, 582, 630, 673, 674, 686, 734, 766, 774, 779, 818, 854, 866, 873, 877, 878, 879, 881, 882, 883, 884, 885, 888, 890, 892, 895, 896, 897, 898, 899, 900, 904 918, 923; Plate 30 on p. 286; Plate 149 on p. 887 Archie classification 283 biomoldic 886, 889 boring 280 breccia 280, 883 burial 882 burrow 280 cavern 206, 279, 280, 283, 802 channel 280, 283 Choquette and Pray classification 281, 282, 283 dissolution 888 effective 278, 881, 889, 898 enhanced 279 eogenetic 279 fabric-selective 280, 281, 283, 284 fenestral 280, 283 fracture 280, 283, 288, 289, 880, 883, 886 framework growth 279, 280, 283, 288. 289, 304, 832, 837, 880 in thin section 280 intercrystalline 74, 278, 280, 283, 284, 287, 288, 888, 898, 900 intergranular 24, 37, 38, 73, 102, 113, 279, 324, 344, 352, 607, 714, 747, 879, 880, 882 interparticle 68, 152, 192, 236, 279,280, 282, 283, 284, 288, 289, 290, 292, 304, 308, 312, 322, 325, 350, 354, 584, 698, 700, 824, 881, 898, 923 interskeletal 268, 438, 673 intracrystalline 271, 326, 333 intraparticle 255, 280, 283, 284, 322, 886, 898, 923 intraskeletal 284, 528, 837 keystone vug 284 measurement 282, 881 mesogenetic 279 minus-cement 278, 283, 312 moldic 68, 102, 280, 283, 284, 304, 496, 528, 879, 880, 886, 888 non-fabric selective 279, 281, 284 oomoldic 152, 284, 889 pre-cementation 279 primary 279, 284, 288, 292, 866, 880, 881, 882, 888 secondary 279, 284, 288, 386, 389, 816, 881 shelter 178, 193, 280, 283, 673, 714, 835
Index
shrinkage 280 solution 279, 288, 289 telogenetic 279 terminology 283 total 278 vuggy 280, 283, 288, 304, 880 porosity classification 881 porosity development 272, 282, 308, 310, 366, 912 porosity groups 283 porosity plot classification 283 porosity reduction 301, 302, 311, 800, 885 porosity types 152, 253, 282, 881, 886; Plate 29 on p. 285 porosity-depth curve 313 porosity-permeability interrelationship 288, 295 Porostromata 97, 122, 125, 128, 373, 408, 766 porostromate algae 408, 530, 688, 704, 706, 777 porostromate calcimicrobes 130, 385, 410, 412 porostromate cyanobacteria 140, 252, 373, 374, 384, 385, 438, 500, 704, 706, 714; Plate 53 on p. 411 porostromate oncoid 21, 122 porostromate skeletal oncoid 611 porostromate taxa 408 Portugal 168, 566 Posidonia 88, 220, 528, 537, 748, 790 post-brecciation fracture 232 post-drowning phase 757, 759, 760 post-hurricane situation 601 pot casts 598 potassium ferricyanide 271 Potomogeton 739 pottery 903, 908, 910, 911 power spectral analysis 805 Powow Canyon, Hueco Mountains, Texas, U.S.A. 516 Poza Rica field, Mexico 877 Praealveolina 472 Praecalpionellites dadayi 488 Praecalpionellites filipescui 488 Praecalpionellites hillebrandti 488 Praecalpionellites murgeanui 488 Praecalpionellites siriniaensis 488 Praeglobotruncana 474 Praetintinnopsella andrusovi 488 Pragian (Devonian) 842 Prague Basin, Czech Republic 538 pre-drowning phase 757, 759, 760 pre-hurricane situation 601 Precambrian 14, 16, 103, 135, 143, 144, 153, 163, 166, 170, 182, 192, 296, 328, 370, 373, 377, 378, 379, 392, 404,412, 606, 620. 710, 728, 734, 806, 814, 836, 868, 874, 875, 878, 880 Precambrian-Cambrian boundary 378, 404, 409, 410, 412, 876 precession 804, 805 precipitate 562, 566 abiotic 873 biotically controlled 47, 873 microbial 84, 868 precipitation 15, 16, 18, 19, 25, 47, 80, 81, 84, 95, 117, 118, 132, 160, 200, 206, 208, 270, 271, 276, 290, 292, 295, 299, 302, 304, 321, 332, 352, 364, 369, 370, 371, 379, 404, 412, 436, 529, 620, 652, 653, 661, 710, 713, 735, 771, 788, 836, 838, 863, 868, 873, 883, 893
969
precipitation of aragonite 276 precipitation of calcite 270 precipitation rate 88, 272, 295 precompaction 313 Precordillera, Argentina 378, 644 precursor aragonite 653 precursor mineralogy 78, 274 predation/grazing pressure 175, 348, 384, 393, 416 presence/absence data 263, 581, 724 preservation 18, 106, 233, 310, 374, 402, 418, 428, 484, 514, 530, 534, 538, 546, 548, 563, 580, 610, 614, 619, 624, 649, 680, 717, 779 preservation of fossils 18, 59, 104, 819 preservation potential 49, 71, 105, 408, 493 pressure 277, 301, 313, 316, 322, 334, 504, 779 pressure solution 55, 56, 60, 91, 93, 96, 194, 198, 200, 236, 239, 268, 271, 272, 277, 278, 301, 310, 311, 313, 314, 316, 318, 320, 321, 354, 362, 644, 774, 778, 796, 818, 865, 889, 890 classification 318 grain-to-grain 316 seam 794 significance 321 terminology 319 Prethocoprolithus 541, 542 primary dolomite 356 primary mineralogy 71, 405 of fossils 653 primary pores 288 primary production 47, 613, 623, 624 Prince Patrick Island, Canadian Arctic 860 principal component analysis 805, 807 prism cracks 746, 809 pristane/phytan ratio 614 Proaulopora 504 probitumen 374 problematic algae 413, 422 procaryotes, Procaryota 82, 370, 404, 408, 410, 830, 838 Procladocora tenuis 308 product moment correlation 262 productid brachiopods 519, 520 production rate 478, 480, 486 productivity 49, 91, 587, 613, 623, 625, 654, 656, 660, 769, 790, 794, 808, 816, 849, 866 profundal zone 19 progradation 7, 660, 781, 782, 789, 804, 807 Propeller Mound, Northern Atlantic, Ireland 388 Prosobranchia, gastropods 529 proteins 117 Proterozoic 55, 104, 298, 362, 369, 377, 390, 412, 484, 492, 579, 613, 731, 739, 875 protists, Protista 102, 474, 482, 484 protointraclasts 167, 640 protolith 336 protomylonite 322 protopeloids 118 Protopeneroplis 481 Protopeneroplis striata 466 protozoans 369, 453, 487, 491 provenance analysis 236, 847, 903, 908 provenance area 229, 232, 242 Providence Channel 88 proximal tempestites 595 proximality curve 775, 776
970
Index
proximality-distality trends 594, 847 Prut Valley, Moldavia, Romania 530, 542 pseudo-algae 420 pseudo-bedding 56 pseudo-birdseyes 192 pseudo-breccia 232, 239, 748 pseudo-cracking 239 pseudo-fenestrae 193 pseudo-imbrication 232 pseudo-microkarst 13, 364 pseudo-ooid 116, 143 Pseudoactinoporella fragilis 442 Pseudochaetetes asvapatii 352 Pseudocyclammina 454, 481, 766 Pseudocyclammina lituus 466 Pseudocymopolia jurassica 442 Pseudodonezella 436, 848 Pseudofusulina 236 pseudointraclast *167 Pseudolabechia 500 Pseudolithothamnium 417 Pseudolithothamnium album 352, 444 pseudomatrix 73 pseudomicrite *75, 81, 92, 96 pseudomorphs 214, 299, 333, 497, 622, 625, 648 pseudoncoids 130 pseudonodular 200 Pseudopalaeoporella lummatonensis 424 pseudopellets 113, 116 pseudopeloids 114, 116 pseudoplanktonic 528, 790 pseudopunctate brachiopods 519, 520 Pseudoschwagerina 236 pseudospar, pseudosparite 97, 322, 324, 325 pseudosparitization 192 Pseudovermiporella 562, 564 Pseudovermiporella sodalica 564 pteropod ooze 49, 91, 922 pteropods 41, 49, 91, 106, 440, 529, 532, 538, 793, Ptychochaetetes 502 Ptychochaetetes globosus 502 Puerto Rico 92, 863 Pulmonata, gastropods 529 pumping 117, 178, 304, 332, 588, 806, 837, 893 punctate brachiopod 68, 520 Purbeck, Cretaceous 155, 314, 316, 362, 450, 544, 766 purity of limestone 896 pycnocline 8 Pycnoporidium lobatum 410 Pyramid Lake 20 Pyrenees Mountains, Spain 446 pyrite 54, 170, 210, 316, 333, 484, 533, 548, 559, 644, 646, 857, 858, 860, 921; Plate 109 on p. 645 authigenic 646 framboidal 644, 647, 858 reworked 647 pyritization 460, 484, 614, 644, 646, 647 pyroclastic 292, 802, 864
Q Q-dike 223 Q-fissure 217, 218, 220
Q-mode cluster analysis 262 quadrat-belt method 260 quartz 257, 277, 292, 304, 308, 362, 366, 468, 472, 478, 483, 548, 554, 560, 586, 622, 641, 642, 643, 644, 663, 682, 686, 692, 698, 706, 712, 725, 728, 734, 737, 738, 739, 744, 757, 760, 764, 800, 806, 810, 822, 824, 826, 848, 850, 860, 863, 865, 899, 911, 918, 919, 921 authigenic 622, 644, 732 detrital 591, 592, 642 euhedral 643 idiomorphic 643 megaquartz 734 terrigenous 264, 646 quartz pebble 240, 366 quartz sand 243, 541, 819 quasi-abiotic precipitate 873 Quasifusulina 322 Quaternary 61, 388, 390, 392, 465, 616, 654, 727, 728, 734, 789, 821 Queensland, Australia 138, 600, 660, 821 quick lime 896 quiet-water environment 56, 136, 137, 211, 339, 340, 352, 376, 493, 504, 551, 552, 591, 592, 606, 704, 764 Quinqueloculina 453, 713
R R-mode cluster analysis 262, 263, 264, 266 radiolaria/spicule ratio 484 radiolarian mud 486 radiolarians 4, 49, 91, 186, 218, 221, 230, 312, 336, 358, 402, 453, 481, 482, 483, 484, 486, 490, 534, 570, 580, 618, 624, 643, 647, 662, 682, 684, 780, 782, 794, 795, 796, 797, 800, 801, 849, 908; Plate 76 on p. 485 biozonations 486 classification 482 diagenesis 483, 486 preservation 483, 484 significance 486 radiolarite 483, 484, 486, 797, 801 Radiolitidae, bivalves 608, 526, 920 Radiomura 559, 562 Radiomura cautica 566 Radiosphera 450 raggioni 299 Raibl Formation, Triassic 137, 788 rain 14, 604, 725, 816, 864 ramp 3, 25, 26, 28, 29, 63, 76, 181, 196, 257, 264, 297, 298, 301, 313, 354, 358, 379, 384, 385, 386, 392, 412, 422, 424, 427, 436, 447, 452, 459, 460, 465, 468, 478, 480, 481, 504, 507, 509, 512, 515, 522, 528, 538, 542, 544, 552, 554, 569, 570, 577, 582, 588, 594, 595, 597, 598, 625, 628, 630, 634, 636, 637, 642, 643, 644, 646, 655, 657, 658, 659, 660, 661, 662, 664, 665, 666, 671, 672, 675, 676, 681, 696, 700, 713, 716, 717, 718, 719, 723, 734, 753, 754, 755, 764, 766, 768, 795, 805, 806, 815, 817, 818, 822, 824, 826, 827, 833, 834, 836, 843, 848, 856, 862, 877, 878, 879, 880, 886, 892; Plate 135 on p. 765; Plate 136 on p. 767
Index
back-ramp 666, 718, 719 classification 665 deeper-water ramp 672, 716, 718, 847 distally steepened 27, 29, 37, 578, 580, 665, 666, 716, 717, 719, 722, 753, 754, 755 drowned 794 epeiric 658, 660, 664, 669 homoclinal 27, 29, 37, 638, 665, 672, 675, 698, 722, 753, 768, 770 inner ramp 121, 254, 354, 406, 666, 671, 686, 694, 696, 698, 702, 704, 713, 714, 716, 717, 719, 753, 764, 766, 826 mid-ramp 664, 665, 666, 671, 673, 682, 692, 694, 698, 713, 714, 716, 717, 719, 764, 826 outer ramp 46, 665, 666, 672, 682, 698, 716, 717, 718, 719, 764, 766, 877 subdivision 29 Ramp Microfacies Types (RMF) 717, 718, 719, 722, 764, 766 ramp model 660, 664, 666, 681, 764, 848 ramp types 665 ramp-to-basin 665 ranking 260, 360, 580, 584, 837 Rare Earth elements 67,652 rare element 67, 267, 612, 614, 619, 652, 735, 896 Ras al Khaimah, United Arabian Emirates 158, 278, 424, 468, 470, 472, 542, 643, 644, 646, 647, 710 rasping 4, 390, 831 rate of sedimentation 292, 623, 646 Rauhwacken 238 ray crystals 303, 735 reactate calcite 316 reaction group 672 receptaculitids 422 recovery 428, 600, 881 recrystallization 14, 55, 71, 74, 75, 76, 94, 95, 96, 104, 116, 120, 122, 144, 172, 194, 224, 267, 268, 271, 290, 321, 322, 324, 336, 362, 374, 382, 390, 402, 418, 428, 484, 550, 562, 582, 888, 896, 904; Plate 38 on p. 323 red algae 41, 97, 124, 126, 140, 290, 296, 304, 358, 384, 385, 390, 394, 396, 405, 406, 412, 414, 422, 427, 444, 476, 560, 568, 579, 604, 620, 635, 638, 666,668, 717, 719, 754, 755, 766, 873, 886 articulated 104 melobesiid 819 problematic 420 red algal nodule 140 red algal spores 452 red bed 647, 663 Red Sea 20, 80, 84, 107, 111, 146, 207, 249, 276, 618, 647, 881 redeposition 166, 169, 170, 184, 198, 200, 229, 248, 464, 482, 580, 594, 640, 641, 680, 708, 710, 779, 784, 822 redox potential 277, 369, 624, 644 reducing condition 169, 271, 646, 662 reef 6, 7, 20, 29, 31, 36, 37, 55, 63, 74, 75, 78, 99, 175, 205, 221, 222, 289, 292, 296, 297, 302, 308, 313, 325, 330, 332, 342, 343, 362, 369, 371, 372,373, 374, 377, 378, 379, 380, 382, 384, 385, 389, 390, 392, 393, 394, 402, 406, 410, 414, 416, 417, 418, 422, 424, 426, 428,
971
458, 460, 464, 465, 466, 478, 480, 481, 491, 493, 496, 498, 500, 502, 504, 506, 507, 509, 510, 512, 515, 520, 528, 541, 542, 551, 559, 562, 563, 564, 566, 569, 591, 593, 594, 600, 602, 604, 623, 624, 625, 630, 635, 644, 653, 655, 658, 661, 662, 663, 672, 673, 674, 676, 677, 678, 680, 681, 686, 688, 696, 706, 712, 716, 723, 734, 738, 753, 755, 756, 757, 762, 764, 768, 769, 770, 772, 774, 779, 786, 790, 798, 803, 807, 816, 820, 822, 830, 831, 835, 837, 843, 847, 848, 852, 856, 862, 864, 868, 870, 871, 872, 877, 878, 879, 880, 881, 885, 886, 892, 896 abundance 870, 871, 872 barrier 662, 664, 844 basic constituents 834 bivalve 42, 493, 526, 719 carbonate production 29, 870, 871 case study 836 cement 807, 836 cold-water reef 3, 29, 41, 42, 852, 856 coral 37, 42, 392, 394, 395, 512, 719, 766, 871 coralline algal reef 42 coralline sponge-microbe 872 deep-water 42, 769, 922 definition 830 destruction 390, 831 distribution 830 drowned 533, 796 ecologic 830, 844, 870, 871 framework 541, 588, 600, 830 geometry 593, 864 hydrodynamic 830 knoll 662 Lophelia reefs 42 microbial reef 190, 379, 393, 395, 393, 872 organic 673 patch 659, 665, 675 pinnacle 665, 666, 880, 885 platform-margin 135, 236, 661, 642, 662, 781, 870, 877, 882, 885, 892 Proterozoic 868 rudist 864 seismic 831, 872 serpulid reef 541 shelf-edge 379, 602, 603 siliceous sponge-microbial 871 sponge 42, 392, 466, 542 staked reefs 869 stromatoporoid reef 392, 500, 552, 871 thickness 871, 872 tropical 36, 522 true 830, 871 type 830 upper-slope 661, 662 reef belt 121, 264, 528 reef biota, reef builders 59, 218, 289, 341, 346, 348, 366, 379, 402, 416, 494, 496, 509, 512, 559, 566, 576, 590, 593, 600, 601, 605, 604, 624, 625, 664, 665, 678, 680, 690, 723, 757, 760, 790, 807, 830, 831, 832, 833, 835, 836, 837, 842, 843, 844, 870. 871, 874, 908 reef blocks 218, 602, 605 reef carbonate 218, 340, 342, 354, 418, 423, 575, 628,
972
816, 830 classification 833 practical guide 836 reef cavity 93, 117, 180, 296, 299, 384, 496, 506, 559, 566, 678, 790, 832, 837 reef clast 168, 236 reef complex 243, 250, 253, 500, 582, 604, 653, 664, 676, 696, 700, 779 reef crisis 871 reef debris 343, 354, 438, 605, 676, 807, 830, 834, 837, 870, 872, 920 reef environment 89, 303, 382, 579 reef fabric 343 reef flank 551, 552 reef flat 500, 600, 702 reef framework 346, 379, 382, 386, 389, 416, 498, 528, 623, 831, 832, 844 reef guild 807, 831, 832 reef limestone 4, 88, 223, 341, 343, 344, 354, 364, 384, 402, 409, 418, 576, 603, 690, 833, 834, 837, 843, 904, 908 reef margin 249, 418, 466, 668, 770, 834 reef mound 389, 422, 428, 436, 496, 504, 564, 638, 717, 771, 782, 830, 831, 837, 870, 871, 872, 878, 880, 892 reef pattern 870 reef recovery 600 reef slope 137, 222, 289, 354, 382, 600, 880 reef structure 374, 378, 844 reef thickness 872 reef type 395, 831, 834, 870, 871 glossary 30 reef zone 674, 780, 798 refractive property 896 regression 62, 576. 624, 641, 642, 753, 808, 882 Regularia, echinoids 554 Reichelina 462 Reichenhall, Northern Calcareous Alps, Germany 710 Reisskofel, Carinthia, Austria 157 relay 723, 724, 843 relief 65, 198, 597, 647, 816, 866 Remaniella cadischiana 488 Remaniella catalanoi 488 Remaniella colomi 488 Remaniella duranddelga i488 Remaniella ferasini 488 Remaniella sp. 487 Renalcis 372, 399, 409, 410, 412, 504, 507 Reophax giganteus 470 replacement 95, 271,274, 308, 321, 330, 362, 648, 649, 650, 726 replacement dolomite 238, 328, 332 replacement of silicates 644 reprecipitation 120, 272, 364, 753, 775, 866, 867 reptiles 558 Reptonoditrypa cautica 516 resedimentation 47, 169,170, 507, 627, 634, 769, 800, 802 reservoir 74, 78, 128, 181, 189, 279, 282, 288, 615, 659, 756, 766, 803, 877, 878, 879, 884 depositional setting 878 distribution 880 fracture-controlled 883, 884 parameter 659
Index
porosity 881 reservoir extent 877 reservoir geometry 879 reservoir heterogeneity 878, 880, 885 reservoir model 582, 877 reservoir porosity 206, 881 reservoir potential 1, 43, 188, 197, 224, 528, 718, 879 reservoir prediction 885 reservoir pressure 881 reservoir rock 1, 3, 67, 73, 99, 102, 136, 197, 206, 225, 278, 339, 340, 354, 414, 453, 528, 566, 604, 686, 718, 734, 817, 877, 879, 881, 882, 888 reservoir rock quality 91, 320, 604, 878, 881, 885, 886 reservoir rock types (RRTs) 885, 890 reservoir storage capacity 225 residual clay 169, 644, 647 residual soil 55 resin 3, 387, 388, 881, 886, 889 Resita, Apuseni Mountains, Romania 442 468 resource, limestone 99, 657, 670, 700 Retiophyllia 364, 690, 810 Retiophyllia clathrata 577 return flow 594, 595 reverse grading 54, 96, 178, 718, 775 reworking 188, 194, 210, 229, 478, 482, 523, 597, 608, 637, 650, 680, 700, 710, 784, 801, 809 Rhaetian (Triassic) 4, 364, 393, 432, 464, 497, 502, 559, 564, 566, 577, 677, 678, 688, 690, 698, 700, 708, 798, 908 Rhapydionina liburnica 470 rhizoconcretion 12, 14, 727, 742 rhizogenic calcrete 727 rhizoid, rhizolith 423,726, 727, 742 rhizome 89, 493 rhizoturbation 206 rhodalgal grain association 138, 819 rhodocline 611 rhodoechfor grain association 627, 670 rhodoid 36, 121, *122, 123, 134, *137, 138, 139, 140, 142, 248, 406, 414, 416, 420, 579, 589, 612, 636, 717, 766, 806, 875, 888, 889; Plate 12 on p. 141 abundance 139 branched 138 composite 124, 139 deeper-water 138 describing 139 distribution 139 economic importance 142 growth pattern 138, 139 shape 138 significance of 140 temperate 138 tropical 138 rhodoid beds 139 rhodoid clast 137 rhodoid limestone 137, 142 rhodolith 44, 122, 123, 135, 264 rhodolith pavement 44, 138, 416 rhodophycean algae 140, 386, 404, 412, 420, 458, 828, 853 rhynchonellid brachiopods 519, 520, 522 rhythmic bedding, succession 185, 866
Index
rhythmite 893 Ribadresella, Spain 544 ribbon counting 247, 254, 255, 256, 261 Richthofenia 520 Ricla, Zaragoza, Spain 208 Ries, Germany 361, 448, 735, 740 Rif Mountains, Morocco 420, 850 rifting 218, 220, 801, 858, 889 rigidity 805, 896 Rigidocaulis 568 rim 333, 334, 660, 661, 665 circumgranular 270 isopachous 270 rip-up clasts 167, 482 ripple bedding, ripple, ripple lamination 56, 595, 596,597, 667, 746, 773 ripple mark 56, 58, 177, 593, 598, 647, 746 river 15, 172, 245, 246, 447, 450, 643, 772, 820, 880, 906 Rivularia 409, 410 RMF see Ramp Microfacies Types rock color 53, 54, 61, 62, 64, 97, 205, 357, 613, 808, 847, 884, 895, 896, 897, 899, 905, 913 rock resource 895 rock type 356, 657, 884, 888 Rock-Color Chart 54 rock/petrophysical classes 283, 889 rockfall 48, 49, 688, 771, 772, 781, 784 rockground 189, 204, 606 Rogenstein, Early Triassic 150, 374 Roman mosaic stones, Kraiburg, Germany 644 Roman mosaic stones, Hemmaberg, Austria 326 Roman quarries, Chemtou, Tunisia 322 Roman quarry, Tentschach, Carinthia, Austria 322 Romania 442, 468, 470, 530, 542 root mold 185, 190, 206 root rock 748 root structure, classification 727 root tubule 364 root void 749 rootlets 170, 190 roots 12, 14, 37, 137, 185, 190, 364, 381, 478, 726, 726, 727, 747, 748, 762, 818 Rosengarten Mts., Dolomites, Italy 782 Rosskofel, Carnic Alps, Austria 551 Rosso Ammonitico, Jurassic 198, 218 Rotaliacea, Rotaliina 455, 456, 458, 459, 465, 472, 476, 478, 616, 618, 677 Rotliegend, Permian 742 Rötelstein, Styria, Austria 350 Rötelwand, Salzburg, Austria 676 Rothpletzella 134, 140, 382, 409, 410, 500 roughness 897 rounding 174, 357, 587, 589, 674, 688, 717, 772, 834 roundness 173, 174, 175, 176, 232, 304, 590, 591, 592, 673, 686, 864, 911 roundness classes 174 roundness/sphericity 174, 230, 236 roveocrinoids 553, 556 rudist bivalves 346, 526, 528, 576, 608, 666, 686, 717, 772, 814, 826, 827, 828, 834, 864, 865, 869, 871, 888, 920; Plate 88 on p. 526 rudist limestone 528, 576 rudist reef 350, 385, 632
973
Rudnigsattel, Carnic Alps, Austria 462 rudstone 214, 244, 341, 349, 354, 356, 364, 366, 414, 523, 542, 589, 596, 597, 605, 608, 663, 670, 675, 686, 688, 690, 696, 700, 708, 710, 712, 723, 747, 757, 759, 760, 762, 766, 781, 786, 788, 807, 813, 824, 834, 848, 850, 853, 858, 886, 898, 919, 920; Plate 45 on p. 355 Rugendorf, Frankenwald, Germany 560 rugose corals 509, 510, 512, 833, 848, 919; Plate 83 on p. 511 rupturing 221, 889 Russia 532
S S- and Q-fissures 217, 220, 222, 223 Saar-Nahe, Germany 737, 742 Sabaudia minuta 470 sabellarid polychaetes 42, 856 sabkha 25, 330, 622, 644, 647, 648, 663, 716, 745, 748, 749, 766, 843 Saccamminopsis 848, 638 Saccocoma 553, 554, 556, 563, 776, 796 Sacramento Mountains, New Mexico 675, 836 saddle dolomite 278, 301, 326, 332 Sagalassos, Turkey 903 Saiwan, Central Oman 528 Sakmarian (Permian) 428, 560, 564 Salda Golu Lake, Turkey 21 salebrids 675 Salento Peninsula, Puglia, Italy 472, 702 Salina Janubio, Lanzarote, Canary Island 152 salinity 18, 19, 20, 21, 42, 71, 80, 117, 125, 136, 147, 210, 217, 292, 301, 302, 360, 364, 376, 405, 406, 416, 423, 432, 438, 442, 450, 455, 456, 466, 515, 520, 541, 544, 569, 579, 587, 607, 612, 616, 617, 618, 619, 620, 622, 626, 632, 634, 636, 643, 652, 653, 656, 659, 660, 662, 663, 677, 694, 698, 702, 747, 754, 756, 850, 863, 864, 881; Plate 104 on p. 621 Salpingoporella 430, 444 Salpingoporella annulata 440 Salpingoporella dinarica 442 Salpingoporella ex gr. muehlbergi 442 salt 20, 25, 49, 53, 622, 897, 899 salt flat 708 Salt Range, Pakistan 462 salt-and-pepper texture 168, 169 saltation 247, 251 sample size 62, 575 sampling 53, 62, 63, 64, 653, 778, 847, 885, 895 error 255 procedure 575 San Pietro, Apennines, Italy 542 San Salvador Island, Bahamas 38, 164, 284 sand, calcareous 6, 7, 37, 58, 146, 248, 249, 250, 267, 295, 356, 364, 390, 396, 416, 423, 424, 442, 448, 478, 528, 600, 601, 602, 604, 605, 606, 607, 610, 630, 637, 663, 665, 666, 667, 694, 696, 698, 702, 713, 714, 719, 762, 770, 774, 784, 788, 877, 879, 880, 914 apron 668 bank 666
974
beach 290 bryozoan 515, 630 coarse- and very coarse-grained 249 dolomite 333 dune 290 fine- and very fine-grained 250 lithoclastic 667 medium to fine-grained 249 pyroclastic 292 siliciclastic 248, 362, 366, 416, 541, 865 skeletal 608, 713, 747, 789, 795 windblown 647 sand bar 29, 464, 770, 848 sand bottom 438, 606 sand flat 762 sand production 37, 817 sand shoal 302, 661, 662, 663, 664, 665, 666, 820, 848, 878, 880, 885, 896 sand-size 368, 778, 786, 858 Sandberg model 272, 273, 874 Sandfang 177 sandstone 560, 598, 728, 732, 742, 826, 827, 863, 905 sandy allochem limestone 366 sandy limestone 53 Santonian (Cretaceous) 474, 526 Sarcodina 453, 482 Sardinia, Italy 206, 560, 630, 892 Sargassum 90 Satonda, Indonesia 620; Plate 104 on p. 621 saturation of CaCO3 40, 95, 272, 283, 889 saturation-height 890 Saudi Arabia 80, 877 Saulgau well, Germany 362 Sauvagesia 526 Savonnières-en-Pertois, France 900 Saykamt, Turkey 560 Scaglia Rossa, Cretaceous 684, 797 Scandinavia 211, 215 scarid parrotfish 387 Schammental, Suebian Alb, Germany 542 schizohaline 371, 530, 618 Schizophoria 68 Schizosphaerella 78, 91, 798 Schlern, Dolomites, Italy 186, 788 Schnegglisteine 125 Schrecksee, Allgäu, Germany 226 schwagerinid fusulinids 462 scleractinian corals 108, 509, 512, 514, 517, 626, 676, 831, 835, 853, 856, 922 sclerites 494, 509 sclerobiont 381 sclerosponges 346, 354, 492, 500, 502, 504, 506, 781 scolecodonts 101, 106 Scottish shelf 42 scour structure 56, 596, 746, 773 sculptural stone, sculpture 900, 907 scyphozoans 509 Scythian (Triassic) 532, 872 Scytonema 706 sea bottom, sea floor 8, 10,14, 26, 40, 47, 75, 91, 99, 169, 188, 189, 222, 223, 250, 386, 453, 474, 483, 486, 524, 541, 547, 550, 562, 594, 595, 606, 607, 646, 659, 662, 664, 665, 666, 669,
Index
712, 747, 755, 776, 816, 868, 875 sea grass 81, 88, 89, 90, 96, 170, 177, 288, 302, 346, 382, 455, 480, 541, 602, 606, 610, 626, 656, 667, 668, 669 sea ice 42, 630 sea level 24, 25, 58, 60, 203, 222, 470, 497, 630, 641, 664, 675, 728, 730, 732, 745, 753, 754, 756, 769, 804, 805, 807, 816, 862, 870 sea pens 854 sea urchins 91, 386, 550, 554 sea weeds 515, 853 sea-level curve 379, 874 sea-level cycle 816 sea-level fall 660, 760, 809, 816, 818, 817, 819, 822 sea-level fluctuation, sea-level change 3, 14, 34, 37, 50, 101, 136, 142, 156, 162, 169, 196, 211, 216, 236, 247, 264, 330, 379, 382, 428, 447, 625, 635, 654, 655, 658, 659, 660, 661, 724, 734, 745, 753, 756, 759, 760, 769, 771, 780, 784, 789, 792, 803, 804, 805, 807, 808, 809, 810, 813, 815, 816, 817, 818, 821, 822, 824, 847, 862, 863, 881, 892; Plate 141 on p. 823; Plate 142 on p. 825 sea-level highstand 344, 789, 800, 808, 821 sea-level indicator 223, 658 sea-level lowstand 759, 770, 789, 816, 820 sea-level rise 210, 211, 659, 660, 756, 809, 817, 818, 822, 824, 827, 832, 838, 864 sealing 225, 766, 884, 899 seam 316, 318 seamount 42, 44, 46, 48, 134, 200, 206, 304, 374, 522, 552, 553, 648, 756, 772, 793, 794, 796, 800, 836, 917 search sampling 63 seasonal change 15, 42, 127, 616, 740 sea urchin 917 seawater 210, 292, 617, 623, 663, 801, 874, 882 seawater chemistry 95, 404 seawater dolomitization model 331, 332 seawater temperature 520, 616 seaweed 34, 90, 916 sector zoning in calcite 67 secular variation 6, 99, 101, 154, 270, 272, 273, 817, 868, 872 sediment 42, 50, 122, 210, 247, 370, 389, 573, 635, 640, 659, 753, 754, 806, 866, 874 argillaceous 484 biological 369 internal 731 siliciclastic/pyroclastic 864 stable-bottom 89 sediment export 753, 789 sediment input 7, 99, 196, 247, 381, 492, 496 sediment production 386, 416, 624, 663, 667, 713, 753, 754, 755, 792, 804 sediment supply 804, 816, 817, 862 sediment transport 24, 28, 247, 594, 597, 769, 885 sediment-water interface 189, 206, 455, 537, 588, 624, 649 sedimentary fabric 581 sedimentary structure 53, 55, 57, 58, 61, 63, 126, 177, 337, 357, 379, 590, 591, 637 sedimentation 47, 301, 308, 340, 370, 374, 390, 658 allochthonous 50
Index
background 712 particle-by-particle 793 pelagic 50, 794, 863 reduced 131, 698 siliciclastic 864 sedimentation method 244 sedimentation rate 3, 55, 80, 101, 184, 186, 189, 201, 205, 206, 211, 223, 224, 292, 302, 304, 360, 382, 384, 386, 460, 483, 515, 523, 528, 542, 587, 590, 606, 610, 623, 630, 636, 642, 646, 794, 795, 805, 820, 837, 838, 897 sedimentation unit 56, 182 Sedir Island, Turkey 913, 914 Seebühel, Switzerland 484, 684 Seekreide 20 Seeland Alm, Cortina d’Ampezzo, Italy 418 seep and vent carbonate 3, 352, 803, 857, 858, 873 seep carbonate 859, 860 seep community 860 seep mound 858 seepage 710, 858, 859, 893 seepage-reflux model, dolomitization 330, 331 Seewen, Switzerland 474, 566 Seiser Alm, Dolomites, Italy 182, 230 seismic activity 779, 922 seismic data 661, 756, 815, 874, 878, 884 seismic event 167, 228, 587, 640 seismic interpretation 884 seismic profile 772, 820 seismic reef 831 seismic reflection 296, 815, 816 seismic shock 167, 169, 640, 772 seismic survey 661 seismites 597, 640 Seletonellaceae 435 Sella Platform, Dolomites, Italy 787, 786, 788 Selliporella donzellii 440 SEM 64, 65, 66, 71, 73, 78, 104, 121, 267, 330, 371, 386, 387, 388, 395, 396, 399, 414, 452, 483, 484, 519, 534, 554, 566, 568, 620, 635, 655, 800, 881, 888 semiarid climate 643, 708, 734, 747, 749 Senonian (Cretaceous) 470 sequence 6, 56, 58, 63, 523, 546, 644, 647, 648, 805, 815, 816, 821 depositional 815, 816, 821 fifth order 804, 817, 821 first order 805, fourth order 804, 817, 821 second order 805 third order 804 sequence analysis 215, 803, 816, 821 sequence boundary, stratigraphic boundary 2, 216, 724, 728, 815, 817, 818, 819, 820, 821, 822, 824, 826, 827, 828, 885 Type-1, Type-2, Type-3 818 sequence stratigraphic break 847 sequence stratigraphy 1, 7, 101, 177, 203, 206, 215, 216, 236, 379, 634, 658, 660, 672, 728, 730, 803, 815, 816, 818, 822, 824, 826, 878, 885, 888; Plate 141 on p. 823; Plate 142 on p. 825 sequence tract 818 serial peel 65 serpentinite 800, 801, 865
975
Serpula socialis 542 serpulid limestone 586 serpulid polychaetes 4, 34, 41, 44, 90, 97, 104, 105, 123, 128, 130, 140, 178, 214, 208, 220, 374, 381, 382, 384, 385, 404, 452, 518, 541, 542, 559, 584, 608, 620, 632, 650, 667, 717, 771, 785, 836 853, 854, 856, 860, 916, 917, 922; Plate 92 on p. 543 serpulite 541, 542 Serraduy, Pyrenees, Spain 727 Serravalian (Tertiary) 871 sessile benthic fossils 624 Sestrostomella robusta 498 setting 7, 50, 102, 478, 579, 682 basinal 217, 379, 452, 468, 486, 504, 540, 552 carbonate-siliciclastic 804 cold-water 625, 854 cool-temperate 90 cool-water 569, 635, 661, 667 depositional 189, 263, 384, 460, 480, 528, 576, 642, 652, 660, 680, 725, 803, 877 lacustrine 75, 236, 358 Sexten Dolomites, Italy 236 shale 313, 464, 595, 598, 647, 659, 662, 686, 717, 732, 764, 794, 796, 828, 840, 857, 863, 868 shaliness 884 shallow burial 275, 277, 297, 628 shallow-marine 3, 7, 8, 26, 28, 29, 34, 37, 41, 50, 58, 221, 243, 334, 374, 406, 424, 438, 450, 456, 491, 544, 547, 548, 551, 558, 560, 564, 566, 584, 588, 594, 634, 642, 643, 646, 660, 665, 667, 678, 680, 682, 700, 725, 728, 757, 773, 774, 790, 806, 809, 814, 816, 817, 822, 863, 865, 868 shallowing-upward 63, 128, 136, 357, 528, 576, 595, 641, 642, 648, 653, 659, 660, 724, 745, 750, 752, 766, 804, 805, 806, 807, 808, 817, 818, 819, 820, 821, 882 Shamovella 568 shape parameter 173 Shark Bay, Australia 21, 25, 102, 372, 376, 618, 664 shark teeth 560, 650 Shartymophycus 420 shear, fracture, shearing 49, 224, 226, 200, 310, 320, 322, 607, 771, 780, 786 shedding 769, 789, 790, 792, 800 sheet cavity 197 sheet crack 197, 298, 746, 747, 762, 812 sheet flow 770, 775 shelf 2, 3, 25, 31, 32, 40, 181, 298, 358, 378, 382, 414, 427, 428, 436, 440, 446, 447, 453, 455, 456, 459, 460, 464, 465, 466, 468, 478, 480, 486, 500, 504, 509, 510, 512, 514, 515, 520, 522, 523, 533, 536, 540, 541, 547, 550, 553, 564, 584, 569, 587, 593, 594, 598, 600, 602, 604, 608, 612, 613, 623, 625, 626, 628, 630, 632, 634, 635, 636, 644, 647, 648, 653, 657, 662, 664, 675, 692, 696, 698, 716, 732, 748, 771, 774, 795, 800, 806, 812, 815, 817, 818, 836, 843, 844, 848, 862, 866, 877, 878 cold-water, non-tropical 37, 667, 852 deep 515, 659, 662, 686, 678, 712, 713, 716, 822, 847 inner 667, 862, 863
976
lateral classification 28 mid- 526, 667 non-rimmed 26, 27, 658, 754 open 34, 44, 436, 667, 754, 819, 852 pericontinental 46 rimmed 26, 27, 28, 34, 660, 664, 665, 672, 680, 681, 754, 755, 818 temperate 40, 480 unrimmed 28 shelf break 10, 28, 29, 31, 62, 452, 636, 665, 667, 675, 818, 874 shelf carbonates 31, 34, 62, 406, 418, 420, 422, 424, 426, 432, 452, 458, 459, 464, 625, 626, 627, 628, 638, 647, 819, 867, 875 shelf lagoon 34, 236, 570, 848, 850, 880 shelf margin, edge 29, 40, 222, 428, 476, 523, 526, 528, 560, 600, 603, 604, 657, 659, 664, 667, 686, 770, 810, 822 shelf platform 40, 428 shelf-to-basin transect 658, 661 shell beds 59, 60, 522, 523, 534, 595, 597, 606, 900 shell classification 519 shell concentration, accumulation 211, 524, 526, 534, 597, 598, 680, 696, 712, 720, 857 shell debris, fragments 76, 304, 352, 452, 596, 608, 682, 669, 812, 822 Shell Standard Legend 174, 283, 884 shellground 189, 606 shells 40, 68, 312, 358, 395, 399, 400, 402, 519, 598, 612, 619, 644, 655, 688, 692, 766, 781, 784, 785, 798, 801, 824, 854, 888, 904, 911, 922 shellstone 342 shepherd’s crook 306, 548 sherd, ceramics 910 shoal 36, 436, 438, 458, 478, 480, 604, 659, 661, 664, 666, 694, 698, 702, 712, 714, 717, 719, 762, 766, 770, 826, 827, 842, 877, 880 shoaling upward sequence 162, 804, 820 shoaling/deepening trend 817 shore-to-basin transect 657 shoreface 428, 666, 716, 827, 828 shoreline 13, 14, 24, 28, 184, 593, 622, 636, 664, 666, 669, 698, 717, 742, 812, 818, 863 Shotori Mountains, Iran 694, 784 shrinkage 218, 728, 912, 908 shrinkage crack 56, 710 shrinkage pore, void 192, 196, 747, 809, 810 Shuguria 409 Sicily 170, 178, 215, 220, 354, 460, 465, 534, 564, 677, 780, 781 sickle cell 731, 735 siderite 214, 292 Siderolites 632 Siderolites calcitrapoides 472 Sierra de Aulet, Huesca, Spain 468 Sierra de Urbasa, Navarra, Spain 632 Sierra Nevada, Granada, Spain 702, 836 sieve data 211, 244, 247, 251, 668 Sigmoilina 678 silcrete 11, 731, 734 Siliana, Maktaris, Tunisia 474 silica 11, 12, 102, 482, 483, 484, 491, 497, 533, 559, 644, 662, 774, 889
Index
cement 643 silicates 379, 641, 644, 653 siliceous 65, 402 siliceous ooze 643 siliceous sponges 492, 494, 496, 504, 566, 608, 625, 831, 838, 871; Plate 78 on p. 495 siliciclastic 19, 58, 220, 362, 366, 385, 551, 641, 642, 647, 648, 658, 659, 713, 732, 756, 769, 771, 773, 788, 803, 815, 816, 820, 846, 857, 858, 860, 862, 863, 864, 866, 906, 913, 914 siliciclastic influx 135, 385, 386, 769, 626, 660, 827, 863, 864, 896 siliciclastic sediment 19, 613, 644, 656, 816, 817, 863, 864 silicification, silification 106, 190, 268, 271, 304, 320, 546, 560, 643, 644, 774; Plate 109 on p. 645 silt /siltite 54, 74, 95, 108, 167, 218, 220, 248, 252, 297, 299, 303, 344, 352, 362, 541, 606, 642, 714, 739, 778, 764, 774, 806, 819, 824, 830 silt size 95, 778, 838 silt/clay ratio 642 siltstone 732, 827, 828, 857 Silurian 90, 94, 95, 134, 154, 378, 384, 394, 396, 412, 416, 418, 422, 448, 450, 459, 490, 496, 500, 510, 522, 530, 532, 534, 538, 544, 546, 548, 552, 563, 566, 686, 780, 814, 831, 838, 842, 872, 878, 880 similarity coefficient, matrix 658, 724 simulation model 62, 658, 866 Sinai, Egypt 68, 128, 147, 824, 826, 827, 828 Sinemurian (Jurassic) 440, 524, 551, 682, 737, 798 sinkhole 13, 14, 730 sinter 11, 13, 15, 16, 67, 364, 731, 734, 735, 740 SiO2-Al2O3-CaO plot 904 siphonocladalean green algae 620 Siphonodendron 510 sipunculids 90, 389, 392 Sirmione, Lake Garda, Italy 797 size distribution, 248, 249 skeletal and non-skeletal grains 42 skeletal composition 652 skeletal debris 97, 458, 578, 682 skeletal grain 3, 37, 38, 40, 74, 101, 104, 146, 163, 166, 176, 189, 248, 251, 268, 270, 290, 354, 399, 406, 452, 491, 570, 579, 580, 610, 622, 628, 630, 649, 663, 665, 666, 682, 694, 700, 706, 708, 712, 713, 717, 718, 719, 723, 738, 753, 754, 759, 770, 774, 777, 779, 781, 782, 784, 795, 820, 821, 827, 842, 850, 852, 868, 874, 875, 888, 912; Plate 9 on p. 109 skeletal grain association, assemblage 41, 616, 853; Plate 105 on p. 629; Plate 106 on p. 631; Plate 107 on p. 633 skeletal grain category 582 skeletal mineralogy 101, 103, 104, 402, 420, 502, 515, 579, 580, 616, 874 skeletal stromatolites 122, 373, 374, 408 skeletonization 260, 831, 832 skewness 245, 250, 252, 779 slacking, slack 896, 898, 911 slake durability 898 Slaughter Canyon, Guadalupe Mts. 843 sliding 49, 58, 64, 229, 771, 779, 781, 790, 792, 858, 893 sliding and slumping 780
Index
slip plane 49, 336 slope 10, 26, 29, 48, 50, 93, 99, 169, 181, 186, 200, 217, 222, 236, 297, 354, 374, 426, 455, 464, 466, 472, 478, 480, 481, 486, 496, 520, 522, 523, 526, 551, 552, 554, 560, 566, 569, 578, 582, 594, 600, 601, 602, 603, 613, 644, 647, 655, 658, 659, 660, 661, 662, 664, 671, 674, 686, 688, 690, 696, 700, 755, 759, 760, 769, 770, 772, 773, 774, 775, 779, 780, 781, 782, 784, 785, 786, 788, 792, 793, 794, 795, 798, 801, 802, 815, 817, 818, 821, 822, 834, 837, 843, 844, 848, 849, 852, 874, 878, 880, 885, 892 accretionary 769 angle 667, 669, 770, 786 break 28, 29, 665, 666 bypass 50, 769, 785 depositional deposit 50, 218, 296, 402, 661, 757, 769, 770, 780, 781, 782, 784, 785, 786, erosional 769 gradient 664 gullied 769 low-angle 784 model 769 stability, stabilization 224, 786,787, 788 type and composition 769 slope and basin, diagnostic criteria 771 slope geometry 769 Slovenia 476, 807, 808 Slowakia 444 slump 49, 58, 226, 374, 418, 593, 597, 660, 665, 702, 704, 771, 779, 800, 834 slumping 47, 49, 58, 196, 197, 229, 230, 771, 779, 780, 790, 893 Smackover Formation, Jurassic 877, 879, 880 small-scale cross bedding, lamination 57, 184 smaller foraminiferal assemblage 459, 460 smectite 277, 279, 644 SMF (Standard Microfacies Type) 680 definition 712 smokers 51 SMOW standard 654 snowball earth 362, 868 snuff-boxes 131 sodium 619, 653 soft algae 382, 604, 624 soft bottom, sediment 188, 197, 272, 344, 494, 504, 523, 554, 603, 605, 606, 610 soft chalk 211 soft corals 42, 44 softground 186, 189, 212, 602, 606 soil 10, 12, 14, 55, 84, 169, 185, 369, 625, 641, 656, 725, 726, 727, 809, 817 soilstone crust 169 Solar Lake, Israel 20 sole mark 595, 596, 597, 773 solemyid bivalves 859 Solenolmia manon 498 Solenomeris 480, 886 Solenomya 857 Solenopora 420 Solenopora cassiana 418 solenoporacean red algae 236, 352, 382, 410, 412, 413, 418, 436, 438, 678, 692, 908, 910; Plate 55 on
977
p. 419 Solnhofen, Bavaria, Germany 78, 358 Solnhofen limestone, Jurassic 54, 75, 92, 93, 358 Solnhofen-Oberalm model 78, 93 solubility 200, 318, 455, 648 solution 4, 175, 184, 198, 207, 208, 217, 220, 229, 276, 282, 283, 289, 292, 303, 320, 321, 334, 402, 405, 465, 623, 643, 653, 775, 880, 883 breccia 232 cavity 206, 290, 296, 690, 809, 810, 837 film 106, 270, 298, 316, 529 pore 149, 278, 288, 308 seam 230, 314, 316, 320 void 4, 59, 108, 192, 299, 306, 644, 886 solution-evaporite-collapse breccia 232, 648 solution/precipitation 55, 149, 333 sonic log 884 Sonnwend Mountains, Tyrol, Austria 223, 678 soot particle 900 Sorby Principle 7 sorting 14, 37, 74, 92, 175, 176, 178, 181, 198, 235, 239, 244, 245, 246, 248, 250, 357, 366, 480, 523, 552, 554, 578, 584, 587, 589, 590, 591, 592, 596, 607, 610, 673, 674, 684, 698, 717, 725, 731, 762, 771, 772, 773, 774, 775, 835, 895 sorting classification 252 Sosio, western Sicily 170, 178, 354, 564 soupground 189, 606 source area 233, 600, 642, 661, 847 source mixing model 863 South African shelf 138 South China Sea 142, 886 South Florida Shelf 660 Southern Franconian Alb, Germany 326, 566 spaeloid 158 Spain 214, 366, 432, 436, 446, 468, 470, 544, 570, 630, 634, 796, 848, 849 spar 59, 371, 714, 842 sparite 16, 73, 324, 325, 352, 357, 358, 364, 366, 368, 427, 578, 694, 708, 726, 740, 918 sparmicritization 94, 97 sparstone 361, 362 Spatangoidea, echinoids 556 Spearman’s Correlation Coefficient 264 species diversity 444 species richness 604, 619 specific surface 18, 896, 897 spectral analysis 67, 792, 805, 807, 813, 817 speleothem 11, 13, 14, 158, 160, 167, 299, 361, 364, 656, 730, 731, 732, 734 Sphaeroschwagerina carniolica 462 sphericity 173, 174, 175, 176, 304, 590, 686, 688 sphericity-form diagram 127 spheroids 143 spherulitic 105, 118 sphinctozoan sponges, sphinctozoic sponges 4, 236, 346, 382, 385, 392, 492, 494, 496, 498, 497, 502, 504, 506, 563, 678, 835; Plate 79 on p. 498 spicules 483, 486, 494, 496, 497, 504, 506, 507, 580, 644, 782, 854, 908 spiculite 496, 682, 853 Spinaporella andalusica 702 spines 520, 548, 580, 616
978
spiriferid brachiopod 519, 520 Spirillilina 466 spirorbid worms 108, 124, 530, 542 Spitsbergen 42, 44,738 splash zone 745 split core samples 188 sponge borings 113, 302, 389, 393, 394 sponge buildup, mound 42, 47, 493, 494, 496, 497, 569, 690 sponge spicules 4, 42, 76, 184, 195, 336, 484, 486, 494, 496, 593, 643, 644, 662, 667, 675, 682, 717, 719, 759, 760, 782, 785, 796, 797, 798, 800, 849, 900 sponges 4, 34, 41, 42, 44, 82, 180, 182, 195, 268, 381, 382, 384, 385, 386, 389, 394, 395, 396, 404, 422, 491, 492, 494, 500, 507, 508, 520, 559, 563, 564, 566, 568, 569, 591, 600, 601, 604, 606, 607, 618, 625, 638, 666, 667, 668, 692, 704, 717, 754, 771, 781, 785, 798, 835, 838, 844, 908 calcareous 4, 492, 498 chaetetid sponges 384, 497, 502, 506, 516 clionid sponges 90, 386 coralline sponges 105, 402, 438, 497, 502, 504, 506, 560, 717, 910 definitions 492 diagenesis 497 freshwater 492, 618 glass sponges 494 inozoan sponges 497, 498, 502, 504, 506, 843 siliceous sponges 117, 492, 493, 494, 504, 624, 625, 717, 757, 853, 854 soft sponges 76, 496, 853 stromatoporoid sponges 402, 502 subdivision 492 Spongiomorpha ramosa 502 spongiomorphid sponges 236, 502, 576, 910; Plate 81 on p. 503 Spongiostromata 122, 373 spongiostromate 124, 125, 128, 135, 373, 374 spongiostromate oncoid 122 spongiostrome 371, 704 spores 101, 436 Sporolithon 414, 444 spring 14, 15, 16, 734 hot, thermal, mineral 735, 736 spring mound 735 spring tuffs 408 spumellarian radiolarians 482, 483, 782, 797 spur and groove 593 squamariacean red algae 352, 381, 406, 417, 428, 444, 620, 635 Sr 66, 274, 321, 330, 617, 619, 652, 653, 735, 857 Sr/Ca ratio 270, 617, 619, 654 Sr/Mg ratio 274, 619 SS structure, stromatolites 125 SS-C structures, stromatolites 125, 127 SS-I structures, stromatolites 127 SS-R oncoids 127 St. Margarethen, Burgenland, Austria 381, 414 St. Paul Monastery, Wadi Araba, eastern Egyptian D 230 St.Barbe Quarry, Dinant Basin, Belgium 510 stability 268, 272, 590, 607, 788, 858 stability of substrate 381
Index
stability ranking 268 stabilization 194, 271, 330, 374, 382, 542, 769, 830 stable isotope analysis 2, 4, 68, 70, 71, 80, 205, 267, 290, 306, 321, 330, 616, 619, 641, 652, 654, 655, 725, 735, 737, 742, 789, 903, 906, 912, 913 Stacheia 422, 638 stacheinacean algae 412, 413, 420, 422, 436, 438 Stacheoides 420, 422 Stachyodes 362, 500, 673 stacking sequences 25, 659, 817 staining 57, 65, 66, 271, 282, 401, 519, 735, 900 stalactite, stalagmite 13, 206, 731 standard calpionellid biozonation 487 standard deviation 246, 250, 252, 256, 779 Standard Facies Model 627, 657, 660, 875 Standard Facies Zone (FZ) 470, 478, 524, 526, 528, 657, 661, 662, 663, 672, 822, 824 standard lithologic logging system 61 Standard Mean Ocean Water 654 Standard Microfacies 551 concept discrimination, definition 680, 681 discussion 712 key, guide 720, 729 problems 714; Plate 127 on p. 715 Standard Microfacies Types (SMF) 121, 228, 263, 540, 551, 575, 582, 627, 657, 661, 664, 712, 716, 718, 723, 813, 822, 824, 848, 874 SMF 1 to SMF 26: 76, 114, 140, 150, 158, 164, 186,190, 306, 344, 346, 350, 352, 354, 358, 374, 424, 428, 438, 450, 474, 476, 484, 504, 520, 526, 530, 538, 548, 554, 628, 670, 673, 682, 681, 682, 684, 686, 688, 690, 692, 694, 696, 698, 700, 702, 704, 706, 708, 710, 712, 713, 714, 716, 719, 720, 721, 723, 725, 759, 760, 762, 764, 766, 790, 813, 819, 822, 824, 844; Plate 112 on p. 683; Plate 113 on p. 604; Plate 114 on p. 687; Plate 115 on p. 689; Plate 116 on p. 691; Plate 117 on p. 693; Plate 118 on p. 695; Plate 119 on p. 697; Plate 120 on p. 699; Plate 121 on p. 701; Plate 122 on p. 703; Plate 123 on p. 705; Plate 124 on P. 707; Plate 125 on p. 709; Plate 126 on p. 711 Stanley-Hardie model 273, 874 statistical analysis 62, 229, 243, 245, 247, 256, 464, 593, 652, 805, 807 statistical sampling 62, 229 Steinernes Meer, Salzburg, Austria 808, 810, 812 steinkern 108, 113, 212, 647, 712 Steinplatte, Tyrol, Austria 190, 298, 299, 502, 577, 581, 598, 678, 696, 700 stenohaline 618, 662, 850 Stenolaemata, bryozoans 515 Sterzing marble 322, 913 Stevns Klint, Denmark 852, 856 Stomiosphaera 452 Stomiosphaera sphaerica 566 stone conservation, decay 897, 899 storm 8, 76, 180, 184, 186, 197, 230, 249, 250, 354, 428, 460, 465, 480, 523, 544, 588, 589, 593, 594, 595, 596, 597, 598, 600, 604-606, 637, 646, 659, 663, 665, 666, 702, 708, 712-714, 716, 718, 723, 738, 739, 745, 748, 750, 762, 766, 804, 842, 863, 879 storm deposit, reworking 186,193, 523, 594, 595,
Index
596, 597, 598, 604, 666, 667, 714, 718, 778 storm layer 59, 60, 178, 594, 596, 597, 598, 834 storm wave 178, 229, 460, 523, 595, 596, 662, 696 storm wave base, storm wave weather base 8, 26, 46, 196, 551, 588, 594, 595, 635, 636, 637, 659, 664, 665, 666, 667, 694, 712, 717, 759, 833, 848 storm wave shell concentration 178 strain 224, 320, 336, 358 Straits of Florida 50, 856 strata-bound mineralization 892 stratal unit 815 stratification 37, 55, 57, 58, 61, 229, 232, 239, 648, 746, 772, 806, 807 stratified cementation 185 stratiform mineralization 892 stratigraphic boundary 653, 672 stratigraphic break 130, 136, 203, 734, 827 stratigraphic condensation 211, 648 stratigraphic trap 880 stratigraphic zonation 406, 556 Straza, Bled, Slovenia 226, 303, 325, 382, 516 strength 226, 896, 897 stress, compressive, tensional 224, 314, 316, 320, 385, 587, 619, 748, 816, 856, 857, 896 stromatactis 6, 96, 178, *190, *192, 193, 194, 195, 196, 197, 598, 636, 637, 838, 842; Plate 20 on p. 191 microbial model 195 sponge model 195 stromatactis mound 195, 196 stromatactoid fabric 704 stromatoid 374 stromatolites 16, 20, 21, 80, 82, 118, 122, 126, 131, 161, 207, 212, 276, 314, 342, 344, 364, 369, 370, 371, 372, 373, 374, 375, 376, 377, 378, 379, 396, 408, 496, 578, 618, 635, 644, 704, 738, 739, 740, 742, 748, 749, 750, 752, 806, 810, 830, 868, 869, 892; Plate 50 on p. 375 agglutinated 372, 374,376 classification 377 decline 378 definition 374 freshwater tufa 21, 372, 376 lacustrine 21, 376, 739 Logan et al. classification 376 micritic 374 saltlake stromatolites 21 skeletal 122, 372, 373, 376 Stromatolithi 122 stromatolitic 19, 83, 130,132, 378, 560, 796 Stromatoporella 362 stromatoporoid sponges, stromatoporoids 74, 104, 140, 289, 346, 356, 362, 372, 374, 382, 384, 386, 392, 395, 492, 500, 504, 506, 508, 510, 552, 590, 644, 673, 675, 692, 717, 750, 752, 764, 768, 842, 871, 918, 919; Plate 80 on p. 500 labechiid stromatoporoids 500 microstructure 500, 506 strontium 278, 420, 617, 652, 653, 655 strontium isotopes 654, 903 strophomenid brachiopods 519, 520 structure grumeleuse 371 Sturiella oblonga 488 Stylaster 496
979
stylasterid hydrozoans 508, 853, 854 Styliolina 537, 538 styliolinids 184, 334, 537, 538, 539, 540, 598, 794, 795, 796; Plate 91 on p. 529 stylobedding 185, 318, 778 stylobreccia 230, 232, 233, 239, 316, 318, 778 styloconglomerate 239 stylocumulate 316, 318, 362 stylolaminated 185, 316, 318, 362 stylolite 225, 278, 288, 311, 312, 313, 314, 316, 318, 332, 362, 644, 647, 696, 708, 889, 890, 898, 904, 919; Plate 37 on 317 stylolite model, formation 311, 316, 318, 320 stylolite seam 55, 56, 226, 362 stylolite set 185 stylolite swarms 362 stylolitization 198, 205, 271, 277, 314, 316, 318, 774, 896 stylomottled 318 stylonodular 318 Stylophyllum 578, 835 Stylophyllum polycanthum 577 styloreactate 316 Stylosmilia corallina 512 subaerial exposure 2, 11, 14, 37, 70, 95, 149, 157, 168, 169, 192, 203, 205, 206, 297, 306, 310, 334, 360, 364, 428, 514, 647, 655, 660, 663, 688, 706, 710, 728, 731, 734, 742, 744, 745, 746, 749, 757, 759, 760, 762, 806, 808, 809, 813, 818, 820, 822, 877, 881, 882, 888, 892 subarctic 3, 44 subdivision of shelves and ramps 28, 29 subglacial crusts 14 subhedral 74, 96 sublittoral 9 submarine dissolution 170, 200 submarine fillings 221 submarine fissures 553569 submarine volcanic swells 196 suboxic 613 subphotic 843 subpolar 630 subsidence 470, 658, 659, 660, 804, 807, 864 subsolution 204 subsolution clasts 170, 208 substrate 4, 56, 99, 117, 210, 295, 304, 348, 360, 381, 382, 389, 393, 395, 405, 438, 442, 450, 455, 465, 492, 500, 506, 515, 544, 569, 570, 587, 592, 602, 603, 605, 606, 607, 608, 610, 636, 671, 677, 753, 848, 850; Plate 103 on p. 609 firm substrate 576 hard 667; Plate 103 on p. 609 soft substrate 576; Plate 103 on p. 609 substrate stability 515, 607, 608, 610, 800 substrate type 205, 208, 450, 606, 607, 610 subsurface 314, 358, 384, 410, 418, 442, 448, 468, 470, 500, 516, 524, 530, 532, 541, 554, 628, 644, 646, 647, 680, 728, 768, 834 subsurface prediction 874 Subterranophyllum 414 subtidal 9, 13, 25, 46, 75, 184, 186, 201, 330, 344, 354, 369, 370, 377, 428, 444, 476, 524, 526, 541, 552, 596, 606, 611, 632, 635, 636, 659,
980
Index
669, 692, 704, 706, 708, 710, 745, 746, 747, 748, 749, 750, 752, 766, 806, 809, 810, 812, 813, 824, 827, 863, 875, 878 subtropical 31, 36, 616, 632, 661, 816, 863, 869 subunconformity 877 succession 382, 464, 625, 804, 806 succession pattern 62, 384, 807 Sulawesi, Indonesia 886 sulfate 63, 333, 484, 641, 646, 647, 648, 857 sulfate mineralization 64, 289, 648, 858 sulfate reduction 83, 118, 332, 369, 484, 620, 646, 656 sulfides 198, 326, 369, 379, 623, 641, 858 sulfur 51, 588 sulfur bacteria 370, 504, 647 sulfur isotopes 612 Sulu Sea, Philippines 414 Sulzfluh, Graubünden, Switzerland 409, 410, 440, 462, 466, 468, 713, 714, 762, 768 Sulzfluh limestone, Jurassic 176, 236, 252, 253, 760; Plate 134 on p. 763 Sumbawa, Indonesia 620 Sunda Sea 669 supersaturation 80, 118, 295, 302, 620 supersequence 869, 870 supralittoral 9 supratidal 9, 25, 192, 193, 197, 288, 344, 408, 564, 611, 644, 659, 663, 675, 704, 708, 745, 746, 747, 748, 749, 750, 752, 766, 804, 806, 809, 812, 818, 824, 850 surf 174, 588, 589, 600, 748, 750 surface karst 13 suspension 184, 247, 251, 257, 299, 356, 667, 771, 773, 864, 919 suspension feeders 493, 514, 523, 606, 607, 610, 623, 624, 757, 759, 760, 832 Svalbord 42, 138 Sverdrup Basin, Canadia 42, 138, 436, 630, 860 swallow tail 648 swamp 364, 646, 668, 742 swell 538, 665, 676, 734, 756, 794, 796, 798, 827 Sweden 186, 212, 516, 548, 795 Switzerland 306, 448, 450, 474, 484, 534, 570, 576, 682, 696, 706, 766, 802, 821 Sycidium 448, 450 symbiosis, symbionts 456, 458, 459, 512, 564, 612, 624, 636, 857, 858, 873, 920 syndepositional 194, 297, 892 syngenetic 892 syntaxial cement 300, 333 syntaxial overgrowth 3, 225, 298, 304, 306, 550, 554, 556, 918 Syringopora 764 Syringostromatida, stromatoporids 500 systems tract 646, 658, 672, 724815, 816, 817, 820, 822, 824, 826, 827, 828 T Tabas,Iran 784 tabulate corals 512, 918 tabulozoans 502, 506, 507 Tafilalt, Anti-Atlas, Morocco 534 Takkanawa Fall, Yoho Valley, Canada 316, 334 Tanganyika 84 tangential grain contact 311
Tanzania 863 taphofacies 807 taphonomic processes 59,195, 211, 261, 464, 544, 551, 579, 753, 831, 875 taphonomic loss 104, 853 taphonomy 104, 712, 717, 875; Plate 9 on p. 109 Tarvis breccia, Permian 236, 241 Tasmania 40, 516, 630 Taurus Mountain Range, Turkey 424, 432, 436, 564 taxonomic composition 59, 406, 478, 605, 768 taxonomic differentiation 418, 476, 547, 552, 568, 578 taxonomic diversity 260, 384, 405, 673 taxonomic loss 870 tebagite 560 technological property 99, 226, 897 tectonic activity 220, 223, 224, 230, 242, 805 tectonic breccia 238 tectonic deformation 172, 224 tectonic fault breccia 230 tectonics 200, 218, 222, 740, 756, 769, 770, 771, 786, 803, 804, 816, 849, 884, 892 telescoping 538, 540 telogenetic 279 temper grain 910, 911 temperate 3, 4, 31, 40, 308, 414, 442, 452, 455, 520, 542, 544, 546, 623, 626, 628, 630, 854 temperature, water temperature 7, 14, 16, 18, 31, 40, 42, 67, 71, 80, 138, 277, 289, 292, 301, 302, 308, 313, 322, 332, 333, 334, 336, 405, 406, 408, 442, 455, 456, 497, 504, 512, 515, 540, 587, 593, 612, 615, 616, 617, 623, 626, 632, 634, 635, 636, 637, 648, 652, 654, 656, 659, 660, 663, 735, 754, 756, 816, 874, 881, 910, 914 tempestite 65, 169, 170, 551, 552, 554, 594, 595, 596, 597, 598, 604, 666, 696, 719, 750, 764, 777, 778, 821; Plate 102 on p. 598 Tentaculites 184, 538 tentaculitids 105, 181, 537, 538, 540, 593, 684, 794, 796, 868; Plate 91 on p. 529 tepee 58, 161, 299, 710, 735, 746, 747, 749, 810, 818, 819, 848 Terebella 385, 541, 566, 569, 785 Terebella lapilloides 542, 690 terebratulid brachiopods 519, 520, 522, 860 terra rossa 14, 55, 205, 730, 746 terrane 849, 864 terrestrial 2, 6, 11, 16, 56, 242, 361, 366, 386, 591, 624, 641,642, 698, 710, 820, 862, 873 terrestrial carbonate 352, 820 terrestrial infill 218, 220; Plate 24 on p. 219 terrigenous input, influx 10, 34, 43, 71, 97, 137, 166, 263, 476, 630, 663, 642, 700, 740, 817, 862, 864 Tertiary 3, 7, 13, 49, 131, 135, 137, 154, 208, 296, 364, 370, 382, 384, 392, 396, 410, 412, 414, 417, 418, 422, 423, 424, 426, 442, 448, 450, 452, 457, 458, 459, 474, 478, 480, 491, 508, 515, 530, 532, 541, 547, 558, 562, 563, 570, 597, 616, 626, 628, 630, 644, 664, 669, 672, 718, 727, 728, 734, 771, 806, 814, 852, 854, 856, 864, 871, 872, 874, 878, 880, 881, 888, 898, 908, 914 tessera 905, 906 Tethys 41, 131, 134, 135, 211, 215, 380, 424, 426, 438,
Index
440, 442, 465, 470, 487, 508, 522, 523, 524, 528, 553, 554, 556, 559, 560, 564, 570, 677, 678, 716, 844, 920 Tetrataxis 460 Tetrataxis nanus 678 Teufelsschlucht, Karawanken Mountains, Slovenia 418 Teutloporella 430 Teutloporella herculea 438 Teutloporella nodosa 438 Teutoburger Wald, Germany 450, 524 Texas 1, 120, 136, 222, 560, 836, 843, 844, 879 Textularia 678 textulariid foraminifera 455, 458, 465, 466, 468, 470, 472, 478, 616, 618, 628, 677, 762, 798 textural criteria 65, 66, 250, 328, 349, 575, 578, 591, 600, 896 textural inversion 357, 362, 694, 822 textural maturity 357 texture 6, 12, 16, 53, 56, 58, 61, 62, 68, 73, 76, 205, 277, 288, 361, 370, 377, 395, 576, 584, 607, 608, 635, 665, 666, 668, 673, 674, 680, 713, 754, 769, 786, 806, 826, 864, 890, 903 Thailand 424 thalamid sponges 492, 506 Thalassia 88 thalassinid crabs 547 thallus 413, 414, 416, 423, 430, 432, 448 Thamnasteria gracilis 512 Thamnopora 510, 692 Thamnopora boloniensis 764 Thamnopora reticulata 764 thamnoporid tabulate corals 362 Thanetian (Tertiary) 444, 650 Tharama 409 Thartharella 541 Thaumatoporella 385, 560, 562, 566, 568 Thecosmilia longimana 512 Theelia florida 554 thermal, thermic 67, 74, 334, 374, 623, 842, 900, 912 thermocline 8, 31, 94, 615 thin section 1, 62, 64, 65, 73, 221, 247, 310, 578, 832, 836, 893 thin-section diversity 579, 580 thin-section fossils 399, 480, 570, 571 thin-section measurement 244 thin-section microproblematica 559 thixotropy 774, 780 Thomas Church, Leipzig, Germany 900 thoracosphaerids 91, 800 thrombolite 76, 82, 83, 118, 364, 371, 372, 373, 374, 378 379, 385, 395, 408, 496, 742, 784, 830, 833 Thurburbo majus, northern Tunisia 553 Thuringia, Germany 742 tidal 24, 25, 57, 299, 369, 377, 514, 596646, 659, 663, 668, 698, 809, 893 tidal channel 75, 169, 170, 176, 235, 428, 708, 745, 750, 879 tidal current 229, 606, 663, 669, 702, 762, 820, 879 tidal flat 25, 31, 55, 60, 95, 185, 186, 193, 301, 302, 371, 604, 612, 665, 666, 681, 696, 704, 708, 714, 745, 746, 747, 762, 768, 804, 812, 817, 879, 880, 882 tidal flat island model 669 tidal intraclast 170
981
tidal zone 192, 509, 588, 641, 663, 704, 706, 745, 780 tidalite 745, 849 tide 8, 9, 96, 250, 332, 442, 635, 639,646, 659, 664, 665, 669, 804 tiering 610 tile 905 tillite 42, 630 time-averaging 105 time-series analysis 805, 807 tintinnids 440, 487, 490 Tintinnopsella carpathica 488 Tintinnopsella longa 488 titan yellow alizarine S 271 Tithonian (Jurassic) 218, 352, 358, 409, 410, 432, 440, 466, 483, 484, 490, 502, 512, 553, 556, 684, 714, 760, 766, 777, 872 Titriplex-III 66 Tivoli, Rome 16, 68 Toarcian (Jurassic) 215, 524, 528, 757 toe-of-slope 26, 229, 464, 662, 682, 686, 696, 716, 759, 771, 772, 784, 788 Tolypammina 382, 454 Tolypammina gregaria 678 Thommotian (Cambrian) 507 tool mark 56 top-and-bottom criteria 62, 177, 178, 180 Tornoceras 534 Torrecilla-Barranco, Zaragoza, Spain 698 touching-vug pores 283 Tournai, Belgium 150 Tournaisian (Carboniferous) 170, 432, 459, 460, 516, 842 trace element 37, 205, 290, 299, 587, 588, 616, 617, 619, 624, 641, 652, 653, 655, 737, 808, 866, 903 trace fossils 56, 60, 177, 185, 208, 313, 387, 388, 389, 395, 396, 589, 596, 610, 613, 614, 636, 748, 773, 797, 800, 802 traction, traction carpet, traction current 56, 247, 251, 253, 257, 775 trade route 903, 910 transect shelf-basin 720 shoreline-to-basin 659 transgression 168, 216, 223, 240, 452, 623, 625, 660, 731, 753, 804, 806, 809, 816, 838, 869, 882 transgressive 465, 641, 642, 659, 750, 808, 815, 892 transgressive lag 428 transgressive sequence 136, 820 transgressive surface 818, 819, 826 transgressive system 478, 523, 554 transgressive systems tract (TST) 675, 816, 819, 820, 826 transitional-marine 680, 708, 745, 875 Transmission Electron Microscopy (TEM) 66 transport 21, 104, 105, 150, 173, 174, 175, 182, 198, 230, 232, 240, 244, 248, 354, 356, 464, 468, 480, 482, 510, 523, 526, 551, 552, 576, 579, 591, 593, 596, 624, 630, 634, 641, 642, 646, 659, 660, 664, 666, 667, 673, 675, 694, 696, 700, 714, 716, 717, 723, 762, 766, 770, 774, 779, 788, 831, 842 trapping 19, 25, 177, 184, 342, 370, 371, 374, 376, 404, 690, 713, 747, 856, 868 travertine 2, 11, 15, 16, 18, 68, 361, 364, 371, 372,
982
410, 710, 734, 735, 736, 746, 898, 905; Plate 2 on p. 17 trees 381 Tremadoc (Ordovician) 869 Trepostomata, bryozoans 515 Tressenstein limestone, Jurassic 252, 253 Trevijano, Sierra de la Demanada, Spain 863 Triactoma kellumi 483 Triasina 678 Triasina hantkeni 464 Triassic 4, 88, 118, 134, 135, 136, 164, 166, 186, 221, 222, 296 298, 299, 306, 374, 378, 379, 380, 384, 385, 390, 392, 393, 395, 396, 406, 407, 412, 424, 426, 428, 438, 442, 444, 447, 464, 465, 474, 486, 494, 496, 498, 506, 508, 512, 515, 523, 524, 526, 530, 541, 553, 554, 558, 559, 562, 564, 566, 568, 576, 597, 604, 614, 640, 653, 657, 661, 664, 676, 680, 688, 690, 694, 696, 708, 716, 780, 781, 782, 788, 789, 790, 792,796, 798, 801, 803, 808, 810, 813, 814, 831, 835, 836, 864, 865, 871, 872, 874, 906, 907, 908 Trieste, Italy 526 trilobites 74, 104, 178, 184, 186, 196, 306, 404, 507, 510, 518, 538, 542, 547, 548, 598, 612, 737, 752, 764, 795, 918, 919, 921; Plate 94 on p. 549 in thin sections 548 microstructure 547, 548 Trinocladus perplexus 444 triple junction 290, 322 Triploporella 435, 440 Triticites 694 Tritrabs ewingi 484 Trnovo, southwestern Slovenia 268, 702 Trochammina alpina 678 Trochiliscus 448 Trocholina alpina 466 Trodoos Massif, Cyprus 800 Troglotella 559 Troglotella incrustans 563 Troms District, Norway 44 Tromsö/Norway 138 Tropfbruch Adnet, Salzburg, Austria 4, 306 trophic 341, 587, 624, 670 tropical 4, 31, 34, 40, 267, 270, 414, 426, 433, 442, 450, 458, 509, 569, 611, 615, 616, 623, 625, 626, 628, 630, 632, 634, 661, 710, 816, 830, 863, 869, 873, 874 trottoirs 416 Trucial Coast 14, 80, 81, 664 truncation surface 204, 779 Trypanites 392 TST 817, 820, 821, 824, 827, 828 TST/HST condition 824 tsunami 169, 588, 640 tube worms 51, 857 Tuberitina 454, 460 tuberoid 117, 362, 690 Tubiphytes 170, 222, 241, 346, 385, 428, 542, 559, 560, 562, 566, 568, 569, 781, 782, 785, 844 Tubiphytes morronensis 566, 569 Tubiphytes obscurus 564, 568 tubular fenestrae 193 Tubuliporata 515
Index
tufa , freshwater 11, 14, 15, 16, 21, 361, 364, 371, 410, 734, 735, 898 tufa stromatolite 374 tufa thrombolite 372 tunicates 104 Tunisia 182, 189, 394, 472, 474, 478, 502 turbidite 29, 46, 48, 50, 65, 118, 178, 184, 253, 354, 464, 486, 540, 554, 597, 659, 660, 662, 665, 686, 688, 698, 723, 759, 769, 770, 771, 772, 773, 774, 775, 777, 778, 779, 780, 784, 786, 789, 790, 796, 797, 798, 800, 821, 878, 880, 888, 908 crinoidal 552, 776 criteria 776, 778 distal 773, 775 practical advice 776 proximal 773, 775 turbidity 394, 587, 591, 597, 863, 864 turbidity current 19, 47, 49, 59, 172, 406, 480, 606, 646, 698, 714, 773, 774, 775, 781, 788, 802, 834 turbulence 49, 208, 384, 416, 607, 772 Turkey 18, 21, 334, 394, 414, 474, 510, 524, 526, 541, 560, 903 Turneffe Islands, Belize 668 Turonian (Cretaceous) 186, 396, 448, 450, 474, 526, 566, 827, 866, 871 Turrispirillina 678 Turtle Mountains, Frank Slide, Alberta, Canada 60, 516, 554 twinning 225, 297, 301, 306, 322, 325, 538, 554 Type-1, Type-2, Type-3 sequence boundary 818 types of skeletons 104 typhoon 594 U Udotea 427 udoteacean green algae 88, 405, 406, 422, 423, 424, 426, 427, 428, 444, 566, 582, 628, 635, 668, 686, 702, 764; Plate 57 on p. 424 ultra-thin section 57, 350 ultramicrite *91 ultraplankton 93, 684 ultrapure limestone 896 ultrasound transmission velocity 904 ultrastructure 66, 68, 74, 104, 106, 404 umber 801 unconformity 168, 169, 177, 203, 204, 205, 208, 212, 217, 334, 648, 734, 756, 757, 805, 816, 818, 819, 879, 882, 892; Plate 22 on p. 209 undersaturation 279, 292 Ungdarella 420, 422, 638 ungdarellacean red algae 406, 412, 413, 420, 422, 459, 638 United Arab Emirates 470, 524, 644, 646, 647, 884 Unken, Salzburg, Austria 466 Untersberg marble 242 Untersberg, Salzburg, Austria 512, 524 upward-shallowing 14, 817, 821 upwelling 31, 40, 49, 140, 455, 612, 623, 624, 625, 626, 634, 648, 656, 756, 798 Urals, Russia 379 Uranium 614
Index
983
Urgonian (Cretaceous) 465 UV epifluorescence 67 Uvanella 382, 498 Uyombo, Kenia 308 V Vaccinites vesicularis 528 Vaceletia 496 vadoid 13, *157, 158, 162, 619, 710, 824 vadose cap 813 vadose diagenesis 13, 120, 205, 604, 694, 812, 817 vadose ooids 157 vadose silt 96, 97, 178, 180, 196, 706 vadose zone 67, 76, 205, 276, 299, 746 Valanginian (Cretaceous) 442, 446, 468, 490 Valentintörl, Carnic Alps, Austria 316, 548, 630 valvulinid foraminifera 465 Variostoma crassum 790 Vaucheria 376, 735 vegetation 12, 731 vein, calcitic vein, veinlet 70, 172, 216, *222, 223, 224, 225, 232, 238, 320, 325, 889, 905 Velebit, Croatia 164, 306 Velebitella 401, 430 velocity-deviation log 815, 884, 885 Vendian (Precambrian) 135, 408, 875 Venetian Alps, Italy 797 Venezuela 877, 884 vent, vent and seep 623, 857, 858, 612 vent community 51, 850, 859 verbeekinid fusulinids 459, 464 Vercorsella camposauri 470 Vercorsella laurentii 470 Vercorsella scarsellai 470 vermetid gastropods 42, 530, 764, 856 vertebrates 404, 558, 649, 650, 735 Verticillodesmis 562 Vesicocaulis reticuliformis 498 vesicomyid bivalves 51, 859 viscosity 881 Viséan (Carbonifeous) 114, 146, 158, 379, 410, 420, 422, 432, 436, 450, 454, 459, 460, 462, 491, 510, 554, 637, 638, 730, 831, 842, 843, 850, 871 Visher method 251 visual comparison charts 257, 259, 282 vitrinite reflectance 337 void 16, 38, 63, 96, 118, 190, 217, 221, 290, 299, 301, 320, 358, 364, 497, 673, 704, 706, 708, 730, 735, 744, 747, 810, 824, 833, 835, 843, 883 void filling 271, 734 Vöksen/Deister, Lower Saxony, Germany 502 volcanic 55, 236, 289,623, 667, 736, 801, 863 volcaniclastics 542, 688, 782, 784, 801, 802, 862, 863, 864, 917 volcanism 47, 289, 533, 886 Volvocales 452 Vrhnica, Slovenia 128, 432, 440 vug 70, 206, 279, 283, 320, 326, 819, 882 W wackestone 74, 75, 89, 117, 135, 88, 196, 214, 312, 320, 348, 349, 350, 352, 362, 364, 366, 416,
427, 432, 436, 452, 459, 460, 464, 480, 482, 484, 486, 507, 523, 524, 526, 533, 534, 542, 548, 551, 576, 589, 595, 596, 597, 598, 614, 627, 637, 640, 659, 662, 663, 665, 666, 668, 672, 675, 676, 680, 682, 684, 692, 694, 698, 706, 708, 712, 716, 717, 719, 732, 737, 739, 740, 745, 747, 750, 752, 759, 760, 762, 764, 766, 785, 786, 788, 795, 796, 797, 798, 800, 801, 806, 807, 808, 809, 810, 812, 813, 819, 822, 824, 826, 827, 828, 834, 837, 842, 844, 849, 850, 852, 856, 858, 865, 882, 884, 899, 908; Plate 44 on p. 353 Wadell coefficient 174 Wadi Aday southwest Muscat, Oman 510 Wadi Al Ghadaf well, Jordan 692 Wadi Ala east of Jabal Kawr, Oman Mountains 523 Walanae Formation, Miocene 886 Walnut Canyon, Guadalupe Mountains 158, 843, 844 walnutoid 157, *158 warm-temperate 31, 446, 544, 450, 625, 626, 628, 630, 632, 634, 819 warm-water carbonates, environment 2, 34, 35, 78, 455, 593, 616, 625, 626, 627, 628, 632, 634 Wasserwaagen 178 water absorption 904 water depth 9, 121, 138, 308, 334, 379, 386, 394, 396, 406, 416, 442, 455, 456, 465, 486, 544, 576, 579, 587, 588, 600, 602, 603, 604, 628, 634, 635, 636, 637, 658, 662, 665, 669, 672, 675, 677, 681, 713, 724, 755, 787, 789, 794, 796, 798, 805, 806, 813, 817, 824, 848, 864 water energy 24, 40, 99, 101, 113, 136, 138, 153, 156, 250, 262, 304, 344, 346, 348, 356, 357, 379, 382, 424, 438, 442, 458, 460, 464, 492, 493, 500, 515, 559, 569, 570, 579, 587, 588, 589, 590, 591, 592, 593, 607, 666, 672, 674, 682, 700, 714, 717, 724, 740, 754, 755, 756, 806, 807, 848, 897 water energy index, level 96, 162, 356, 590 water escape structure 194, 224, 640 water table 12, 13 water temperature 7, 18, 512, 540, 612, 615, 616, 617, 626, 632, 634, 637, 816, 914 waterfall deposit 15 Waulsort, Belgium 842 Waulsortian facies, Carboniferous 196, 552, 638, 696, 718, 836, 842 Waulsortian mound 196, 379, 786, 605, 842, 843; Plate 144 on p. 841 wave 184, 250, 348, 526, 540, 588, 600, 606, 664, 665, 739, 746 wave abrasion 8, 588, 667, 697, 754, 806 wave action 24, 90, 174, 339, 354, 416, 480, 589, 591, 660, 672, 694, 879 wave base 8, 40, 551, 588, 593, 636, 637, 659, 662, 665, 667, 694, 712, 717, 754, 764, 766, 770, 806, 844 wave sweep base 354 Wealden facies, Cretaceous 450 weathering 11, 55, 60, 62, 63, 66, 83, 226, 278, 333, 623, 643, 728, 734, 866, 897, 898, 899, 900; Plate 150 on p. 901 weathering category 899 weathering resistance 895, 897, 898
984
Index
weathering type 896 weathering-potential index 898 well-log response 789, 885 Wenlockian (Silurian) 560, 871 Wentworth-Udden grain-size scale 243, 244 Wetheredella 382, 620, 921 Whiskey Creek Seep Carbonates, Eocene 859 Whiteinella 474 whitings 80 Wiestal quarry near Hallein, Austria 615 Wietersdorf, Carinthia, Austria 472 Williston Basin, Dakota 877, 879 Wilson Model 657, 658, 660, 661, 764, 768 wind, wind transport 12, 14, 92, 442, 450, 594, 742 winnowing 121, 194, 195, 211, 354, 361, 480, 523, 526, 589, 592, 606, 607, 663, 667, 696, 713, 725, 794, 795, 820, 850, 862, 896 Wirbelau near Weilburg/Lahn, Germany 289 wireline log 816, 877, 884, 885, 890 wispy seam 316 wood 404, 546, 559, 560, 908; Plate 97 on p. 561 Wordian (Permian) 424, 428, 462, 510, 563, 564, 690, 780 works of art 903, 907 worms 112, 185, 188, 390, 393, 394, 395, 396, 399, 427, 540, 541, 542, 564, 606, 690, 746, 785, 858, 859, 860 Würgau, northern Franconian Alb, Germany 541, 542 X X-ray diffraction 64, 70, 7271, 274, 735, 911 X-ray fluorescence 911 X-ray technique 2, 394, 458
Y Yabeina syrtalis 462 Yellowstone National Park 15 Yoho Valley, British Columbia 150 Yucatan, Mexico 12, 14, 24, 28, 252, 600, 602, 664 Yucca Mound, Sacramento Mts., USA 686 Yumakle, southern Turkey 334 Z Zapala, Neuquén Basin, Argentina 352 zebra fabric *196, 197 Zechstein, Permian 102, 136, 114, 284, 314, 379, 515, 516, 704, Zelenica, Begunjscica Mountains, Slovenia 374, 704 Zellia heritschi mira 462 zinc 857, 892 Zlambach beds, Triassic 432,497, 530, 688, 690 Zogelsdorf, Eggenburg, Austria 544, 630 zonation 68, 406, 430, 438, 459, 460, 468, 480, 515, 538, 671, 678, 864 zonation pattern 68, 418, 570, 635, 671, 789, 844, 844, 856; Plate 111 on p. 679 Zondarella communis 644 zone fossils 450 zone X, zone Y, zone Z 669 zooclast 101 zooplankton 93, 483, 487 zoospores 556, 563 zooxanthellate 51, 922 Zostera 88, 748
Filling the last page with a picture of a very ancient fossil: Maldeotaina (a stromatoporoid-grade sponge) from limestones of Precambrian-Cambrian boundary interval (Upper Krol Formation) of the Lesser Himalaya (India). Flügel and Singh (2003). Scale is 5 mm.
