Diagenesis in Sediments & Sedimentary Rocks (v. 2)

DEVELOPMENTS IN SEDIMENTOLOGY 25B

DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCKS, 2

FURTHER TITLES IN THIS SERIES VOLUMES 1, 2, 3, 5, 8 and 9 are out of print 4 F.G. T I C K E L L THE TECHNIQUES OF SEDIMENTARY MINERALOGY 6 L. V A N D E R P L A S THE IDENTIFICATION O F DETRITAL FELDSPARS I S. D Z U L Y N S K I and E.K. W A L T O N SEDIMENTARY FEATURES O F FLYSCH AND GREYWACKES P.McL.D. DUFF, A . H A L L A M and E.K. W A L T O N 10 CYCLIC SEDIMENTATION 11 C.C. R E E V E S Jr. INTRODUCTION T O PALEOLIMNOLOGY 12 R.G.C. B A T H U R S T CARBONATE SEDIMENTS AND THEIR DIAGENESIS 13 A.A. M A N T E N SILURIAN REEFS O F GOTLAND 14 K.W. G L E N N I E DESERT SEDIMENTARY ENVIRONMENTS C.E. W E A V E R and L.D. P O L L A R D 15 THE CHEMISTRY O F CLAY MINERALS H.H. R I E K E III and G.V. C H I L I N G A R I A N 16 COMPACTION OF ARGILLACEOUS SEDIMENTS M.D. PICARD and L.R. HIGH Jr. 17 SEDIMENTARY STRUCTURES O F EPHEMERAL STREAMS 18 G. V. C H I L I N G A R I A N and K.H. W O L F COMPACTION O F COARSE-GRAINED SEDIMENTS 19 W. S C H W A R Z A C H E R SEDIMENTATION MODELS AND QUANTITATIVE STRATIGRAPHY 20 M.R. W A L T E R , Editor STROMATOLITES 21 B. V E L D E CLAYS AND CLAY MINERALS IN NATURAL AND SYNTHETIC SYSTEMS 22 C.E. W E A V E R and K.C. BECK MIOCENE OF THE SOUTHEASTERN UNITED STATES 23 B.C. H E E Z E N , Editor INFLUENCE O F ABYSSAL CIRCULATION ON SEDIMENTARY ACCUMULATIONS IN SPA.CE AND TIME 24 R.E. GRIM and N . G U V E N BENTONITES 25A G. L A R S E N and G . V . C H I L I N G A R I A N , Editors DIAGENESIS IN SEDIMENTS AND SEDIMENTARY ROCK, 1 26 T. SUDO and S . SHIMODA, Editors CLAYS AND CLAY MINERALS OF JAPAN 27 M.M. M O R T L A N D and V.C. F A R M E R INTERNATIONAL CLAY CONFERENCE 1978 28 A. N I S S E N B A U M , Editor HYPERSALINE BRINES AND EVAPORITIC ENVIRONMENTS 29 P.TURNER CONTINENTAL RED BEDS T. SUDO, S . S H I M O D A , H. Y O T S U M O T O and S. A I T A 31 ELECTRON MICROGRAPHS OF CLAY MINERALS 32 C.A. N I T T R O U E R , Editor SEDIMENTARY DYNAMICS OF CONTINENTAL SHELVES

DEVELOPMENTS IN SEDIMENTOLOGY 25B

DlAGENESlS IN SEDIMENTS AND SEDIMENTARY ROCKS, 2 EDITED BY

GUNNAR LARSEN Department of Geology, University of Aarhus, Aarhus (Denmark)

AND

GEORGE V. CHILINGAR University o f Southern California, Los Angeles, Calif (U.S.A.)

ELSEVIER SCIENTIFIC PUBLISHING COMPANY 1983 AMSTERDAM -OXFORD -NEW YORK

ELSEVIER SCIENTIFIC PUBLISHING COMPANY Molenwerf 1, P.O. Box 211,1000 AE Amsterdam, The Netherlands Distributors f o r the United States and Canada: ELSEVIER/NORTH-HOLLAND INC. 52, Vanderbilt Avenue New York,N.Y. 10017

Librar? of ('ongrras ('ntaleging in tpUhlicptioti bata (Revised)

Main e n t r y under t i t l e :

Diagenesis i n sediments

and sedimentary rocks.

) (Developments i n s P d i a e n t o l o g y ; 2 5 A E d i t i o n of 1967 published under ti-tle: Diagenesis i n sediments. I n c l u d e s b i b l i o g r a p h i e s and i n d e x e s . 1. Diagenesis. I. Larsen, Gunnar, 192811. C h i l i n g a r i a n , George V . , 1929j oirit a u t h o r . 11. T i t l e . 111. S e r i e s . QE471.LsJ 1979 552' .5 78-23961 ISBN 0-444-41657-9 ( v . 1)

ISBN 0-444-42013-4 (Vol. 25B) ISBN 0-444-41238-7 (Series) Elsevier Scientific Puhlishing Company, 1983 All rights reserved. No part of this publication may be reproduced, stored in a retrieval system or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher, Elsevier Scientific Publishing Company, P.O. Box 330, 1000 AH Amsterdam, The Netherlands 0

Printed in The Netherlands

Dedicated to R.G.C. BATHURST, L. C A Y E U X K.O. EMERY, G.M. FRIEDMAN, A.P. LISITZYN, G,V. MIDDLETON, G. MILLOT, C. PRAY, S.G. SARKISYAN, H.C. SORBY and E. STEIDTMANN for their important contributions to the field of diagenesis

VI

CONTRIBUTORS

L. BUBENICEK

Societt Nationale Elf Aquitaine (Production), Pau (France)

G.V. CHILINGAR

University of Southern California, Los Angeles, Calif., U.S.A.

H.E. COOK

U.S. Department of the Interior, Geological Survey, Branch of Oil and Gas Reserves, Menlo Park, Calif., U.S.A.

R.M. EGBERT

U.S. Department of the Interior, Geological Survey, Branch of Oil and Gas Reserves, Menlo Park, Calif., U.S.A.

R.W. FAIRBRIDGE Department of Geological Sciences, Columbia University, New York, N.Y., U.S.A. H.J. KISCH

Department of Geology and Mineralogy, Ben Gurion University, University of the Negev, Beer Sheva, Israel

G . LARSEN

Department of Geology, University of Aarhus, Aarhus, Denmark

G . MULLER

Institut fur Sedimentforschung, Universitat Heidelberg, Heidelberg, Germany

A. SINGER

Department of Soils and Water, The Hebrew University of Jerusalem, Rehovot, Israel

VII

CONTENTS VI

Contributors . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 1. Introduction G.V. Chilingar and G. Larsen . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 2. Syndiagenesis- Anadiagenesis-Epidiagenesis: Phases in Lithogenesis R.W. Fairbridge . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . CHAPTER 3. Diagenesis in Argillaceous Sediments A. Singer and G. Miiller . . . . . . . . . . . . . . . . .

1

17

...........

115

CHAPTER 4. Diagenesis of Deep-sea Carbonates H.E. Cook and R.M. Egbert . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

213

CHAPTER 5. Mineralogy and Petrology of Burial Diagenesis (Burial Metamorphism) and Incipient Metamorphism in Clastic Rocks H.J .Kisch . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

289

APPENDIX A. Diagenesis of Iron-Rich Ores (Illustrated by the Role of Diagenesis in Oolitic Iron Ores) L.Bubenicek . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

495

APPENDIX B. Mineralogy and Petrology of Burial Diagenesis and Incipient Metamorphism in Clastic Rocks-Literature Published since 1976 H.J .Kisch . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

513

References Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

543

Subject Index . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . .

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1

Chapter 1 INTRODUCTION GEORGE V. CHILINGAR and GUNNAR LARSEN

THE EXPANDING INSIGHT

A revolutionary expansion of knowledge about sediments and sedimentary rocks has occurred during the last decade. This is largely due to the development of plate tectonics concepts, results of deep-sea drilling, and intensive studies by petroleum geologists. In the first edition of this work, the first one ever published on the subject of diagenesis, in 1967, it was pointed out that although the word “diagenesis” had been in the technical language for almost a century, it was not even listed in the general index of Encyclopedia Britannica: In the meantime, a new edition of the latter has appeared and the omission was corrected. Whereas in the first half of the present century. the subject could have been well-covered by a list of one hundred references, today we find well over that number appearing annually. According to Dr. Rhodes Fairbridge (personal communication, 198I ) , the revolutionary change in the knowledge and interest in diagenesis comes partly from need and partly from serendipitous discovery. The need emerged from the increasing emphasis being placed on the understanding of economic forces involved in the intensified search for petroleum, natural gas. coal, water, and other sedimentary raw materials. The discovery came hand-in-hand partly with the emergence of plate tectonics theory and, specifically, with the deep-sea drilling program. Rarely in the history of any discipline has its progress rested so squarely on the back of one single operation, the core-drilling of the “Glomar Challenger”. Repeatedly, from all parts of the world, this vessel has brought back long cores that not only contain the deep-sea stratigraphic history of the last hundred million years or so, but also disclose the actual development of diagenesis, an evolution that could previously only be deduced (R. Fairbridge, personal communication, 1981). It is also fortunate that when the cores get into the laboratory there are now a host of new techniques available for their study. In addition to the sieve-shaker, the binocular microscope, and the mineral stains of the earlier years, the big advance recently has been the development of the scanning electron microscope (SEM), which three-dimensionally supplements the information obtained by using the earlier-designed transmission electron mi-

2 croscope (TEM). There is also the luminoscope and the sophisticated equipment for isotopic analysis and trace-element identification. Arising in part from the surge of data, knowledge, and ideas, there has also appeared a number of textbooks (for example, Friedman and Sanders, 1978), important monographs, and the continuing series of Deep-sea Drilling Reports (DSDP). Also useful for those who need a quick capsule of information, alphabetically organized, is the Encyclopedia of Sedirnentoiogy of Fairbridge and Bourgeois (1978). A very convenient reprint collection of basic material of diagenesis compiled by Doris Curtis (1976) is entitled Sedimentary Processes: Diagenesis, and includes such classical works as those of ZoBell (1942), which brought out the role of bacteria, and Krumbein (1942), which established the multiplicity of physical and chemical changes. An important contribution is that of Davies and Supko ( 1973), which summarizes the results obtained during the first years of the DSDP-project. Another collection, volume 40 in the “Benchmark Series”, which was edited by Van der Lingen (1977), is devoted specifically to the hitherto little-known field of diagenesis of deep-sea biogenic sediments, now made accessible through the DSDP-drilling. The special type of diagenesis near sea-floor spreading centers is associated with hydrothermal systems, of particular interest to economic geologists. It has also been provided with a “Benchmark” treatment (Rona and Lowell, 1980).

ON THE DELIMITATION OF DIAGENESIS

An universally accepted definition of the term diagenesis does not exist. Many geologists define diagenesis as “all processes occurring between deposition and metamorphism”. This definition, however, is open for criticism, because in nature there are two distinct and different stages of alteration of rocks subsequent to deposition and prior to metamorphism: (1) conversion of sediments into sedimentary rocks, and (2) changes occurring in sedimentary rocks prior to metamorphism. One of the editors (G.V.C.) prefers to call the first stage “diagenesis” and the second stage “catagenesis”. These two stages differ not only in the character of physicochemical processes occurring in the materials composing sediments and sedimentary rocks, but also in the sources of energy required for the processes. The source of energy for diagenetic processes is the combination of various mineral components of the sediments and interstitial fluids. During the first (diagenetic) stage, an important role is played by the organic components (organic acids, etc.), microbial activity, solar energy, excess amounts of water present, etc. During the second (catagenetic) stage, on the other hand, the major role is played by temperature and pressure in

3 the alterations, which occur in the already formed sedimentary rocks. Briefly stated, diagenesis of sediments and catagenesis of rocks differ in sources of energy, physicochemical processes, and resulting products of alteration. As pointed out by Davies and Supko (1973), however, “many of the details of the vital link between the soft sediments and lithified sedimentary rocks remain a mystery.. .”. In the present book, the editors have found it necessary to leave the definition of diagenesis to individual contributors.

EFFECT OF SEDIMENTATION RATE ON DIAGENESIS

Studies on the influence of sedimentation rate on geological processes in contemporary sedimentary basins indicate that the degree of diagenetic alterations of sedimentary mineral components and associated organic matter in sedimentary deposits is a direct function of sedimentation rate and its duration (see Nazarkin, 1979). Sedimentation rate controls the degree of compaction of sediments and, consequently, determines direction, dynamics, and degree of organic matter alteration during diagenesis, which influence the intensity of hydrocarbon generation. As the sediments compact, dehydration occurs as a result of release of interlayer and adsorbed water. This dehydration intensifies hydrocarbon generation and migration. Associated liberation of heat supplements the petroleum-generation potential. Thus, in the case of low sedimentation rate one cannot expect presence of large oil and gas accumulations. Other factors being favorable, petroleum-generating potential of source rocks increases with increasing rate of sedimentation and its duration. Porosity and permeability of reservoir rocks would also be greater. Nazarkin ( 1979, p. 286) proposed the following three categories of petroleum-generating potential of paleobasins: (1) High petroleum-generating potential: average sedimentation rate ranges from 300 to 900 tons kmP2yr-I. (2) Average petroleum-generating potential: average sedimentation rate ranges from 160 to 300 tons km-2 yr-I. (3) Low petroleum-generating potential: average sedimentation rate ranges from 60 to 160 tons kmP2yr-’. The critical rate of sedimentation necessary for generation of sufficient quantities of hydrocarbons for accumulation of economical deposits is 60-100 tons kmP2 yr-I. The rate of sedimentation determines many diagenetic processes. In order to determine these relationships, however, they should be thoroughly studied.

4 SYNDIAGENESIS. ANADIAGENESIS A N D EPIDIAGENESlS

I n Chapter 2, Professor Rhodes W. Fairbridge recognizes three phases of diagenesis: syndiagenesis (penecontemporaneous), anadiagenesis (during burial and orogeny), and epidiagenesis (following emergence) as proposed by Fairbridge ( 1967). Slightly different (temporal) emphasis is expressed by the almost synonymous Choquette and Pray ( 1970) terms of eogeneric (early), mesogeneric (middle), and telogmetic (late) stage. The phase system expresses three conditions: time, place, and process. Initially, there is penecontemporaneous exchange with ambient media. Then there is a long-term exchange mainly with ascending and circulating connate and compaction waters, with an increasing role of temperature and pressure. Finally, there is the late-stage interplay with meteoric water, which often sinks to considerable depths with increased mixing at depth. Pervading each phase in the system is pore water, connate or meteoric, which is characterized by its dissolved components and by pH- and Eh-values. Besides precipitation and solution. there is reprecipitation (neoformation), authigenesis, and pseudomorphism. The first-phase waters start as normal lake or sea water, but become modified rapidly, either becoming supersaline as one goes down into the anadiagenetic realm, where temperature and pressure concurrently increase. or becoming fresher with depth due to compaction mechanism. According to Fairbridge, with uplift and re-exposure, the cycle reverses: temperature and pressure decrease and fresh water replaces or dilutes the saline fluids. In ancient formations, the cycle of anadiagenesis and epidiagenesis may be repeated again and again, but only one syndiagenesis cycle is possible. Cementation may, thus, be followed by leaching and then by recementation once more. In each phase, an equilibrium or steady state may be achieved. Events of geologic history, however, abruptly terminate the evolutionary development in any one of these phases, recycling or bypassing one or another. Carbonate diagenesis (treated in detail in Chapter 4 of this volume and Chapters 6 and 7 of Vol. I) is briefly discussed in Chapter 2. Carbonates are “peculiarly subject to many stages of diagenesis” (Wilson, 1975, p. 16) and are sensitive, with rapid response, to varying types of pore water. Remarkable advances have been made in this field during the last two decades, ranging from the actualistic settings and paleogeography to petrography and geochemistry. Bathurst ( 1 975, 2nd ed.) has provided the definitive study, whereas further treatments are available from Milliman (1974) and others. A contribution by Dr. L. Bubenicek on diagenesis of iron-rich ores is presented i n Appendix A.

5 DIAGENESIS IN ARGILLACEOUS SEDIMENTS

In Chapter 3, entitled “Diagenesis in Argillaceous Sediments”, Arieh Singer and German Miiller apply the term diagenesis to all changes which take place in a freshly deposited sediment until it reaches the stage of metamorphism. The following subjects are covered in that chapter: ( 1) initial composition of Recent argillaceous sediments, (2) initial (pre-burial) porosity and structure of argillaceous sediments, (3) changes in chemistry and mineralogy during diagenesis, (4)changes in structure during diagenesis, and ( 5 ) correlation of mechanical and chemical-mineralogical changes with depth of burial. The transitional zone between diagenesis and metamorphism is also discussed in that chapter. According to Singer and Miiller, argillaceous muds, with an initial porosity of 70-90%, are compacted during shallow-burial (burial depth down to 500m) diagenesis to mudstone (or shale if fissile) with a porosity of about 30%. Argillite with a porosity of < 10% forms on further compaction. The decreases in porosity and water content are rapid down to about 250-300 m burial depth and slow below that depth. Diagenesis passes into low-grade metamorphism when recrystallization of sheet-silicates occurs at about 200°C and about 10,000 m burial depth. Changes in porosity and water content are mainly a function of maximum overburden (burial depth) and of time. They are, however, also affected by geothermal gradients, tectonic stresses, lithology, depositional environments, overpressured fluid zones, and diagenesis. Under compaction, the individual particles or “domains”, making up the “card-house”-microstructure of dispersed muds in non-saline environments, attain a high degree of parallel orientation. The degree of orientation upon compaction of the much thicker domains (or oriented aggregates) making up the “book-house”-structure of flocculated muds, is less marked. During the pre-burial stage of marine diagenesis, ion exchange is the major chemical-mineralogical process. Exchangeable CaZi on the clays is exchanged for M g 2 + , K + and N a + from the sea water. Part of the adsorbed cations become non-exchangeable. Chlorite and illite-like minerals may thus be formed from degraded precursor clay minerals. Clay-mineral lattice transformations at this stage are minor. In the supersaline marine environment, transformations are more pronounced and include the transition of illites to chlorite through an intermediate stage of corrensite formation. In the supersaline lake environment, zeolites are the most prominent diagenetic products. Glauconite and manganese nodules are other products of pre-burial marine diagenesis. Zeolites, silica-minerals, palygorski te, and sepiolite form both during the pre-burial and shallow-burial diagenetic stages. Palagoni tc. smectites, and oxides form as a result of the submarine alteration of basalts and pyroclastics.

6 During deep-burial and compaction, the total electrolyte content of pore fluid decreases and its composition changes. With increasing depth, sediment pore water exhibits a strong depletion in Mg2+ and K + and a corresponding enrichment in C a 2 + . Dissolution of plagioclase and formation of Mg-rich smectites appear to be responsible for these changes. The deepest pore waters are enriched in dissolved SiO,. The major mineralogical process during deep-burial diagenesis is the progressive conversion of smectite into illite or into a mixed-layer illitesmectite, having a high proportion of illite layers, with increasing burial depth. This process is occurring in association with other minor changes, such as decomposition of coarser-grained mica and K-feldspar, decomposition of kaolinite, and formation of chlorite or chlorite interlayers. The fixation of K in newly-formed illite layers involves an increase in the net negative charge of the expandable layers, resulting either from the substitution of Al’+ for Si4+ in the tetrahedral layer or by the substitution of divalent for trivalent cations in the octahedral layer. The K cation necessary for the conversion is not supplied by pore solutions from a distant source, but is derived from K-feldspar and mica within the sediment. The K is thus being redistributed between the detrital and diagenetic phases. The conversion process appears to be directly related to temperature. Thus, reaction stages are controlled by burial depth and geothermal gradients. The magnitudes of the activation energy for the conversion process, determined experimentally, suggest that breaking of chemical bonds in the tetrahedral sheet is involved. Kaolinite is eliminated during deep-burial diagenesis, while chlorite and/or corrensite is formed, possibly by the interlayer adsorption of amorphous Fe20, and/or A1,0, onto expandable layer silicates. The deepest stage of diagenesis is characterized by a uniform clay-mineral association “illitechlorite”. During the transition to metamorphism it changes to a paragenesis quartz-albite-muscovite-chlorite. The burial diagenesis of silicate minerals is correlated with that of organic matter.

DIAGENESIS OF DEEP-SEA CARBONATES

As pointed out by H.E. Cook and R.M. Egbert in Chapter 4, four major processes modify physical properties, state of fossil preservation, texture, and geochemistry of deep-sea carbonates. These include: ( 1) gravitational compaction (mainly within the first 200m of burial); (2) dissolution (at the sediment-water interface, at shallow-burial depths, and some within the water column); (3) pressure-solution (mainly during the deep-burial phase); and (4) cementation (beginning within the first few centimeters of burial and

7 continuing with increasing burial depth and age). Pelagic carbonates are widespread and comprise about 70% of the carbonates deposited worldwide during the past 100 m.y. Deep-sea carbonates consist mainly of planktonic foraminifera, coccoliths, and discoasters whose tests are composed of stable low-magnesium calcite. This is in sharp contrast to the highly reactive aragonite and high-magnesium calcite that comprise shallow-marine carbonates. One of the most unique features of deep-sea carbonates is that they are generally not contacted by fresh water and, if they are, this does not occur until very late in their diagenetic history. In shallow-marine carbonates, early cementation, produced in part by contact with fresh water, preserves much of the original fabric and retards compaction in many cases. In deep-sea carbonates, on the other hand, because fresh water is absent and stable low-magnesium calcite is the main carbonate constituent, cementation is usually somewhat delayed and lithification by compaction is the rule within part of the sediment column. At shallow-burial depths of 50-200 m, gravitatiorial compaction is the dominant mechanism for reducing porosity from about 80 to about 60% with cementation being a subordinate pore-reducing process. Compaction in this shallow-burial stage takes place by closer packing of grains, some crushing of microfossil tests, and disintegration of delicate planktonic forams and coccoliths into abundant micron-sized crystals through dissolution along sutures. Discoasters appear to be the first organisms to develop secondary calcite overgrowth. This can occur within the first 35 m of burial before gravitational compaction produces a grain-supported texture, which suggests that dissolution-diffusion-reprecipitation processes occur early, while the grains are still dispersed. With increasing burial depth, a grain-supported fabric develops and pressure-solution becomes a major process in producing calcium carbonate for cementation, with gravitational compaction being a subordinate process. According to Cook and Egbert, the most advanced diagenetic state observed in Deep-sea Drilling Project cores is evidenced by the following features: (a) virtually all fossil grains are covered with subhedral to euhedral calcite cement, (b) the central areas of coccoliths are filled, (c) foram chambers are filled with granular calcite, (d) large amounts of interparticle cement are found, and (e) grain interpenetration and welding are common. Thus, at burial depths on the order of 1000- 1 200 m, porosities have been reduced to about 20%. Continued diagenesis produces an “ameboid mosaic” or “pavement mosaic” texture of completely interlocking grains. The degree to which a deep-sea carbonate becomes cemented, probably depends to a large extent upon the diagenetic potential the carbonate

8 sediment had at the time it was buried. Diagenetic potential is a measure of how much more diagenesis a sediment can undergo in the normal course of its history. There is a trend towards a decreasing Sr2+ content and increasing S " 0 values (negative) in pelagic carbonates with increasing burial depth. This is to be expected as most deep-sea carbonates are diagenetically altered after burial. With increasing burial depth, dissolution-reprecipitation processes reduce porosity. The Sr2+ originally present in the biogenic calcite is virtually excluded from the secondary cement derived from the dissolution and pressure-solution of calcite microfossils. The S " 0 values of this cement become increasingly more negative with burial depth, due to precipitation of the cement from pore fluids at increasingly elevated temperatures along the geothermal gradient. Acoustic reflectors, which are probably related to the degree of cementation in the sediments, may be predetermined in their broad characteristics by major oceanic events, which in turn largely control the nature (diagenetic potential) of the biogenic material reaching the sea floor. Deep-sea chalks form excellent petroleum reservoirs under certain circumstances: (a) when their porosity is retained at deep-burial depths if they are hydrostatically overpressured, which decreases grain-to-grain stress and, consequently, retards pressure-solution processes; (b) when early oil emplacement may retard dissolution-reprecipitation processes; (c) by lack of deep burial depths; and (d) when fracturing enhances the chalk's permeability.

BURIAL DIAGENESIS AND INCIPIENT METAMORPHISM

In Chapter 5 , Dr. H.J. Kisch reviews the mineralogical changes taking place during burial diagenesis and incipient metamorphism from the point of view of a metamorphic petrologist, stressing the attainment of equilibrium, the compositional constraints on the reactions, and the relation of the stages of burial during which the various changes take place. The similarities and largely gradual differences between diagenetic and metamorphic reactions are discussed, particularly regarding the degree of attainment of chemical equilibrium and the extent of mineralogical and textural reconstitution. Although persistence of metastable clastic relics and of clastic textures is more common at low temperatures due to kinetic factors (slow reaction rates), the processes operating in burial diagenesis are concluded to be essentially metamorphic in nature. Kisch gives a general introduction to the concept of diagenesis, the divergent uses of the term and its synonyms. He points out the absence of a

9 natural delimitation between late diagenesis and metamorphism and stresses the arbitrariness of such delimitations. Particular attention is paid to the various petrographic, textural, and mineral-facies criteria proposed for the distinction between diagenetically altered sedimentary rocks and metamorphic rocks. Following Coombs (1954, 1958, 1960, 1961, 1971), Kisch uses the degree and type of textural modification for a loosely defined subdivision of deep-burial phenomena into: ( 1) burial metamorphism, and (2) incipient regional metamorphism. Coombs’ term “burial metamorphism” is used for partial or complete reconstitution of clastic rocks on a regional scale, characteristically without development of a penetrative fabric. Some investigators consider burial metamorphism a synonym of burial, late, or epigenetic diagenesis (epigenesis or catagenesis of some authors I ) . “Initial metamorphism” is used for the more advanced stages of mineral modification, of which slaty cleavage is a common, though not an essential, attribute. It has very general mineral-facies connotations, including most of the prehnite-pumpellyite facies, includes terranes showing the attributes of Kubler’s (1967b) anchimetamorphism, and is considered by some authors to be approximately equivalent to the stage of “early metagenesis”. The second part of Chapter 5 deals with the modification of clay mineralogy of clastic sedimentary rocks upon deep burial. The following burialdiagenetic processes are reviewed: (a) The progressive replacement of smectite by illite through random illite-smectite mixed-layers and, commonly, through a subsequent stage of regular or partly ordered illite-smectite mixed-layers. Depth and temperature of the onset of mixed-layering, the non-continuous nature of the dehydration process, and the related question of the stability of illite-smectite mixed-layers are also discussed. (b) The progressive replacement of smectite by chlorite through chloritesmectite mixed-layers, including the regular mixed-layer corrensite. (c) Changes in the polytype of kaolinite-group minerals. (d)The replacement of kaolinite by illite and/or chlorite and, less commonly, by dioctahedral, aluminous chlorite. (e)Changes in the polytype of illite, and its dependence on the nature of its predecessor (illite-smectite mixed-layer; kaolinite) and on illite composition. ( f ) Changes in composition and polytype of trioctahedral chlorite. The

’

Possibly “early rnetagenesis” of Soviet authors (see discussion in the Introduction chapter of Vol. I of this book).

10 dependence of the various modification processes on compositional variables is stressed, including the effect of bulk composition, particularly the breakdown of clastic minerals such as biotite, K-feldspar, and plagioclase. The importance of the composition of interstitial solutions and, hence, of porosity and permeability is discussed. The following ongoing changes are reviewed during the incipient metamorphsm or anchimetamorphism, when random illite-smectite mixed-layers and kaolinite have disappeared and illite and chlorite have become the predominant layer-silicate minerals: (a) Composition and polytype of illite-muscovite. (b) Determination of the progressive increase in illite crystallinity (and its complicating factors), and the use of illite crystallinity in defining an anchimetamorphic zone. (c) Appearance of pyrophyllite and conditions of the kaolinite pyrophyllite-quartz equilibrium. (d) Appearance of Na-rich illite-smectite mixed-layers (rectorite or allevardite) and Na-rich, paragonitic micas. (e) Appearance of some other minerals locally associated with the anchimetamorphic zone, such as stilpnomelane and chloritoid. (f) The rank of coal associated with anchimetamorphism. Finally, tentative schemes of lowest-grade mineral facies for clastic rocks are presented. The third part of Chapter 5 deals with lowest-grade metamorphic facies in volcanic and lithic-feldspathic sedimentary rocks, as defined by the appearance of diagnostic zeolites such as laumontite and, at more advanced grades, by prehnite, pumpellyite, etc., without zeolites. Some well-documented cases of mineral zoning are described, including the classical terranes of South Island, N.Z., and Honshii, Japan. Particular attention was paid by Kisch to the various compositional controls on the zonal distribution of various burial-metamorphic zeolites and other Ca-Al-silicate minerals and telescoping of their distribution zones. He discusses: ( 1) selective replacement of different primary mineralogicpetrographic constituents, such as the preferential replacement of silicic pyroclastics by alkali-zeolites and analcirne; (2) persistence of laumontite into the prehnite-pumpellyite zone in silicic tuffs; and (3) dissimilar extent of reconstitution in rocks of contrasting mineralogic composition, including the notion of mineral facies. The various zeolite-facies equilibria are discussed, followed by a discussion of the constraints upon their application, which include: ( 1) reduced fluid pressure (osmotic conditions) and the consequent effect of porosity and permeability differences; (2) salinity and alkalinity of the fluid phase; (3) oversaturation in silica and ionic activity gradients controlled by hydrolysis

11 and solution of acid volcanic glass upon burial; (4) ionic equilibria and coupled solid-solution reactions, and ( 5 ) variation of the pco,/p H , O ratio (hydrous Ca-Al-silicates may be suppressed at high ratios). A number of distinctive types of progressive successions of lowest-grade metamorphic mineral assemblages or “facies series” are distinguished, based on the experimental evidence on the pressure dependence of diagnostic equilibria in the pumpellyitic facies. The final sections of Chapter 5 deal with the rank of coal and with associated clay minerals in the burial-metamorphic facies. Kisch concludes that the laumontite zone of the zeolite facies is associated with coals of not higher than low-volatile bituminous rank, and that it is entirely within the stage of “deep epigenesis” of Kassovskaya and Shutov (1963) or the zone of “diagenesis” of Kubler (1964, 1967a,b, 1968, 1970). The prehnitepumpellyite facies starts at a similar or somewhat lower rank than the stage of “early metagenesis” or the anchimetamorphic zone. Relations between illite crystallinity and coal rank and the onset of pumpellyitic facies are different in high-pressure, lawsonite-bearing terranes. Coordinated studies of the different kinds of parameters of burial diagenesis and incipient metamorphism are likely to contribute to the understanding of pressure- temperature gradients in burial diagenesis and incipient metamorphism, and to contribute to the recognition of thermal events and tectonic movements.

THE DOLOMITE PROBLEM

Although dolomitization was discussed in detail in Vol. I of this book, the following information sheds additional light upon the formation of dolomite, which is still one of the great puzzles of geology. Sea water is supersaturated with respect to dolomite and yet it does not appear to precipitate there. In addition, dolomite has not been synthesized in the laboratory under the present-day earth-surface conditions. Much of the dolomite forming at the present time and that of the geological record, at least the Phanerozoic, has formed by replacement. One particular dolomite problem concerns the dominance of dolostones over limestones in the Precambrian. Two basic hypotheses have been put forward to explain this: (1) age: being older, the original CaCO, sediments have had more chance to come into contact with dolomitizing fluids; or (2) different composition of sea water and of atmosphere: Precambrian sea water and atmosphere, unlike those of the Phanerozoic, permitted dolomite precipitation and/or extensive dolomitization of limestones. New information on the Precambrian dolomite problem has recently been presented by Dr. Maurice Tucker (University of Newcastle-upon-Tyne;

12

personal communication, 198l), based on a petrographic-isotopic study of the Late Proterozoic Beck Spring Dolomite of eastern California. Tucker’s results show that this Precambrian dolomite, and possibly many others too, are different from Phanerozoic dolomites in terms of their fabrics and isotopic ratios. The Beck Spring Dolomite (? 0.9 to 1.2 billion years old) is a typical platform carbonate with intertidal and subtidal facies of cryptalgal laminites, stromatolites, micrites, and grainstones (of peloids, ooids and pisolites). The most conspicuous feature of this dolomite is the preservation of fabric details. On a microscopic scale, all fabric details are preserved, even though the carbonate mineral is dolomite. There is no pervasive or selective dolomite replacement of calcite, as is usually the case with Phanerozoic dolomites (see figs. 7-41 to 7-50 in Vol. 25A in the Series “Developments in Sedimentology”). Ooids have a radial-fibrous and concentric structure identical to the calcitic ooids of the Phanerozoic. Diagenetic cements are directly comparable to those of limestones with the early isopachous fibrous dolomite crusts, which are present around grains and line cavities. There are also later sparry dolomite cavity fills, with delicate growth zones. The petrography shows that Phanerozoic-type dolomitization of an original limestone has not taken place. Carbon and oxygen stable isotope analyses of depositional grains (pisolites and micrite), early fibrous dolomite cements, and later sparry dolomites confirm that dolomitization has not taken place. The various components possess distinctive isotopic ratios. They exhibit a distinct trend towards lighter, more negative 6I8O and, to a lesser extent, towards lighter 6I3C, from the grains, through the fibrous dolomite, to the dolomite spar (Fig. 1-1). This isotopic trend is directly comparable to that obtained from the equivalent components in the Phanerozoic limestones (e.g., Walls et al., 1979). The 6I8O and 613C values of the depositional grains and micrite reflect the isotopic composition (and temperature) of sea water at the time. The ratios for the fibrous dolomite cements indicate a marine origin with some later recrystallization in lighter waters. The dolomite spar with more negative 6 ’ * 0 reflects precipitation at higher temperatures during burial. The decreasing S13C indicates a contribution of ” C from diagenetic decomposition of organic matter. Although many geologists believe that practically all Recent and ancient dolomites are replacive in origin, the isotopic and petrographic data of Maurice Tucker (personal communication, 1981) from the Beck Spring Dolomite of eastern California are best interpreted in terms of primary dolomite precipitation. Dolomitization is now believed to be by most geologists as being a wet process involving dissolution-precipitation, and in as much as water is involved, isotopic exchange will occur. Retention of

13

tl

+E

t6

15

+5

+l

6 1

t3

k 3

+2

,2

+1

11

0

0

-1

-1

-2

-2 I

-3 17

10

19

20

21

22

23

2L

&”o

25

26

27

28

29

30

31

-3

32

SMOW

Fig. 1-1. Carbon and oxygen isotope results from the Beck Spring Dolomite, late Proterozoic. eastern California. Symbols: 0 =pisolites; M =micrite; A =cryptalgal laminite (also micrite); =fibrous dolomite; V =sparry dolomite; I S =internal sediment. (Courtesy of Dr. Maurice Tucker, personal communication, 198 1.)

original isotopic ratios effectively rules out a major dolomitizing event, as does the fabric evidence for lack of replacement. An interpretation of original dolomite for this Late Proterozoic dolostone implies that the composition of the Precambrian sea water, which permitted primary dolomite precipitation, was somehow different. Higher Mg/Ca ratio, higher p C 0 , and lower SO:- content are all possibilities. Finally, it should be mentioned in this Introduction that the present Volume I1 of “Diagenesis in Sediments and Sedimentary Rocks” will be succeeded by a third volume, which will cover the following subjects: Compactional Diagenesis, Diagenesis of Ore Deposits, Diagenesis of Evaporites, and Diagenesis of Organic Matter and Generation of Oil.

REFERENCES A N D BIBLIOGRAPHY Ali, S.A. and Friedman, G.M. (Compilers), 1977. Diagenesrs of Sandstones. Am. Assoc. Petrol. Geol., Reprint Ser.. 20: 239 pp.

14 Aoyagi, K., 1979. Paleo-temperature analysis by authigenic minerals in sedimentary rocks. J. Jpn. Assoc. Petrol. Technol., 44(6): 1-5. Aoyagi, K. and Kazama, T., 1980. Sedimentary mineralogy of argillaceous sediments from deep-sea drilling project holes 436, 438A, and 439, Japan Trench. In: Honza et al. (Editors), Initial Reports of the Deep-sea Drilling Project, 56/57. U S . Gov. Print. Off., Washington, D.C., pp. 101 1-1017. Bathurst, R.G.C., 1975. Carbonate Sediments and Their Diagenesis. Elsevier, Amsterdam, 2nd ed., 658 pp. Choquette, P.W. and Pray, L.C., 1970. Geologic nomenclature and classification of porosity in sedimentary carbonates. Bull. Am. Assoc. Pet. Geoi., 54: 207-250. Coombs, D.S., 1954. The nature and alteration of some Triassic sediments from Southland, New Zealand. R. SOC.N. Z. Trans., 82( 1): 65- 109. Coombs, D.S., 1958. Zeolitized tuffs from the Kuttung Glacial Beds near Seaham, New South Wales. Aust. J. Sci., 21: 18-19. Coombs, D.S., 1960. Lower-grade mineral facies in New Zealand. Rep. Int, Geol. Congr., 21st Sess. (Norden), 13: 339-351. Coombs, D.S., 1961. Some recent work on the lower grades of metamorphism. Aust. J. Sci., 24(5): 203-215. Coombs, D.S., 1971. Present status of the zeolite facies. Ado. Chem. Ser., 101: 317-327. Coombs, D.S. and Whetten, J.T., 1967. Composition of analcime from sedimentary and burial metamorphic rocks. Geol. Sac. Am. Bull., 78: 269-282. Coombs, D.S., Ellis, A.J., Fyfe, W.S. and Taylor, A.M., 1959. The zeolite facies. with comments on the interpretation of hydrothermal syntheses. Geochim. Cosmochim. Acta, 17: 5 3 4 07. Coombs, D.S., Horodyski, R.J. and Naylor, R.S., 1970. Occurrence of prehnite-pumpellyite facies in northern Maine. Am. J . Sci., 268: 142-156. Curtis, D.M. (Compiler), 1976. Diagenesis. S.E.P.M. Reprint Ser. 1: 216 pp. Davies, T.A. and Supko, P.R., 1973. Oceanic sediments and their diagenesis: some examples from deep-sea drilling. J. Sediment. Petrol., 43(2): 38 1-390. Fairbridge, R.W., 1967. Phases of diagenesis and authigenesis. In: G. Larsen and G.V. Chilingar (Editors), Diagenesis in Sediments. Elsevier, Amsterdam, pp. 19-89. Fairbridge, R.W. and Bourgeois, J., 1978. The Encyclopedia of Sedimentologv. Dowden, Hutchinson and Ross, Stroudsburg, 901 pp. Friedman, G.M. and Sanders, J.E., 1978. Principles of Sedirnentoiogy. Wiley, New York, N.Y., 792 pp. Kassovskaya, A.G. and Shutov, V.D., 1963. Facies of regional epigenesis and metagenesis. Izv. Akad. Nauk, S.S.S.R., Ser. Geoi., (7): 3-18. Krumbein, W.C., 1942. Physical and chemical changes in sediments after deposition. J. Sediment. Petrol., 12: 111-1 17. Kubler, B., 1964. Les argiles, indicateurs de metamorphisme. Reu. Inst. Fr. Pet., 19: 1093- 11 12. Kubler, B., 1967a. La cristallinite de I’illite et les zones tout a fait superieures du metamorphisme. In: Etages tectoniques. A la Baconnikre, Newhatel, pp. 105- 121. Kubler, B., 1967b. Anchimetamorphisme et schistosite. Bull. Cent. Rech. Pau, l(2): 259-278. Kubler, B., 1968. Evaluation quantitative du metamorphisme par la cristallinite de M i t e . Bull. Cent. Rech. Pau, 2(2): 385-397. Kubler, B., 1970. Crystallinity of illite. Detection of metamorphism in some frontal parts of the Alps (abstr.). Fortschr. Mineral., 47 (Beih. 1): 39-40. Kubler, B., Martini, J. and Vuagnat, M., 1974. Very low-grade metamorphism in the Western Alps. Schweiz. Mineral. Petrogr. Mitt., 54(2/3): 46 1-469.

15 Milliman, J.D., 1974. Marine Carbonates, 1 . Recent Sedimentary Carbonates. Springer, New York. N.Y., 375 pp. McBride, E.F. (Compiler), 1979. Diagenesis of Sandstone: Cement - Porosity Relationships S.E.P.M. Reprint Ser., 9: 233 pp. Nazarkin, L.A., 1979. Influence of Sedimentation Rate and Erosional Sections on Oil and Gas Potentials of Sedimentary Basins. Saratov Univ., 336 pp. Rona, P.A. and Lowell, R.P. (Editors), 1980. Seafloor Spreading Centers: I+drothermal Systems. Dowden, Hutchinson and Ross, Stroudsburg, Pa., 424 pp. Van der Lingen, G.J. (Editor), 1977. Diagenesis of Deep-sea Biogenic Sediments. Benchmark Papers in Geologv, 40. Dowden, Hutchinson and Ross, Stroudsburg, Pa., 385 pp. Walls, R.A., Mountjoy, E.W. and Fritz, P., 1979. Isotopic composition and diagenetic history of carbonate cements in Devonian Golden Spike reef, Alberta, Canada. Bull. Geol. SOC. Am., 90: 963-982. Wilson, J.L., 1975. Carbonate Facies in Geologic Histoty. Springer, Berlin, 471 pp. ZoBell, C.E., 1942. Changes produced by micro-organisms in sediments after deposition. J . Sediment. Petrol., 12: 127- 136.

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17 Chapter 2

SYNDIAGENESIS-ANADIAGENESIS-EPIDIAGENESIS: LITHOGENESIS

PHASES IN

RHODES W. FAIRBRIDGE

INTRODUCTION

Diagenesis of sediments is understood by the writer as those physical and chemical changes that the sediment undergoes during and after deposition and lithification, without introduction of great heat or pressure. Three phases of the overall rate process were defined (in the 1967-edition of this work) bn a temporal and spatial basis. In an ideal model, a dynamic evolution takes place within the constraints of increasing time, depth of burial or exposure, and changing hydrologic-geochemical systems. Syndiagenesis is the early phase, synchronous with deposition and early burial, being strongly influenced by biochemical agencies in a fluid regime dominated by the entrapped waters of the sea- or lake-floor. Anadiagenesis is the deep-burial phase marked by compaction and maturation, strongly influenced by increasing pressure and the upward expulsion or transit of connate waters that are often highly mineralized or saline. Epidiagenesis follows uplift or emergence that brings surficial, meteoric waters into circulation often displacing pre-existing fluids and reversing numbers of geochemical processes. Eustatic fluctuations or tectonism may accelerate or short-circuit the ideal triple-phase evolution. The term “diagenesis” has been in the geological literature for over a century, and it is hardly surprising, therefore, that it has evolved in its meaning. First introduced by Von Guembel (1868, p. 838), it comprised all post-sedimentary modifications of mineral constituents, ultimately to the level of metamorphism. This very broad definition was never adopted, and indeed Van Hise (1898, 1904), probably in ignorance of Von Guembel’s term, proposed that “metamorphism” for its part, should cover the identical field. By tacit rather than explicit agreement among geologists, metamorphism generally is taken to begin where diagenesis ends, i.e., at a temperature of around 200°C. There is clearly an area of overlap of about 10O-30O0C, depending upon the individual writer’s point of view (see, for example, Fyfe et al., 1958). In the new glossary of Bates and Jackson (1980), “diagenesis” is divided into a mineralogical usage and a sedimentological usage. In the former, it refers to any geochemical or crystal rearrangement of minerals (here they

18 specifically cite clay minerals, but this exclusivity is not necessary) that may affect the particles either before or after burial. In the sedimentological usage -the more usual approach-it means “all the chemical, physical and biological changes undergone by a sediment after its initial deposition, (both) during and after its lithification, exclusive of surficial alteration (weathering) and metamorphism”. Further.. . “it embraces those processes (such as compaction, cementation, reworking, authigenesis, replacement, crystallization, leaching, hydration, bacterial action, and formation of concretions) that occur under conditions of pressure (up to 1 kb) and temperature (maximum range of 100°C-300°C) that are normal to the surficial or outer part of the Earth’s crust; and it may include changes occurring after lithification under the same conditions of temperature and pressure”. In this sense the term was effectively utilized and popularized by Walther (1894, pp. 693-71 l), the “father of sedimentology”, as an actualistic science. In this meaning it was widely adopted in the English-speaking world (Twenhofel, 1932), in Germany (Correns, 1950), and France (Dunoyer de Segonzac, 1968). In the U.S.S.R., however, most geologists restrict the expression to cover only the initial, uncompacted phase (Fersman, 1922); after lithification, the term “epigenesis” (or more recently “catagenesis”) is employed (pers. comm. G.V. Chilingar). Some western writers have also adhered to this narrow approach, but the Bates and Jackson (1980) synthesis is preferred by the present author. Diagenesis is restricted to sediments and sedimentary rocks, but the concept of lithification may refer also to igneous rocks. Diagenesis embraces both post-depositional mechanical modification and geochemical reorganization. It has widespread application; thus, glaciologists speak of the diagenesis of snow, in the sense of compaction, secondary cementation, and recrystallization (Anderson and Benson, 1963), whereas organic geochemists speak of the diagenesis of the products of organic metabolism leading to the formation of petroleum hydrocarbons (Breger, 1960). One may note that the physical state of the latter may be solid, liquid or gaseous.

LITHIFICATION

Lithification is the most striking aspect of diagenesis. I t is the “conversion of a newly deposited, unconsolidated sediment into a coherent, solid rock, involving processes such as cementation, compaction, desiccation, crystallization ... concurrent with, soon after, or long after deposition” (Bates and Jackson, 1980). Walther ( 1894) recognized that soft sediments did not become hard and “lithified” into rock merely by the action of time. Hutton ( 1788) believed that all lithification was due to heat and pressure that

19 involved partial fusion following deep burial. With his far greater experience of the younger formations and contemporary sedimentation, Walther was able to discern a continuum that led to his deduction that there were specific chemical and physical processes characteristic of diagenesis that took place under low temperature and pressure conditions. Lithification, according to Grabau (1913, p. 75 1) includes the following (with comments by the writer): (1) Congelation (e.g., the physical dehydration and hardening of silica or organic gels to form an amorphous or cryptocrystalline solid). (2) Crystallization (the primary reorganization of compounds, as in unstable pyroclastic sediments, but without participation of the interstitial water or of other minerals). (3) Recrystallization (a secondary crystallographic reorganization of the minerals, under increasing stress or other influence, e.g., the inversion of aragonite to calcite). ( 4 ) Compaction, welding and pressure cohesion (as a result of progressive loading, interstitial fluids being squeezed out and grains brought into contact, sometimes leading to local contact solution and redeposition in the voids). ( 5 ) Cementation (filling or partial filling of voids by cements, mainly CaCO,, SiO, and Fe203, derived from circulating waters. Additional factors in diagenesis have been noted by Andree (191 1). They include: (a) formation of concretions (both in the pre- or post-lithification phases), and (b) desalinification (a post-lithification phenomenon involving the leaching and sluicing out of connate waters by vadose waters, i.e., circulation of meteoric origin, ground water and artesian circulation. Hydrothermal or pneumatolytic circulation is certainly excluded). Special geochemical factors in diagenetic mineralization, recognized partly by Grabau (1913, p. 750) and others, include: ( 1) Low-temperature metasomatism which embraces mineral replacement (e.g., limestone by silica and vice versa; also dolomitization). ( 2 ) Hydration and dehydration (e.g., the transformation of gypsum, CaSO, .2H,O, to anhydrite, CaSO,, which is thought to occur under a load of around 100 m of sediments). ( 3 ) Ion exchange (typical of the clay and mica families of minerals). (4) Polymerization and depolymerization (e.g., natural catalytic “cracking” and other organic chemical reactions characteristic of the hydrocarbons). AUTHIGENESIS

After lithification, the next most striking aspect of diagenesis is the formation - of new minerals within an enclosing sediment or sedimentary

20 TABLE 2-1 More common authigenic minerals (excluding halides) I Mineral

Formula

Frequency

Usual development synanaepidiagen- diagen- diagenetic etic etic

Anatase Anhydrite Ankerite Aragonite Azurite Barite Bornite Brookite Calcite Celestite Cerussite Chalcedony Chalcopyrite Chamosite Chlorite group Collophane Dahllite Dolomite Galena Glauconite Greenalite Gypsum Halite Hematite Hydromagnesi te Illite

X

X

X X X

X

X

X X X X

X

X X X X X

X

X

X X

X X

x

X

X

X

X

.

X X

X X X X X

X

X X X

X

X

X X X

X

X X X

X

X X

Kaolinite Leucoxene Limonite

X'

X

X Magnesi te Malachite Marcasite Mon tmorilloni te Muscovite Natron Nesquehonite Opal

X

X

X

x

X X X X

X X

X X

21 TABLE 2-1 (continued) Mineral

Formula

Frequency

Usual development synanaepidiagen- diagen- diagenetic etic etic

Orthoclase Phillipsite Plagioclase Psilomelane (Wad) Pyrite Pyrolusi te Quartz Rhodochrosi te Rutile Siderite Sphalerite Strontianite Sulphur Tourmaline

KAISi,O, (Ca,Ba,K,Na),Al x(Al,Si)2 SiloO,o. 15-20 H,O (Ca,Na)(Al,Si)AlSi ,O,

Wi therite Zeolites (Phillipsite, Heulandite, Laumontite, Chabazite, Natrolite, Analcime) Zircon ZrSiO,

C

X

R C

X

C C C C

X

R C R R R

X X X X

R R

X

C R

X

X

’ In part, after Twenhofel (1950, p. 288) and Teodorovich (1961). ’ R=rare; C=common. rock. It may be achieved by replacement (metasomatism), recrystallization, or by the filling of voids. Some authors would include secondary enlargement as in quartz overgrowths. As proposed by Kalkowsky (ISSO), this is authigenesis, the process of forming any new or secondary mineral, but now specifically in a sedimentary environment. It is also termed “neoformation” or “neogenesis” in the U.S.S.R. and France (Millot, 1970). We exclude from the term “authigenesis” such processes as solution and re-precipitation phenomena, simple cementation, decementation and intrastratal solution, including, for example, cone-in-cone, and overgrowths (addition of like ions). Pettijohn ( 1957, p. 650) distinguished diagenetic metasomatism from authigenesis. Diagenetic metasomatism seems to be part and parcel of his “diagenetic differentiation”, which involves the redistribution of materials within

22 a sediment, such as the formation of nodules and concretions. In such event, an existing mineral type, e.g., calcite may assume a new form or position (as distinct from a new type of mineral), but the solutions can well be supplied from without, i.e., from the motion of connate waters or even vadose waters. Tester and Atwater (1934) emphasized that such minerals must be regarded as discrete crystallographic units, rather than rock-forming components. There has thus been some tendency to restrict “authigenesis” to refer to the generation of “exotic” minerals, other than those forming the bulk of the rock (e.g., the clays), but this narrow interpretation is in no way implied by its original definition. A comprehensive listing of such minerals and discussion of their origin has been provided by Teodorovich (1961). During diagenesis, new minerals often form as a result of a reaction between the ions of the interstitial water and the primary particles. For example, the reactions may involve the simple addition of some new and different ions (adsorption), the exchange of ions, or the replacement of certain ions by some new ions (metasomatism). The secondary overgrowth of some new ions onto an existing mineral of the same composition does not give rise to a new mineral. Some of these new minerals are so characteristic of primary, magmatic or high-temperature metamorphic phases that their “exotic” appearance rarely fails to cause surprise. Pustowaloff (1955) has drawn attention especially to zoisite, clinozoisite, epidote and sphene, whereas Wetzel ( 1955) has noted cinnabar (HgS) and barium minerals. Authigenic reactions are here interpreted as all those reactions leading to new mineral formation. They include: (a) metasomatism (ionic replacement); (b) ion exchange and adsorption (addition of new, and/or exchange of different ions, especially base exchange); (c) replacement (complete molecular substitution); (d) hydration and dehydration (addition to or release of H,O from the molecule or from solid solution); (e) oxidation and reduction (addition or release of 0, or hydroxyl ions); ( f ) polymerization and depolymerization (construction and breakup of hydrocarbon chains). Common authigenic minerals are indicated in Table 2-1. Inasmuch as any fortuitous mixture of minerals and ionic solutions, such as exists in the fresh sediment, is not likely to be in chemical equilibrium, a train of events is set in motion to establish such an equilibrium, at rates and in directions that are controlled by the environment.

BOUNDARY LIMITATIONS

There are limiting factors for diagenesis, but there are passage zones into metamorphism that vary according to the primary composition of the

23 sediment. Thus, the transition from limestone to marble takes place generally at lower temperatures and pressures than that from sandstone to quartzite. The limits of such passage zones may be defined basically in terms of chemistry and physics. Geochemical parameters In fresh, particulate sediment one deals with two components, namely, the solid sediment particles and the enclosing liquid. The latter is always present initially in marine deposits, but sometimes is absent at first from terrestrial sediments such as dune sands. As pointed out by Goldschmidt (1954), the chief controlling chemical factors in sedimentary petrogenesis are: (a) hydrogen ion potential (pH); (b) oxidation-reduction potential (Eh); (c) ionic adsorption phenomena. A world-wide study by Baas Becking et al. (1960) has shown that there seems to be virtually no environment found anywhere at or near the earth’s surface where the pH-Eh conditions are unacceptable for some form of organic life (Fig. 2-1). As a corollary, one must conclude that there is no environment near the earth’s surface (other than volcanic) that is not in some way modified by organic metabolic processes. Inasmuch as CO, is the principal by-product of organic oxidation and is also the principal raw material of plant and bacterial photosynthesis, it is to be expected that it plays an all-pervading role. Thus, C 0 2 reactivity on the earth’s crust will be related to the rate of organic metabolism. Inasmuch as the latter is thermophylic within the ecologic limits of the various phyla, provided that adequate water is present, the most reactive regions of the earth’s surface will be tropical. An advanced theoretical approach to the geochemistry of the early phase of diagenesis was published by Berner (1980). He derived a general “diagenetic equation”, with quantitative consideration of diffusion, compaction, pore-water flow, burial advection, bioturbation, adsorption, and radioactive decay. Special attention was given to decomposition of organic matter, cementation, and replacement. It is, therefore, unnecessary in this chapter to more than touch on a few environmental aspects. A good basic reference to solutions and equilibria is that of Garrels and Christ (1965). ( a ) Hydrogen ion potential. CO, dissolves freely in H,O, creating a bicarbonate ion and a free hydrogen ion. The hydrogen ion concentration in pure water at 20°C is lo-’ equiv./l (pH = 7), but upon saturation of water with CO, it rises to (pH = 5). CO, is thus involved with carbonic acid and the bicarbonate ion in sea

24 1,000

FI, Sp Oxidot ion

BOO

600

400

200

Eh mV 0

-200

-400

-600

\

I

I ACID

-800

I

0

pH

I

4

\

I

I

6

8

4*

%.\‘

I1 ALKALINE I

10

I

12

Fig. 2-1. Catenary diagram illustrating limits of natural environments in terms of pH and Eh, especially the sites of syn-, ana-, and epidiagenesis. (Based on works by L.G.M. Baas Becking and R.M.Garrels.) It is called a “catenary” diagram because the boundary is like a linked chain (“catena”).

water in equilibrium conditions as follows:

H 2 0 + C 0 2 = H , C 0 , = H C O ~ + H + = 2 H++CO;The corresponding equilibrium pH values in pure water (from left to.right) are 5, 6.3, and 10.3. In a closed system, the CO, reactions tend to move to the right as temperatures increase. Sea water in contact with COz will have a pH of 8.2-8.3. According to Sillen (1961), sea water is constantly buffered by the presence of clay-mineral particles in the ocean. Regardless of latitude and temperature, the oceans are thus always alkaline, whereas fresh waters tend to be acid.

25 Rainwater is normally saturated with CO, and, thus, has a pH of 5.5 or less, but it is usually buffered quite rapidly by soil carbonates except in acid situations. Rainwater p H is often much less (3-4) in the industrially polluted and volcanic areas, as well as downwind, for distances of up to 1000 km, where the “acid rain” is frequently toxic to lake and river life. Under organic control, however, very considerable modifications may be introduced, and the pH range in sediments may extend from about 2 to 12 (Baas Becking et al., 1960). It is also important to consider the ionic potentiul of the various components. The ionic potential is defined as the ratio of the ionic charge Z to the ionic radius r . According to Wickman (1944), these potentials fall into three categories : ( 1 ) Z / r = 0-3, soluble cations, which stay in true ionic solution even up to a very high pH, e.g., Na’, K + , M g 2 + , F e z + , M n 2 + , C a z t , S r 2 + , Ba2+; their hydroxides have ionic bonds and are, therefore, soluble. (2) Z / r = 3-12, elements of hydrolysates, which are precipitated by hydrolysis, e.g., A13+, Fe’+, S i 4 + , M n 4 + , etc.; these have hydroxyl bonds which makes them susceptible to hydrolyzation. ( 3 ) Z / r = > 12, soluble, complex ions, which form “complexes”, i.e., complex anions containing oxygen, and as a rule give true ionic solutions, e.g., B’+ , C 4 + , N 5 + , P 5 + , S 6 + ,Mn” ; they have hydrogen bonds, which also. like group ( l ) , lead to soluble compounds. These relationships have been clearly explained for geologists by Barth (1962). To quote from Barth (p.29): “Most natural waters go through an evolution of increasing pH, until they eventually empty into the sea, which is slightly alkaline. Silica becomes more soluble with increasing pH and is therefore often delivered into the sea. But aluminium hydroxide is precipitated in mildly acid solutions near the point of neutrality ... The difference in behavior of ferric and ferrous iron is of interest. Ferric iron is soluble only in rather strongly acid solutions; i t is, therefore, precipitated before aluminium, but the separation is usually not clean. Ferrous iron remains longer in solution in equilibrium with carbon dioxide in oxygen-free waters. Similarly tervalent and quadrivalent manganese ions are precipitated before bivalent manganese”. In the interplay between high and low pH in natural waters, the two principal players which rank (in total quantitative terms) far higher than all the other elements are silicon and calcium. The reasons for this will appear on considering the solubilities of the principal elements of the earth’s crust. After oxygen, which occurs mostly in combination, there are only seven quantitatively important elements: Si4+, A13+, Fe2+ or F e 3 + , C a 2 + , Na’ , K + , and M g 2 + , in that order (see Table 2-11). Inasmuch as natural waters are everywhere subjected to organic inter-

26 ference (largely reflected by the concentration of CO,), one may observe in cool. humid climates with acid soils (high CO, content and low pH) that calcium (with Al, Fe, etc.) is mobilized, but that silica, such as comprises quartz sand, remains stable and becomes progressively cleaner and cleaner (e.g., podzolization). In contrast, in a highly alkaline soil, characteristic of warm. rather dry “Mediterranean” climates (low in their supply of H,O and COz), the pH is high (8-9), so that calcium is precipitated, and results in the well-known lime “caliche” or Ycalcrete” crusts, whereas silica is mobilized and generally carried into the river system during the brief wet season, partly in colloidal form. In drier places, it may simply rise by capillarity to the surface, there to replace calcrete or to be reprecipitated on desiccation as a “silcrete” crust. Thus a specific geological formation, subjected through time to different paleoclimates, may be affected by an alteration of ground and artesian waters (both at the surface and at depth) from high to low pH, leading to complex intergrowths and respective replacements of quartz by calcite and vice versa (Runnels, 1969; Fairbridge, 1975; Fairbridge and Bourgeois, 1978). Extreme swings, from one absolute p H boundary to the other are to be

TABLE 2-11 Common elements in the earth’s crust and their solubility at 25°C with respect to pure H,O (pH=7), or modified by solution of CO, ( p H = 5 ) or C0:- ion ( p H = 9 ) Element

Crustal abundance At p H 5 (at 25°C) (moles/l) (parts/1000, or g/kg)

Si Al Fe3+ Ca Na

277

K Mg

81 50

36

281

2 . lop3* 1.4. lo-’ 6 . lop9 very soluble

At p H 7 (moles/l)

At pH 9 (moles/l)

4.5. 1.4. l o - ”

6 . lo-’

6. at p H 12

1.4, 6 . lo-*’

3.2. lo-’

pH at which hydroxide begins to precipitate

**

2 4

2.5 12

very soluble (the hydroxide will not precipitate)

26

21

very soluble

1.1.10-‘

10.5

* Approximation, based on curve by Correns (1949, p. 210). Somewhat different according to Siever (1959). SiO, is in the form of the oxide, not hydroxide.

** This value may be too low. Due to the amphoteric nature of AI(OH),, i t begins to dissolve i n strongly alkaline solutions, forming the complex AI(OH,)+OH- =AI(OH), . At pH = 10. the solubility of AI(OH), increases sharply. Equally rapid (in the other direction) is the increase in solubility at ca. pH=4.

27 seen under exceptional conditions in some deserts. Normally, the dilution of all solutions by rain water (pure H 2 0 ) tends to bring the pH within one unit of neutrality (pH = 7), and thus minerals that may develop during diagenesis in the extreme pH ranges are exceedingly rare. ( b ) Oxidation-reduction (“redox ”) potentials. In order to understand the boundary limits of pH it is necessary to consider also the oxidation-reduction or “redox” potential (Eh), which is to some extent reciprocal to the pH, but is influenced by certain other factors (Latimer, 1952: Blatt et al., 1980. p. 240). The pH-Eh relationship for natural environments, as established by Baas Becking et al. (1960), presents a boundary like a distorted shield, with small shoulders (or “ears”) in the low pH-high Eh corner and in the high pH-moderate Eh corner (see Fig. 2-1). The whole shield (i.e., all natural environments) fits between two parallel lines, the slope of which corresponds to -0.059 V/pH unit. These two parallel lines are absolute barriers (“fences”), representing the equilibrium limits of water at or near the earth’s surface. The upper diagonal bounds the upper limit ( H 2 0 / 0 , ) where the partial pressure of oxygen is equal to 1 atm, whereas the lower diagonal marks the lower limit ( H 2 0 / H 2 ) , where p H ,= 1. In the lower diagonal. the left-hand end corresponds to pH of 0.0 a n d E h of 0.0. (See also Fig. 2-2) The value of these relationships to an understanding of diagenetic reactions should hardly need emphasizing. Yet as brought out by Garrels (1960. p. 104), it has taken geochemists nearly half a century to recognize this fully. It is fortunate for the sedimentologist that a reliable electric (battery or line-operated) pH-Eh meter can be purchased for a quite modest outlay. and may then be freely used not only in the laboratory but in the field, permitting readings to be obtained on the spot in natural media. This is important, for it is not easy to obtain a sample of some gas-saturated mud, for example, and transport it to a laboratory without grossly upsetting the original Eh-pH relationships. To the sedimentologist, the pH-meter is what the field pick or hammer is to the hard rock geologist. In short, i t is absolutely basic to sedimentology. Whereas the pH readings are easily obtained and are closely repeatable, the Eh is more difficult to establish in natural media and must usually be regarded only as an approximation. It was through soil studies and bacteriology that pH-Eh relationships were introduced into sedimentology. ZoBell’s study ( 1946) brought to this writer his first inkling of their application. Earlier work has been done in France and Belgium, and an English translation of a book by Pourbaix (1949) presented the thermodynamics of dilute aqueous solutions in terms of pH and Eh. These principles have been excellently applied to geology by Garrels ( 1960).

28 PH

-

7.0 I

-0.3

Fig. 2-2. “Fence diagram” illustrating principal environments of sedimentation and diagenesis, according to Eh and pH. (After Krurnbein and Garrels, 1952.)

In nature, oxygen-consuming organisms are the principal agents in lowering the redox potential, but in the atmosphere or at the sea floor there is usually such constant water circulation that the lower half of Baas Becking’s (1959) “shield” is not involved. This state of affairs changes, however, as the sediment is buried; in clays, for example, a sediment thickness of only a few millimeters is sufficient for diagenesis to begin in earnest. The aerobic bacterial attack on buried organic debris quickly removes all free oxygen from interstitial water, and at a p H of about 7, the Eh is about -0.4. The anaerobic bacteria then take over, as their aerobic brethren have literally eaten themselves to death, and they attack the sulfate anion, the most readily divisible ion containing oxygen. After chloride, SO:- is the most important

29 anion in the ocean (7.68% of the total ions). With reduction to sulfite and then to sulfide, the redox potential steadily drops and the pH shifts to higher values, so that in young sediments at a depth of 1-3 m the pH is often up to 9 or more. Any free CO, has long since passed into CaCO,, so that the only gas phase is H,S. The setting is now appropriate for the pyrite reaction, which is perhaps the most significant in all of diagenesis (see the section on Oxidation and Reduction). The stability fields of Fe,O,, Fe,O,, and FeS, have been illustrated by Garrels (1960, p. 145). Under favourable conditions, pyrrhotite, galena and other metallic sulfides will start to form. Baas Becking et al. (1960) have demonstrated the stability fields of some important naturally occurring non-metallic compounds in laboratory-controlled bacterial studies (Fig. 2-3). ( C ) Ionic adsorption phenomena. According to Barth (1962, p. 30), ionic

adsorption phenomena.. . "take place at low temperature in colloidal phases or phase complexes that are capable of capturing and binding certain ions through adsorption. One example is the binding of potassium ions by the clayey products of the hydrolysis. In a geochemical adsorption process the binding of the ion to the colloidal surface takes place in competition with the

Fig. 2-3. Stability fields of some important naturally occurring non-metallic compounds in terms of oxidation-reduction potentials and pH framed within the limits suggested by this work. Distribution of these compounds in un-ionized states is governed by their dissociation constants. (After Baas Becking et al., 1960.)

30 over-all hydration of the ion in the solvent. It can be shown that the degree of adsorption in an ion is a function of radius, charge, polarizability, and normal potential, as well as the nature of the chemical compounds formed at the phase boundary. Through the processes of adsorption the natural waters are deprived of many of the rarer elements. Most of the ions of the heavy metals, such as ions of lead, zinc, and copper, as well as complex ions of arsenic and molybdenum, are captured by, and coprecipitated with, the colloidal particles, usually hydrolyzates, and thus are supplied to the sediments. The amounts of poisonous metals and metalloids which potentially have been delivered into the ocean from the primary rocks throughout geologic times are so considerable that a serious poisoning of the ocean would have been caused if this process of elimination of poisonous substances had not been in action. Or the evolution of life would have taken a different course, developing organisms not susceptible to our poisonous metals. This statement applies, for instance, to copper, lead, arsenic, selenium, mercury, antimony, and bismuth. In many cases these metals have been removed from aqueous solutions by a means also known in practical medicine, that is, adsorption on freshly precipitated hydroxides of iron. There is considerable concentration of selenium, arsenic, and lead in the sedimentary iron ores. The arsenic content of these ores in most cases is so high that it brings a very notable quantity of arsenic even into iron or steel, from which this element is difficult to eliminate by the usual technical processes of refining. Molybdenum is concentrated in manganiferous sedimentary ore deposits.” Geophysical parameters

The physical boundaries to diagenesis are defined mainly in terms of temperature, pressure and time. ( a ) Temperature. The mean temperature at the earth’s surface through most of geological time for which there are identifiable indicators, say the last 3 . lo9 years (Fairbridge, 1967a), has remained at ca. 20 10°C. Soil temperatures today at depths of 0.1-1.0m fall generally within this range. Seasonal, latitudinal and altitudinal variables increase this range from ca. 60°C down to - 100°C, excluding volcanic phenomena, hot springs, and so forth. Certain algae are adapted to life in hot springs near 100°C, but this is quite exceptional. The mean temperature at the water-sediment interface over most of the deep ocean floor is ca. 2°C. At intermediate to shallow ocean depths, the temperature approaches the world mean, noted above. These very moderate to low temperatures greatly influence the geochemistry of early diagenesis, because in such ranges, crystallization is normally

31 slow and only simple compounds form. The complex, mixed crystals commonly involved at the temperatures of formation of metamorphic and igneous rocks (over 100°C) are encountered only in deep-burial diagenesis (“metagenesis” of Soviet geologists; “anadiagenesis” of the writer). Solubility and pH both change with rising temperature (see Fig. 2-4). The complex mixed lattice of the magnesian-calcite-dolomite series is a notable exception to this generalization, but even this reaction is favored in nature by elevated temperatures (30-40°C). In evaporite basins a complex series of halide salts is also favored by somewhat higher temperatures (Braitsch, 1962; Borchert and Muir, 1964). In geosynclines, that is to say, sedimentary basins or troughs marked by a considerable accumulation of sediments and often extending over periods of the order of 10’ years, temperatures below the surface are found to rise by ca. 0.5 to 1.5”C/100m, due to the poor conduction of the earth’s internal heat. In certain regions, however, the gradient may be much steeper. There are generally two potential causes of this: (1) Abnormal concentrations of igneous or radioactive heat, as near volcanic vents and major fault lines, notably the celebrated “Mid-Ocean

5 0

1

20

1

1

1

40 60

80

I

5000

1

1

1

1

1

1

1

1

100 120 140 160 180 200 220 24(

I

I

I

I 1

IOQOO 15,000 25QOO 20,000 30,000

Feet

Fig. 2-4. Variation in the dissociation constant of water as a function of temperature and depth in a geosyncline. (After Blatt et al., 1972, fig. 7-1). One should note that neutrality varies with temperature, so that at an average burial depth of l0,000m, neutrality (equal concentration of H + and OH- ions) is indicated by pH equal to 6 .

32 Rift” along which the heat flow may rise to 8 . cal. cm-2 sec-’, in contrast with an average of ca. 1 elsewhere. (2) Sedimentary accumulations of minerals, which on oxidizing are exothermic, that is to say, that they generate heat. The oxygen is generally brought in by artesian water and the heat produced is dissipated by its continued circulation. This water, which may be partly connate (that is primary), tends to migrate upwards and outwards as basin sediments compact. One of the principal minerals involved is either marcasite or pyrite (FeS,) and these are commonly present as vast quantities of finely divided particles in any shale that was formed under slightly reducing conditions, or in coal seams. The heat generated may spark off a spontaneous coal fire in mines or landslide areas. In artesian wells at Perth, Western Australia, the water from a depth of only 300m comes from a pyritic shale-silt-sand sequence that brings the water temperature at the well head to 90°F (32°C). In deep oil wells on the Texas Gulf coast at depths down to 25,000 ft (8000m), the temperature through normal heat flow should be high, but may exceed 150”C mainly through this same exothermic mineral oxidation. ( b ) Pressure. In a gradually accumulating sedimentary basin, there is pro-

gressive increase of load pressure (Lane, 1922). This has sometimes been called load metamorphism or static metamorphism (Grabau, 1913, p. 750), as opposed to dynamic metamorphism, which involves tangential stresses as well as simple vertical compressive stress. For this reason, and because temperatures are relatively low, the effect of simple overburden pressure may be considered as “diagenetic”. There is a type of load metamorphism, however, that transcends normal diagenetic changes, because of complete remobilization of ions and formation of minerals beyond even the limits of metasomatism. The results of this are seen in some ancient evaporites, in particular the Stassfurt deposits of the Permian in north Germany (Janecke, 1915; Rinne, 1920; Braitsch, 1962; Borchert and Muir, 1964). The question of the dominance of the role of load or of dynamic metamorphism, or of geothermal heat is not yet resolved, but the temperature of the alteration was probably not over 80°C. ( c ) Time. Time is of course the geologist’s trump card in any argument with physicists and chemists. Some solubilities are so extraordinarily low that they take millions of years to bring about any noticeable effect. I t is, however, rather too easy to delude oneself by this line of reasoning, and it is worth bearing in mind that certain seasonal effects are highly episodic, and lead to short peaks of hydrolysis, pH-Eh oscillations, etc., which may pass unnoticed at other times.

33 There are also brief (in geological terms) episodes of diastrophic activity, such as periods of uplift accompanied by massive fracturing, jointing, and faulting, that would favor extensive recirculation of waters which may previously have lain stagnant (or isolated by low permeability) for extended periods of time. Oil geologists were well aware of this characteristic in the history of a basin’s fluid components to be periodically subject to induced flow and interruption. To generalize, one may say that the average geosynclinal basin experiences progressive downwarping and compaction for lO7-1OXyears, and that this is followed by one or more episodes of uplift (with fracturing and faulting), generally in brief spasms, marked by earthquakes of a few hours. Depending on geotectonic factors and the nature of the underlying crust, the geosyncline may or may not become involved in superficial folding or in deep-seated buckling and compression. In the latter event, the sediments are placed in regions of high heat flow, and then metamorphism, granitization and igneous activity are introduced. Where only superficial folding is involved, however. in a superficial “skin” that may not exceed 3000m, the principal orogenic stress is provided by gravity, and rock alteration is limited to diagenesis except in specific stress-strain zones such as faults. After uplift, extended periods are likely to pass with only episodic and very gentle (epeirogenic) revival of topographic relief. Long-continued exposure to meteoric circulation will be the rule, and some non-metamorphic Precambrian rocks have been so exposed for periods of over 1 . 10’ years.

DIAGENETIC EVOLUTION

Diagenesis begins at the moment a sedimentary particle comes to rest, for example, on the sea floor. It continues to a point in history when either deep burial and orogenic buckling cause the initiation of metamorphism, or when emergence leads to exposure and the initiation of weathering and erosion. Sever (1959) has called these phases, respectively, early, middle, and late diagenesis. It is an almost Davisian evolutionary cycle of youth, maturity, and old age, except that it may be complicated by rhythmic repetitions and “accidental” alternative courses, introduced by interaction with geotectonic, paleoclimatic and other cycles. For clarity, one may employ with “diagenesis” the classical prefixes “syn-” (together with, or synchronous with the sedimentation process), “ana-” (up, again, or thoroughly lithified), and “epi-” (outer or surficial, i.e., modified by surface phenomena). Thus, the three phases have been named by the writer (Fairbridge, 1967b, p. 32) as follows (Fig. 2-5): (a) syndiagenesis (the sedimentation phase, depositional and early burial);

34

RAINFALL

pH 7

+ + + + + + + + ’ + + +

Fig. 2-5. Idealized profile through a continental margin, showing the sites of contemporary marine sedimentation and the three phases of diagenesis. There is: ( 1 ) diffusion potential during syndiagenesis; (2) upward liquid motion in anadiagenesis (compaction fluids); and (3) downward motion during epidiagenesis.

(b) anadiagenesis (the compaction-maturation phase, with deep burial); and (c) epidiagenesis (the emergent-pre-erosion phase). This is a temporal and spatial classification designed to establish the environments of the mechanical and geochemical processes, i.e., time, place, and process. The three phases are evolutionary and dynamic, passing from one to the next across recognizable boundaries. Within each phase there are distinctive hydrologic systems associated with characteristic, though not unique geochemical systems. Thus, for example, lithification is the most common attribute of anadiagenesis, but in some specific environments (e.g., several types of carbonate sites), lithification is syndiagenetic. Dapples ( 1962) designated three geochemical stages, specifically for sandstones, and adopted for that purpose by several workers (e.g., Selley, 1976), as follows: ( 1) Redoxomorphic (mainly oxidation or reduction reactions; a metabolic control, that is most effective in the syndiagenetic phase, which includes both initial and early burial sub-phases). (2) Locomorphic (principally metasomatic, one mineral being replaced by another; this is important during the lithification of the anadiagenetic phase). ( 3 ) Phyllomorphic (characterized by ion exchange such as is associated with clays and micas, to be observed during all phases).

35 The time factor is subject to great variation. Rapid subsidence of the basin can lead to accelerated transition to anadiagenesis. Slow subsidence may expose the fresh sediment to eustatic fluctuations and thus to epidiagenesis without going through the anadiagenetic stage. Choquette and Pray (1970, p. 219), on a basis of time and porosity types, also recognized three phases: eogenetic stage (time of brief evolution in proximity to sea floor), mesogenetic stage (long time period following burial and isolated from sea floor and preceding emergence), and tefogenetic stuge (time following emergence and influenced by processes leading to an unconformity). Essentially these terms, differently expressed, are synonymous with syn-, ana- and epidiagenetic, within each of which there is a trend towards a steady-state or equilibrium condition. In familiar terms one could simply restate the syn-, ana- and epidiagenetic phases as “mud”, “hard rock” and “weathering” stages. In many cases this would be a true designation, but it would also be misleading in some instances: hard, concretionary segregations may form during early burial stage, leaching often occurs during deep burial, and ha,rd-crust formation frequently results from exposure to weathering processes. Emphasis should include not only time and place, but also process. From a quite different point of view, the diagenetic processes involved in the evolution of sand to sandstone include six stages according to Pettijohn et al. (1973, p.387): five stages correspond to progressive depth of burial, 0-nearly 10,000m, with a pressure increase of about 1 bar/4 m; the sixth stage follows uplift and is equivalent to epidiagenesis (Fig. 2-6). Inasmuch as it is the anadiagenetic stage that is most likely to be impinged upon by the geotectonic cycles, this episode may be so reduced in the evolution to the epidiagenetic phase that.one might almost speak of a short-circuit connection when orogenesis has caused uplift of fresh unlithified sediments and led to their rapid erosion. Alternatively, in the same stage but in a different geographic position in the sedimentary basin similar sediments might be trapped in a downbuckling of the crust and initiated into a metamorphic cycle that might indefinitely postpone or eliminate the epidiagenetic phase. Syndiagenesis Defined as the sedimentational, pre-diastrophic phase, syndiagenesis begins at the moment the sedimentary grain touches the bottom, and is marked by the presence of large amounts of trapped interstitial, or connate, water, which is expelled only very slowly. The term “syndiagenesis” was first used by Bissell (1959). In the study of mineral deposits the analogous process is “syngenesis”. This term, as originally defined by Fersman ( 1922), however,

Fig. 2-6. The stages of diagenesis of sandstones in relation to depth of burial and increase of pressure and temperature. (After Pettijohn et al.. 1972. fig. 10-1, p. 387; courtesy of Springer, New York.) Stages 1 and -7 are syndiagenetic. stages 3 to 5 are anadiagenetic, whereas stage 6 is epidiagenetic.

was intended only for primary chemical sediment such as oolite, and he used “diagenesis” only as the writer uses “syndiagenesis”. Two stages are recognized in the syndiagenetic phase. These have been called (Dapples, 1959, 1962): initial stage, controlled by the chemistry of the superjacent water, and early burial stage, controlled by the entrapped, connate water, chemically modified by the bacteria and other subsurface organisms.

( a ) Initial stage. Buried with the sediment is generally a moderate to large amount of organic matter which provides nutrients for burrowing organisms that greatly disturb the surface layers of the fresh sediment and keep them relatively well oxidized, as earthworms do in soil. Indeed, Waksman (1933) described the organic matter in this zone of bioturbation as marine humus. Bader ( 1954) described pelecypod population densities as essentially controlled by what he called the “decomposition coefficient” of the sediment. In basins lacking a free circulation above the sediment-water interface, a stagnant, euxinic (“Black Sea”) condition will lead to poisoning of bottom waters, and metazoic benthos will be excluded, resulting in the nice preservation of original finely stratified bedding planes (Chilingar, 1956a; Caspers, 1957). This is typical of the Black Sea, where the trapping of organic debris

37 leads to diagenesis of petroleum hydrocarbons, but this is more a function of rapid accumulation than it is of euxinic conditions (Smirnow, 1958; Degens, 1967; Degens and Ross, 1974). The bacterial population near the sediment surface will, in the well-ventilated basins, belong to the aerobic families and some may even be photosynthetic autotrophs; others will employ the buried organic matter and the oxygen from the connate water. The result will be a sharp rise in the pco, marked by a drop in pH, which from the surface may pass from 8 to 7 or 6.5 (ZoBell, 1942; Debyser, 1952). This zone extends for a few millimeters down to about 30-50 cm, depending on factors such as depth of water, amount of organic nutrients, rate of sedimentation, etc. (Twenhofel, 1942). The effect on the inorganic sediments is sometimes slight: “cleaning” of quartz sands, for example; or it may be profound: rapid solution of calcite and particularly aragonite grains, destruction of carbonate shelled foraminifera and calcareous spicules, etching of more massive shells and “weathering” of some feldspars and clays. Pelecypods (and other organisms) that inhabit such ,acid bottoms protect themselves with a chitinous covering (periostracum), e.g., Mytilus; but immediately after death solution begins. Experiments by Hecht (1933) demonstrated that even on the surface, empty shells lost 10-20% of their weight per year (in ordinary North Sea water). In the richly organic muddy sediments of the Wadden Sea, however, only casts and moulds of shells are normally found. Experiments of burying shells along with the rotting molluscan remains showed losses of up to 25% of the total shell weight in only two weeks. Gypsum crystals sometimes formed on the meat, illustrating the local reduction of marine SO:- to H,S and its immediate reaction with the Ca2+ of sea water to form CaSO, a2 H,O. On tropical coasts, even on many coral reefs, the almost universal presence of mangrove swamps (populated especially by the genus Rhizophora) provides a rich source of organic debris, leaves, branches, etc., so that the pH in the muds (even at the surface) normally drops to 6.5 or less (Orr and Moorhouse, 1933). In addition to C 0 2 , it is probable that humic and tannic acids are also liberated. Coral reefs of CaCO, are pocked by giant mud-filled pot-holes up to 5 m in diameter, wherever mangrove trees have been situated. ( b ) Early burial stage. Below the oxidizing zone is a reducing zone (Dapples, 1959, “early burial stage”). Here anaerobic bacteria become dominant and the pH rises steadily, often to above 9 (ZoBell, 1942). The Eh drops to -0.4 or - 0.6. Sulfate-reducing bacteria, notably Desuljovibrio desuljuricuns, liberate H,S. CaCO, precipitates freely at a pH of 8.5 and in this Eh range FeS is the stable iron compound (afterwards becoming FeS,; see section Oxidation and Reduction). Somewhat less commonly, siderite (FeCO,) is

38 formed (see stability diagram in Garrels, 1960, p. 130). During diagenesis, the chemical reactions are generally governed by the first part of the van’t Hoff‘s Law, which states that low-temperature reactions usually generate heat (i.e., they are exothermal), and are accompanied by the association of ions. Van Hise (1898) observed that the operation of this law was characteristic of his “upper physicochemical zone”. In the bacteria-rich reducing conditions of the early burial stage, however, many of the larger organic and inorganic molecules are broken down. Vegetable matter disintegrates and only the most stable parts remain; lignin, the principal residue, is extremely stable in the marine realm and may be a useful indicator for rates of sedimentation (Bader, 1956). An important aspect of the break-down of organic matter is the rapidity of the reaction in the aerobic stage. Material that survives this attack and passes into the “early burial” (anaerobic) condition has a much greater chance of preservation in rocks (Abelson, 1959, p. 83), though often further modified to petroleum hydrocarbons and other organic products (Hunt, 1979). The transgressive phases of major eustatic fluctuations are particularly significant (Tissot, 1979; Arthur and Schlanger, 1979). After hydrogen, the most important active element in the early burial stage is sulfur. This element is present in sea water as the anion SO: - , which represents 7.68% of the total dissolved constituents of the ocean, and is the most important after Na+ and C 1 ~ “The . sediment acts as a chemically open system to the sulfate of the overlying water” (Berner, 1964). Sulfur is also an important member of many organic compounds. It shows a valence change of -2 to t 6 during oxidation and reduction. It also has two stable isotopes 32S and 34S, with a 6% mass differential which is easily measured with modern instruments. During the valence changes, for example, from SO,‘- to S’- , the isotopes are fractionated, so that the sulfide ion is enriched in j’S, the more energetic isotope. The sulfur isotope ratio is, therefore, a valuable indicator of passage through the early burial stage (Holland, 1973; Holser, 1977). Thus, H,S and related authigenic minerals formed then show an 32S/34S ratio of 22.1 to 22.7 (with an everage of 22.49 for all sedimentary sulfides), in contrast to a constant 21.76 0.02 for sea-water sulfates (Ault, 1959); evaporate sulfate figures are similar to sea water, but have a wider spread (f0.2). The mean sulfur isotope ratio for magmatic hydrothermal and meteoritic sulfides is about 22.2, which is readily distinguishable from the mean for sedimentary rocks; but unfortunately the spread of values for the sedimentary rocks makes it difficult to use this device to solve the controversy about the metallic sulfide ore deposits. Baas Becking et al. (1960) have shown experimentally that, under certain conditions, marine bacteria can synthesize not only pyrite, but also the common ore sulfides. Deans ( 1950) reported

39 that Westoll had found fossil fish skeletons diagenetically replaced by galena, sphalerite, chalcopyrite and bornite. There is a considerable controversy, therefore, between those who would attribute all ore sulfides to magmatic sources and those who consider them syndiagenetic. Both sources are possible as demonstrated experimentally. The main problem today is to discover the relative importance of the various sources in the different deposits. While synsedimentary origins are now widely accepted, the localization of very high metal concentrations in sea water, from time to time and in rather limited areas, was probably due to the local thermal springs (“exhalative” magmatic hydrothermal sources) so that ultimately magmatic sources were responsible (Dunham, 1952: Williams, 1960). The principal sources of these thermal springs are sea-floor spreading centers, such as demonstrated near the Galapagos Islands in the Pacific (Spiess et al., 1980), or in the currently inactive trench of the Red Sea (Degens and Ross, 1969). The element nitrogen is sometimes forgotten in geological literature, but its role (mainly through ammonia compounds) in the syndiagenetic phase is not unimportant. It can also be a helpful indicator. Whereas the total carbon content decreases sharply with depth of burial, the level of fixed ammonium remains rather constant. Thus, the C/N ratio can be used for environmental reconstruction (Arrhenius, 1950; Stevenson, 1960). The reducing zone generally leaves a characteristic mark on the syndiagenetic phase, because it is inevitably the last environment of a sedimentational stage and thus leaves its imprint on sediments for all time. Although its products may subsequently be modified, the evidence is never totally effaced. Indeed the oxidation stage (“initial stage” of Dapples, 1959) may be bypassed in the euxinic environments and the reducing stage would occur at and above the sediment’s surface (Degens and Ross, 1974). If the oxidizing zone had been present, however, the acidizing experience of the sediments is the one that would have had the more striking effects as seen in the light of day, perhaps a hundred million years later; carbonate fossils are absent and the only obvious traces of former life are the chitinous forms such as conodonts. If the sediment is a coarse-grained one, such as a quartzose sand or silt, there is generally a far greater opportunity for oxidation than in clays. Thus, entrapped organic matter is totally consumed, the soluble carbonates and other minerals are destroyed or modified under low pH, and the sand is thoroughly cleaned. In this case the reduction zone is left with little nutrient for the bacteria and the populations are thus greatly limited in size and variety. In the case of carbonate sediments (that is, where the great bulk of the

40 material, regardless of grain size, is CaCO,) no amount of bacterial CO, production will cause the total solution of their substrate. The response of lime sands p proto-calcarenites") will be rather similar to that of quartz sands, but lime muds p proto-calcilutites") will respond rather like clays, and the resultant limestone may be speckled with marcasite or pyrite concretions. These are quite rare in calcarenites. In the case of fine-grained siliceous material, it is suspected that much of it enters the ocean through rivers in the form of dissolved silica (Bien et al., 1958; Heath, 1974), as desert dust, or by solution from volcanic ash. It is biogenically concentrated as opaline silica from diatoms, radiolaria, sponge spicules, holothuria, and alcyonaria (Riedel, 1959). Penecon temporaneous resolution may occur, but is prevented in rapidly accumulating Globigerina oozes. The silica appears to accumulate in small globules of gel that dehydrate very slowly. Migration of silica often seems to occur while the sediments are still quite soft. Indeed much movement may be expected while sediments are permeable and electromagnetic response is facilitated. It may be borne in mind that the low pH of the initial stage, which may lead to carbonate solution, will favor SiO, stabilization, whereas under the high pH of deeper levels the reverse is true (see recent papers collected in Van der Lingen, 1977). Horizons of flint nodules and chert layers in chalks and limestones are often so regularly displayed that one might take them to reflect a sedimentary rhythm. On the other hand, Sujkowski (1958, p.275) speaks of a diagenetic rhythm, while admitting that a mild sedimentary rhythm might lead ,to a very inconspicuous banding of textural character. On these terms migration would be favored by slightly coarser more permeable layers; these would be predisposed to diagenetic rhythmic bedding. Additional discontinuities are also provided by “hardgrounds” which are often very widespread (Bathurst, 1975; Kennedy and Garrison, 1975). The effect of diagenetic rhythm probably goes further than the formation of concretions. According to Sujkowski (1958): “By separating the compounds of an unstable mixture inside a sedimentary series, diagenesis exaggerated the rhythmic differences pre-existing in a deposit. It is also not excluded that in some texturally homogeneous deposits, diagenetic rhythm is quite a secondary phenomenon resulting only from the unmixing of the different chemical components to the limits of diffusion”. Thus it is evident that primary deposits may be more uniform in composition than the rocks derived from them. The thickness and duration of the syndiagenetic phase are determined by a number of variables such as lithology, organic components, rate of sedimentation, aeration, and depth of water. Generalizing, one may say that the base of the syndiagenetic phase is defined by the lower limit of vigorous

41 bacterial activity, which may range from a depth of ca. 1-100 m. In terms of organic metabolism, syndiagenesis may be taken to last as long as the food holds out. In terms of absolute duration, this may be for 1000-100,000 years, but considerably more research is needed on this aspect. It has been claimed that viable bacteria can be traced back to Carboniferous coal seams (nearly 3 lo6 years old), but possibilities of contamination are so great that it is very difficult to prove. +

A nadiagenesis

Anadiagenesis corresponds to deep burial, involving the compaction and maturation phase of diagenesis, during which the particulate sediment grains (or chemical ions) become once again (Greek: m a - ) lithified. Diastrophism may or may not be involved: this depends upon the particular geotectonic situation of the sedimentary trough or basin. Characteristically this phase is one of slow compaction and concomitant expulsion of connate water. Rising mineralizing waters are often known as hypogene (with,hydrothermal admixture), but it should be emphasized that most anadiagenetic waters are non-magmatic (White, 1957). Dapples (1959) called this the “late stage” of diagenesis, but apparently included with it also the epidiagenetic phase (see section on Epidiagenesis below). During anadiagenesis some of the connate water becomes trapped permanently in the sediments as a result of compaction and cementation to the point of impermeability. It thus becomes “fossil sea water”, though greatly modified from its original form. The name “connate” was proposed by Lane (1909), and indeed only since the introduction of the term has the importance of this phase of diagenesis been appreciated (see Lane, 1927; White, 1957; and Chave, 1960). The effect and weight of sediment-loading was also studied by Lane (1922). Earlier, Van Hise (1904), for example, regarded all interstitial water as meteoric. Some of the economic geologists, in contrast, seemed to have regarded it as almost all magmatic (Schmitt, 1950). Sediments that have passed through anadiagenesis are, therefore, characterized by Cementation, the most common cements being siliceous or calcitic, and more rarely ferruginous (Pettijohn et al., 1975; Bathurst, 1975; Blatt et al., 1980). An important “diagenetic fabric” may be studied on polished surfaces or in thin sections. Use of the universal stage microscope permits the identification of the rarer minerals, particularly the sequence of growth (Glover, 1963). Etching and overgrowth phenomena are most common. In siliceous sediments this phase may not become well developed until considerable depths are reached, but with carbonates the reactions may be extremely rapid, begmning even at the surface. On mature plate boundaries, with a loading of up to = Z O O bars, i.e., in

42

orthogeosynclines, it is possible that the progressive downwarp and filling of the trough leads to burial in excess of 10,00Om, which is the approximate depth at which the geoisothermal level exceeds the normal operative limits of diagenesis (about 100°C). The system is affected firstly by “load metamorphism”, and secondly, by diastrophism which inevitably tends to take place in any segment of the earth’s surface which is depressed by 10-20 km. Gravitational sliding and crumpling probably occur on a geotectonic scale, directed at first inwards, while the basin continues to subside. This may or may not be followed by vertical readjustment if subduction and excessive heating leads to granitization of the roots zone. Uplift then results, and further gravitational slides occur, this time externally directed. Through this orogenic evolution, it is evident that, taken as a whole, the sediments involved in the lower part of the trough, and those that slide into it, will become incorporated in the metamorphosed or granitized roots. The superficial sediments, however, are only involved in the “Jura-type” displacement and “Alpino-type” nappe slides towards the exterior. These generally escape metamorphism, and should be associated only with lower temperature fields and more modest dynamic stresses. Ensuing alteration is thus little more than anadiagenetic. Theoretically it might seem possible to distinguish between the strictly compactional and the dynamic phases in such orthogeosynclines, but in practice it is difficult to draw a sharp line, especially inasmuch as much of the sliding and folding is synchronous with the sedimentary accumulation. On the other hand, in parageosynclines (cratonic basins, i.e., one of Kay’s, 1951, auto-, paralia-, exo-, zeugo-, taphro- or epieu-geosynclines) one is dealing with a basin that has a rather stable underlying crust and, therefore, deep burial is impossible. These basins, on the basis of a world survey (Fairbridge, 1959) cover 32% of the continents and shelves, with a maximum depth averaging 5 100 m and an average area size of 180,000 km2. It is evident that with such limitations anadiagenesis in parageosynclines is unlikely to pass down into metamorphism, except perhaps in localized zones of intense faulting and volcanism. It is one of the characteristics of the parageosynclines, however, that they accumulate episodically, that is to say, there is a phase of subsidence, faulting, and downwarping, followed by a period of stability or brief uplift, which in turn is succeeded by renewed subsidence. Thus, a number of well-known basins have a two-, three- or four-storied structure, each showing progressively advanced anadiagenesis with depth. Emergence is marked by epidiagenesis and surface weathering, only to be succeeded by an unconformable sequence with its own new diagenetic cycle. With each renewed subsidence there will be revived fracturing and jointing. The lower stories thus display multiple generations of joints; this may be called diaclastic revival (from “diaclase”, the classical term for joint or fracture).

43 Geochemically, anadiagenesis is often a de-watering stage. Connate waters are progressively expelled from the lower levels, moving upwards and outwards, and following the dip of the basin. Gradually pores close, permeability is reduced, and the basin becomes more or less sealed. So it remains until re-activated by diaclastic revival, which may be due to diastrophic motion, ranging from further subsidence to general epeirogenic uplifts or orogeny. As waters become progressively displaced, they must pass through overlying or lateral strata, in general following the predictions of hydrodynamic theory (Scheidegger, 1957). They are thus subjected to mixing with other generations (and thus potentially differing classes) of connate waters as well as the varied mineral components, some of which may still be metastable. Nagy (1960) has spoken of a “natural chromatography” and others of “clay filtering” (Millot, 1970), inasmuch as it is essentially the clays that offer ionic adsorption potentials (see further discussion below). The deep waters may become progressively but irregularly more saline (Chave, 1960). Valuable compression experiments have been reported by Chilingarian et al. ( 1973).Whereas syndiagenesis is mainly characterized by initially acid waters followed by reduction, anadiagenesis is marked by increasing alkalinity but more neutral redox potential. Because in many sedimentary sections pervious sands alternate with less pervious shales, a progressive upward mixing of expressed waters may lead to curious anomalies in the pressures and salinities in deep basins. Laboratory studies (Berry, 1960) suggest that the clays while being mainly responsible for the geochemical filtering or natural chromatography act also as semi-permeable membranes subject to the law of osmosis. Sandstone porosity may also be progressively reduced by upward transport and deposition of clay minerals (Almon et al., 1976; Blatt et al., 1980). The depth-time limits of anadiagenesis may be broadly defined as extending from the lower limit of syndiagenesis (ca. 1-100 m) to ca. 10,000 m, and from say 1000-100,000 years to somewhere between lo7 and 10’ years. In orogenic belts, however, it may be short-circuited and grossly curtailed by rapid elevation. Numbers of indicators may be used to judge the time range of diaclasis and other phenomena. For example, the dehydration of silica gels seems to be an extraordinarily slow affair, and what may start as an opaline silica, often ends as chalcedony (the cryptocrystalline dehydrated form) or quartz. Evidence is given in the section on Hydration-Dehydration to suggest that the anadiagenesis of primary silica may occupy 104-10h years. Epidiagenesis Epidiagenesis is the emergent or post-diastrophic phase of diagenesis. The analogous stage in mineral genesis is associated with the terms epigene or

44 supergene. Edwards and Baker ( 1951) described pyrite and marcasite nodules which formed in the syndiagenetic phase as “supergene”, but this is not appropriate, because they originated in connate water, whereas supergene refers to descending (meteoric) water. By the time of the onset of the epidiagenetic phase the sediment has been successively exposed to penecontemporaneous environments, to compactional processes, and now finally to subaerial controls. In the preceding section it was noted how a brief negative eustatic oscillation could shortcircuit the anadiagenetic phases and how the sediment could be exposed immediately after syndiagenesis to epidiagenesis, but this would only be a minor episode or series of episodes in the whole evolution of a subsiding basin. In the epidiagenetic phase, emergence (diastrophic or eustatic) permits deep penetration by ground water, and, in appropriate basins, the establishment of artesian systems, that may in certain regions reach far below the present m.s.1. Inasmuch as meteoric waters are normally saturated in oxygen and CO,, a completely new geochemical cycle is usually initiated. Oxidation becomes very general and the pH will tend to drop, except where the waters are heavily contaminated by connate reserves or where they encounter precipitated soluble salts. Epidiagenesis of the old, stable cratons is a pervasive process in soil development and landscape control, inasmuch as it is constantly interrupted by transgressions and then reintroduced. A global eustatic rhythm of rise and fall of sea level tends to be superimposed on regional epeirogenic crustal warping. The alternation of stabilization and reactivation is the essence of the theory of biorhexistasy (Erhart, 1956; Termier and Termier, 1963). “Biostasy” is the condition marked by low relief, dense forest cover, and chemical leaching, with export of silica and calcium. “Rhexistasy” is the reactivation stage, when stream profiles are lowered and landscapes dissected, generating a large supply of clastics, while loss of vegetation leads to desiccation of plainlands. In tropical regions the clastics often include large supplies of hematite-rich debris from the mechanical breakup of laterites. The primary epidiagenetic effect of biostasy is leaching and that of rhexistasy is capillary concentration and deep cementation (Fig. 2-7). Ground water is divisible into two categories: (a) the uadose water that is associated with the well-aerated zone near the surface, and characterized by episodic saturation and seasonal capillarity; (b) the phreatic water that lies below the water table in the zone of more or less permanent saturation. The upper part of the vadose zone is usually marked by leaching. At the boundary between the two is the water table, and in the capillary fringe above it, especially in seasonally arid climates, there tends to be reprecipitation of carbonates (Bathurst, 1975; Milliman et a]., 1974), silica (Flach et al.,

45 Ca C O j Synqiagenesis lnselberg

Dense forest

/

/

Chemical

I

,

a

Warm, clear, wide

A. BlOSTATlC STATE 1, LOW land masses 2. High sea level (thalassocratic 1 3. Humid climate (high oceanicity) 4 . Rich vegetation 5. Chemical weathering

limited

€3.

RHEXISTATIC STATE 1. Emergent landscape, higher relief 2. Lowered base level (eustatic low) 3 . A r i d t o semiarid climate 4. Reduced vegetation, savanna 5. Mechanical erosion

reduced Shelf sea

Eustatic

leaching of fresh sediments)

RWF 1981

Fig. 2-7. Epidiagenetic effects of biostasy and rhexistasy.

1969; Stephens, 1971; Langford-Smith, 1978) and iron oxides (Turner, 1980; Allen, 1980), which become the “hard pans” of soil scientists and the “duricrusts” of geomorphologists. Rapid penetration of the bedrock by ground water is largely restricted to sedimentary rocks, although it also enters joints and fault zones in igneous rocks. In this way, given long periods of geologic time, weathering gradually penetrates into the most compact rocks, and even into mineral cleavage planes. One of the more spectacular processes of epidiagenesis is the rehydration of anhydrite to gypsum that produces enterolithic expansion with convolutions within thin-bedded layers, or surface geomorphologic effects, such as mounds, rolls, tepees, and chaotic blockfields, in semi-arid country. (Gypsum-anhydrite relationships are discussed further under “Hydration-dehydration”, below.) The rapid oxidation of such mineral compounds as pyrite (FeS,) will, as discussed earlier, tend to raise the temperature of deep ground waters far beyond the limits normally expected from the geothermal gradient, and further re-solution of certain salts will be facilitated. This general geochemical revival permits renewed cementation, and mineralization of fault and joint zones, thus effectively contributing to the final lithification of the rocks. An important aspect of the epidiagenetic phase is the state of permeuhility

46 achieved in the preceding diagenetic episodes. Well-compacted and unjointed shales may, for example, be so well sealed that no epidiagenesis is possible. Fresh pyrite crystals may be broken out of them with no trace of oxidation. A porous sandstone on the other hand may be thoroughly sluiced through. The question arises as to where weathering begins and epidiagenesis ends. Van Hise (1904) subdivided his “katamorphic zone” into two: a “belt of weathering”, and a “belt of cementation” (i.e., the anadiagenetic phase). The “weathering belt” was taken to include all of the zone affected by circulating water, which can be classed as epidiagenetic. Certainly there is a link as Blackwelder (1947, p. 500) brought out. Krumbein (1947, p. 171) had written on “weathering as a diagenetic process”, but evidently referred mainly to soil-forming processes, in other words, to strictly superficial phenomena. He said: “Weathering is essentially a process of delithification, but it is much more than a simple reversal of the reactions and processes of lithification. Weathering is in large part a phenomenon of oxidizing environments, whereas diagenesis proper occurs mainly under reducing conditions”. Interpreted in this light, “weathering” might be considered to embrace the whole phase of epidiagenesis, which, if justified, would do severe damage to generally accepted definitions of weathering: e g , the group of processes, such as the chemical action of air and rainwater and of plants and bacteria, and the mechanical action of changes of temperature, whereby rocks on exposure to the weather change, decay and finally crumble into soil. It might seem wise, therefore, to keep the term weathering for these surface processes, and to recognize that the oxidation zone, often reaching to depths of’5000 m or more, is a special @on-reducing) stage of diagenesis. The nature of the depth-time relations of the epidiagenetic phase are controlled by the accessibility of oxygenated waters, provided that the anadiagenetic cementation has not blocked the permeability of the porous sediments. There will be a tendency for meteoric waters to penetrate as soon as a given basin becomes even partially emerged, and thus to set the last phase in motion. Following hydrodynamic theory, pools or pockets of oil and natural gas tend to become isolated, and even to be pushed into tilted reservoirs by the fluid motion. Some oil basins of this sort are now reached well below sea level at depths over 8000m. In mountain ranges, meteoric waters are encountered in fault zones and in pervious strata, and such waters should, theoretically at least, extend to still greater depths. The duration of such exposure is almost unlimited, and the erosional phase of a mountain system may range from 10’ to lo9 years.

47 AUTHIGENESIS AND SOME DIAGENETIC PROCESSES

Other chapters in this book and its companion volumes deal more completely with almost all diagenetic processes, but it may be useful to review briefly those processes that involve authigenesis more or less in the order of their appearance through the three evolutionary stages of diagenesis. There have been several attempts in the past to restrict the term authigenesis to a particular phase of diagenesis, to the syngenetic mineralization that occurs in the early or syndiagenetic phase, as distinct from epigenetic mineralization of the late phase (summarized by Teodorovich, 1961). Pettijohn (1957, p. 662) drew attention to a Russian classification by Baturin (1937), which distinguished these two divisions as: (a) early diagenesis, marked by authigenic mineralization; (b) late diagenesis, marked by epigenic mineralization. The suggested correspondence has not been widely accepted, however. Geochemical reactions The chemical reactions involved in authigenesis are relatively simple in principle, but are complicated by the extremely “open” type systems, the multiple mixing of constituents, and the uncertain effects of time. As Sujkowski ( 1 958) emphasized, the duration of any reaction may be quite brief, but the diagenetic processes advance in jumps as the chemical media and physical conditions change. Certain metastable mixtures only become reorganized on attaining critical geochemical thresholds. Diagenesis generally tends to lead towards the simplification of the number of rock components. Generally speaking, most of these reactions are those studied under the topic “weathering”, which usually involves the unmixing of complex molecules, though sometimes the reverse is true, i.e., the construction of complex compounds which is the essence of authigenesis. In summary these reactions are: hydration-dehydration, hydrolysis-dehydrolysis, ion adsorption, cation or base exchange, oxidation-reduction, carbonatization. ( a ) Hydration-dehydration. Hydration-dehydration reaction involves the take-up of water on crystallization and recrystallization with loss of water, e.g., the gypsum-anhydrite relationship: CaSO,

+ 2 H 2 0+ CaSO, - 2 H 2 0

Hydration also occurs during hydrolysis, oxidation, and carbonatization (see, for example, Berner, 1971; Blatt et al., 1980; Garrels and Christ, 1965).

48 Hydration is most common during syn- and epi-diagenesis, whereas dehydration is prevalent during anadiagenesis. ( h ) Hydrolysis-dehydrolysis. There is a tendency for water to react with dissolved salts (in the chemical sense). This is hydrolysis. Water plays the role of a base and yields hydroxyl ions in solution; this is in contrast to an acid that yields hydrogen ions, which on reacting with H,O give H , O + , the hydronium ion. Most silicate minerals are susceptible to hydrolysis. Thus olivine hydrolyzes to serpentine:

2 Mg,SiO,

+ 3 H,O * 3 MgO ~2 SiO, . 2 H,O + Mg(OH),

Relative susceptibility is useful in identification of unknown minerals and can be demonstrated experimentally by means of the “abrasion pH” (Stevens and Carron, 1948; Keller et al., 1953). The mineral is ground up with distilled water and the p H is measured. Feldspars give a pH of about 10 and wollastonite gives a p H of 11. It is found that under these conditions metal cations are liberated, e.g., C a 2 + , Mg2+, N a + , and K + . Thus orthoclase may break down to kaolinite, liberating K and silica. Tamm ( 1925) demonstrated that rock flour pulverized by glacial grinding is hydrolyzed even at the stage of its incorporation in glacial melt water; indeed, hydrolysis is the principal reaction of weathering. This continued trend through geological time has largely determined the fact that the ocean is alkaline (Garrels and Mackenzie, 1974; Holland, 1972). Most detrital silicates (notably feldspars, micas, and clays) carried into the ocean are in an incomplete state of hydrolization. Depending upon the pH of the parent rivers (which vary from extreme acidity to alkalinity), the reaction may be driven to the right, involving further hydrolization, or to the left, that is silicates become dehydrolyzed. Upon burial, the p H in a marine sediment is likely to drop immediately from 8 to 6.5 or lower, driving the reaction to the left, but in the early burial and anadiagenetic phases, the pH rises and the reaction swings to the right again. In addition to this hydrolization most of the silicate weathering products also have the capacity of base-exchange and adsorption. ( c ) Ion adsorption. Ion adsorption is a peculiarity, discussed earlier, of many weathering products such as aluminium hydroxide, ferric hydroxide, and the whole family of clays. After organic dyes, H + and OH- are most readily adsorbed, generally followed, in the marine environment, by the cations: C u 2 + ,A13+, Zn2+, M g 2 + , C a 2 + , K + , and N a + ; and by the anions: S 2 - , C1- , and SO:-. The order of replacement of cations is as follows: H, Ba, Sr, Ca, Mg, Rb, K, Na, Li. For example, hydrogen ion will replace the calcium

49 ion, unless the latter is present at a higher concentration (see Chilingarian and Vorabutr, 1981). After the adsorption of H + or O H - , the adsorbent has a free electrical charge, which is a characteristic of colloidal particles. Inasmuch as clays favor the hydroxyl ion, they are often negatively charged, with the result that they tend to adsorb a whole range of rare metallic cations from ocean water (Krauskopf, 1956). It is, therefore, a vital process in authigenesis. Selective adsorption by aging gels leads to completely new minerals (Eitel, 1954, p. 458). ( d ) Cation or base exchange. A long-recognized feature of weathered silicates, such as soil clays, is their capacity to exchange cations with any passing alkaline solution (Tamm, 1925; Kelley, 1939; Russell, 1970). An acid solution (rich in CO, or “humic acids”) will tend to remove the exchangeable bases, leaving an “acid clay” (Graham, 1941). Means of measuring the “free exchange bonding energy” of the cations in soils have been devised quite some time ago (Marshall and Upchurch, 1933), whereas attempts to carry out quantitative measurements in fresh marine sediments are only beginning. It is deduced that, for example, the clay mineral illite, a muscovite degraded with respect to K, becomes enriched with that ion from sea water, whereas other clays obtain it by base exchange. ( e ) Oxidation-reduction. As mentioned earlier (in the section on Boundary

Limitations), the oxidation-reduction balance is closely related to the question of the absolute boundary limitations for diagenesis. Examples are given in a later section (on Hydration-dehydration) and only a few definitions are presented here (see, for example, Friedman and Sanders, 1978; Berner, 1971 ; Krauskopf, 1967). Oxidation implies the loss of electrons. The substances that gain electrons are called oxidizing agents or oxidants, and in gaining electrons they are reduced. In weathering, this oxidation is often effected by means of atmospheric or organic oxygen (liberated in photosynthesis); free oxygen is available in normal sea water and surface sediments. In an aqueous medium, however, it is often impossible to separate the concept of oxidation from hydrolysis. Furthermore oxidation is also involved when metallic iron is “oxidized” to FeS and then to FeS,. Immediately below the oxygenated layer of sediments is the reducing environment of “early burial”, where the above reactions go to the left. Traces of this reduction may be recognized in sediments after a complete diagenetic cycle. It will be recalled, however, that both the initial and end stages are most generally oxidizing, so that the reducing phase may be deduced only from the survival of certain minerals or by direct sampling (bottom-coring) in contemporary sediments.

( f ) Carbonatization. The carbonate ion COj- (or the bicarbonate ion HCO, ) is often known to replace silicates during weathering, particularly

50 with Ca2+ and M g 2 + .Both being always present in the ocean, they are likely to be organically reprecipitated as CaCO,, “magnesian calcite”, or recrystallized later as dolomite. Decarbonatization normally occurs in the presence of solutions of CO, and then of HCO,. The bicarbonate ion can exist in the initial (near-surface) conditions of open oceanic diagenesis, and again in the epidiagenetic phase when it is reintroduced by meteoric waters. It is rarely to be expected in the early burial (reducing) stage of syndiagenesis, and is equally unusual throughout anadiagenesis. Its role may perhaps be traced in the early syndiagenetic alteration of brucite Mg(OH), to MgCO, in dolomitization, which is discussed under “Metasomatism”. The concentration of the weak acid H,CO, in water that is in equilibrium with the atmosphere is independent of pH, but is always low. During syndiagenesis, however, the interstitial waters with pH around 8 develop far higher concentrations of carbonic acid than they do in rain water or in surface sea water. This is mainly due to the CO, of biogenic origin. Inasmuch as the processes in such complex mixtures as ocean waters and random sediments are considerably intertwined, it is convenient to consider them in more detail under headings that, at least in part, characterize the site as well as the nature of the reactions.

HALMYROLYSIS

Sediment particles reach the ocean: ( 1 ) in estuaries and deltas as rivertransported material (Fairbridge, 1980b); (2) from coasts subject to littoral erosion; (3) anywhere at the sea surface when the particles arrive wind-borne, as fall-out (pyroclastic ash) from volcanic eruptions, or by extraterrestrial transport. Reactions may begin immediately while the particles settle. Some may continue, as the sediments slowly drift along the bottom (Ewing’s “nepheloid layer”: see Jacobs, 1978). By virture of their fine grain size and exchangeable cations, the clays are particularly susceptible to electrolytic reactions. The alteration brought about is sometimes referred to as “predepositional diagenesis” or “submarine weathering”, but it is not really included in the original definition of diagenesis, nor is the analogy with subaerial weathering very precise. It seems that the best term for the process is halmyrolysis, an expression proposed by Hummel (1922), from the Greek roots hali- (sea) and myros (unguent), erroneously spelled “halmyrolosis” by Pettijohn (1957, pp. 649,662) in one place and “halmyrolisis” in another. Chemical reaction with sea water begins, of course, immediately after the sediment reaches the ocean, and continues while it is moved over the bottom or is swept along by currents as it settles from suspension. On reaching a point of even temporary stability it will be subject to contact with sea water,

51 possibly for extended periods, and may then be retransported to a deeper position. Some parts of the continental shelf and slope, as well as “sills” between basins, submerged plateaus and guyot (flat seamount) tops, are so constantly swept by currents that fine-grained materials remain only in pockets, or where trapped inside the shells of marine organisms. Inasmuch as sea water takes up elements (e.g., Ca,Si, and P) at the expense of the sediment, the term “submarine weathering” is sometimes quite appropriate (Correns, 1950), but the whole picture is more complicated. Mackenzie and Garrels ( 1966) suggested a sort of “reverse-weathering reaction”, progressively, through geologic time, removing silica and bicarbonate from sea water: (amorphous aluminum silicate) HCO, H4Si04 cations + cation aluminum silicate (mainly clay) CO, H,O. Most marine silica, however, is removed biogenically (mainly by silica-shelled diatoms). Teodorovich (1954; see review by Chilingar, 1955; and discussion by Packham and Crook, 1960) recognized no less than thirteen diagenetic facies that are subject to halmyrolysis and/or processes of early diagenesis. They have been defined in terms of pH-Eh and mineral coinponents. Six fundamental geochemical environments are noted: (1) oxidizing (ferric oxides and hydroxides); (2) weakly oxidizing, Eh = 0 well below sediment surface; (3) neutral (iron chlorites), Eh = 0 slightly below sediment surface; (4) weakly reducing, Eh = 0 precisely at surface; ( 5 ) reducing (carbonates; scattered FeS and FeS,), Eh = 0 slightly above interface; (6) strongly reducing (sulfide zone proper), E h = O well above sediment interface and approaching sea surface.

+

+

+

+

+

Gfuuconite Probably the most characteristic of all halmyrolytic phenomena is that of glauconitization (Porrenga, 1967). The mineral gfuuconite (a hydrolyzate mineral related to illite, a hydrous mica rich in iron and potassium) has long been regarded as an exclusively marine product, authigenic, and characteristic of certain shelf and slope environments. It seems possible, though rarely, that it may even be anadiagenetic (Wermund, 1961). Cloud (1955) believed that glauconitization was favored by cooler waters; this is not correct, for the most widespread occurrence of glauconite today is over the continental shelves north of Australia, as originally described by Murray and Renard (1891) and later extended b ~ Fairbridge r (1953, p. 11). This is one of the most uniformly warm shelf areas of the world, where the sediments are very often tinged green by glauconite, except where they are masked by large supplies of detrital carbonates near coral reefs, or inshore near river mouths. Although Murray and Renard found glauconites down to depths of 4000m in the ocean, it seems likely that they were transported there with

52 other sand-size particles, and that the optimum zone of glauconitization is from wave-base (say 15m) to somewhat beyond the shelf margin (say 500 m). The mineralogy and distribution of glauconite suggest that it is derived from clays, micas, and feldspars by a slow hydration and ion exchange process that is favored by slightly reducing conditions together with a free access of sea water at a p H of around 8. Reducing conditions can be established on the open shelf in a microenvironment, such as the rotting interior of a molluscan shell, and certain foraminifera, or associated with fecal pellets. Organic participation, even in micro-environments seems to be the rule (Ojakangas and Keller, 1964), in the case of existence of correct Fe3+/Fe2+ ratio. The process is evidently very slow, because glauconite is not formed in areas of rapid burial, and is often found at intermediate stages of formation; however, the shelf areas where it is most common have only been inundated during the last 10,000 years, so that the complete authigenesis of a particle 2 m m in diameter may require something of the order of 100- 1000 years. Although one lacks experimental evidence, the glauconite reaction probably lies in the activity field of the heterotrophic anaerobic bacteria ( p H = 7 - 8 and E h = O to -100 mV), and within a temperature range of 25-5°C. There is apparently a reduction in the amounts of glauconite in the pre-Mesozoic oceans (Conway, 1945), which may be related to the postulated higher pco, of the Precambrian and Paleozoic times. Tugarinov and Vinogradov-( 1961) reported Precambrian glauconites back to 1.5 . lo9 years. The older the glauconite, the lower is its Fe3+/Fe2+ ratio (Smulikowski, 1954). One of the complicating factors about glauconite is that under alkaline conditions it remains very stable, so that it may be constantly reworked, over and over again, in marine sands in some areas (e.g., through the Cretaceous and Tertiary of eastern New Jersey), thus rendering it sometimes misleading for K/Ar age determination. As soon as it is exposed in an acid soil, however, it rapidly breaks down, liberating potash (useful as a fertilizer) and iron oxides. Phillipsite and other zeolites Another product of halmyrolysis is the marine zeolite, phillipsite (its formula is sometimes given as (K,Ca)AI,Si,O,, - 4 H,O), which was first discovered in deep-sea environments by the ‘Challenger’ Expedition (Murray and Renard, 1891). These elegant little crystals, often twinned, make up an appreciable part of the “red-clay’’ sediments in parts of the Pacific, where they are not masked or inhibited by rapid deposition of other materials. In

53 places, they exceed 50% of the non-carbonate fraction. On land, this mineral is found only in association with basaltic lavas, and in the ocean it seems to occur downwind of island volcanoes, the ash showers from which have been widely distributed. It is not found in volcanic ash on land, and so appears to be the result of halmyrolysis, but the details of its formation have never been completely elucidated. Some studies suggest that palagonite tuff (basaltic glass) is altered to montmorillonite, and thence to the zeolite. Photographs show phillipsite forming inside palagonite nodules (Bonatti, 1963). Another more restricted deep-sea zeolite is clinoptilolite, which is most common in the Cretaceous-Eocene sediments (Venkatarathnam and Biscaye, 1973). Contemporary zeolitization following deep burial is hardly distinguishable from the first stage of metamorphism, the zeolite fucies (Fyfe et al., 1958; Coombs et al., 1959; Turner and Verhoogen, 1960, p. 532). This phenomenon may occur at burial depths down to about 8000 m (say 2000 bars pressure and 100-200°C), provided that there is suitably metastable parent material, in this case, fresh, rapidly accumulated volcanic ash. Tuffs, deposited from nuees ardentes, rapidly develop zeolites close to the surface, as around Vesuvius (Norin, 1955). Reactions between connate water and volcanic glass may give analcite (Bradley, 1929; Eitel, 1954, p.997). There is also the devitrification of glass to heulandite [(Ca, Na,)Al,Si,O,, . 5 H,O] and laumontite (CaAl,Si,O,, . 4 H,O), and the albitization of plagioclase. Indeed, these phenomena are no more than a bulk interaction of compounds normally found scattered through basins that have passed through the phase of anadiagenesis. One might even question the classification of this zeolite facies under the heading of metamorphism. Thus, in the West Coast ranges (Oregon and northern California) zeolite facies are regarded by Hay (1962) as simply diagenetic (see also Packham and Crook, 1960). Because of its volcanogenic derivation, zeolitization in the geologic record is an especially valuable indicator of plate tectonic activity (Sever, 1979). Clay minerals

A much-studied aspect of halmyrolysis is the authigenesis of clay minerals (reviewed, for example, by Drever, 1976). Clay and colloidal particles are transported mechanically by rivers, in suspension or adsorbed onto organic gels, or by wind, as dust (or loess), to the oceans. When transported in particulate form, the mineral will have its eventual character dictated primarily by the source region (“heritage” of Millot, 1953, 1970). Alternatively, when transported as finer colloids and ions, susceptible to electrolysis, flocculation occurs within a few hours on contact with sea water, and authigenesis can be expected to follow.

54 Russell (1970) showed that some 11-47% of K-ions delivered in solution by rivers to the ocean is removed by clay minerals during a rapid halmyrolysis. This capturing process seems to apply also for Na, but is unimportant for Ca and Mg. The Mg-cation is only taken up in an anaerobic (syndiagenetic) setting, replacing Fe that later forms pyrite (Drever, 1971). Experimental work over a 6-month period showed that the most reactive species were the montmorillonitic clays, whereas kaolinite and illite appeared to be stable within this time frame (Whitehouse and McCarter, 1958). A “concept of dualism is essential to the understanding of clay genesis”, according to Weaver (1958); he refers here to the inherited characteristics on one hand and the depositional environment on the other. The various ancient geosynclines and basins of North America are now rather well differentiated on this basis (Weaver, 1960). Clay minerals most clearly reflect the environment of deposition in young accumulations where the sedimentation rate is low and, thus, where time is available for halmyrolytic reactions. In eugeosynclines where the sedimentation rate is high, rapid burial seals in the components, and source region is most clearly reflected by the minerals eventually stabilized. Upon burial in a largely clayey facies, there is a natural tendency towards the creation of a relatively impervious shield, so that syndiagenetic alteration is restricted by the very limited supply of ions available in the connate waters immediately in contact with each of the clay particles. During early anadiagenesis, an almost closed system may be created and clay minerals buried in this environment show little change in a column of over l000m (Milne and Earley, 1958). With progressive loading and tectonic evolution, fracturing permits recirculation and natural chromatography (see section on Natural Chromatography) is facilitated. Burial in a pervious sandy facies, however, permits very extensive authigenesis. Systematic regional studies of deep oceanic clays have been made by Biscaye (1965) and by Griffin et al. (1968). The broad pattern that is emerging seems to suggest that there are definite geographical, latitudinal provinces, which are dependent upon these two factors: (1) inheritance (source materials and conditions) and (2) authigenesis (local materials and conditions). Grim (1958) has remarked that certain varieties of clay minerals are more “at home” in certain environments than others, and this is the basis of authigenic differentiation. Whereas a very large number of clay minerals are known, only three principal types will be discussed here: illite, montmorillonite, and kaolinite. Illite, the hydrous mica group of clay minerals, was first named by Grim et al. (1937); these are the so-called “low alumina clays” (10-20% A1,0,). Crystallographically, they appear to have essentially the same lattice structure as muscovite, but are degraded with respect to K:

55 where y is around 1- 1.5. An ideal end-member could be:

Illites seem to be formed during weathering in the moderately high-pH soils of cool to temperate climatic belts: for example, the partly weathered till and loess plains of Illinois. In the ocean they are either reworked from older illites or authigenically derived from less stable clays; illite appears to be the most stable of the clays in the marine environment (Drever, 1977). During syndiagenesis it may emerge from the halmyrolysis of montmorillonite or kaolinite (see below). Particularly K is adsorbed in these reactions, which are favored by the presence of C a 2 + ,M g 2 + , and ferrous iron. Halmyrolysis may well be initiated the moment that the clay colloids reach the saline waters of the ocean. Griffin and Ingram (1955) showed how, in a single estuary (the Neuse River in North Carolina) that drains a hinterland of relatively uniform metamorphic rocks, kaolinite is by far the most abundant mineral introduced. As the water becomes progressively saltier, there is considerable mixing with “chlorite”, tb be replaced in turn by illite near the mouth. Some of the kaolinite gets through this transportation phase, however, because it is present in most oceanic sediments, notably in the tropical latitudes, offshore from the deeply weathered cratons (Biscaye, 1965; Porrenga, 1967). Illite is the commonest clay type in the Paleozoic fold belts of the world, and thus it is not surprising that it is the dominant type in the temperate North Atlantic as well as in the Pacific (Griffen et al., 1968). They suggested, incidentally, that a weathered illite which had lost most of its K before reaching the sea could be reconstituted by ion adsorption during early diagenesis. In the tropical Atlantic, the illite: montmorillonite: kaolinite ratio is approximately 1 : 1 : 1 (Biscaye, 1965), but 8 : 1 : 1 in the temperate latitudes. K/Ar isotopic age-dating of these North Atlantic illites give mean ages in the 300-400 m.y. brackets (Hurley et al., 1963), which strongly suggests simple reworking and transport from Paleozoic sources. It would seem thus that authigenesis plays a rather minor role in the production of illite, although it is certainly possible to interpret this “Paleozoic” age as merely the statistical result of sedimentologically “smoothed” mixing of still older clays with a fraction of modern authigenic products. The supply of K available in the connate water would in any case restrict the degree of authigenesis in the halmyrolytic stage. Isotopic dating of Pacific clays (Hurley et al., 1963) gives a Late Mesozoic age, which also suggests a minor role for authigenic clay components during the syndiagenetic phase. This does not preclude further K being taken up during anadiagenesis. It would be interesting to learn if the K/Ar ages of

56 deep-sea illites decrease as one goes down the D.S.D.P. cores. Independent confirmation of the provenance rock age, transported dusts, and deep-sea sediments is provided by Rb/Sr ratio and Sr-isotope data (Biscaye and Dasch, 1971; Biscaye et al., 1974). It is noteworthy that the K values in clays rise on going back into the Precambrian (Nanz, 1953). Conway (1945) suggested earlier that there was a peak of K extraction in the oceans in Late Precambrian time. Whereas he suggested subsequent increased organic activity as the cause, it may perhaps be that progressive rise of oceanic pH is less favorable for illite halmyrolysis today. Also unfavorable to the extensive formation of illites are the widespread desert conditions of today. Montmorillonite formula is generally given as A1,Si,O,o. (OH),nH,O. The structural formula of montmorillonite group of minerals is best expressed as:

where R3+-Al, F e 3 + ,Cr; R2+- F e Z f , Mn2+, Mg, Zn, Cu. In this “expanded” or three-layer lattice clay group, some of the aluminium may be replaced by Mg (suponire) or by Fe (nontronite). A general term for the montmorillonitic clays derived from volcanic ash is bentonite. Because of its expanded lattice, montmorillonite is one of the most reactive types of clay and exhibits a considerable propensity for base exchange and ion adsorption (Kelley, 1939, p. 434). The order of replacement is generally Naf < K f < Mg2+< Ca2+< H + , i.e., that potassium replaces sodium, etc. Its cation exchange capacity is quoted as 60-150 mequiv./lOOg, as compared with 3-15 for kaolinite and 10-40 for illite. Rapid diagenetic reactions in the halmyrolytic stage are thus to be expected, even before the clay particle reaches the bottom; however, when buried and cut off from acid solutions, montmorillonite may even survive from Paleozoic times (although admittedly rather rarely). Burial in excess of 3000m clearly adds heat and pressure in addition to the chemical energy so that the end-product is normally illite (Keller, 1964). Montmorillonite seems to be the product of simultaneous weathering of feldspars and ferromagnesian minerals (Ross, 1943) from mafic igneous rocks, and especially fresh pyroclastics (ash), in waterlogged alkaline soils or newly-accumulated ash deposits of moderate pH but low Eh (poorly oxygenated). This is a feature of intermediate and “Mediterranean” latitudes, and formation of montmorillonite occurs particularly in semi-arid lakes (e.g., North Africa), swamps and littoral lagoons (e.g., Texas Gulf coast), as well as in volcanic regions (with alkali-pyroclastics). Experimental procedures have shown that in addition to pH, the flow rate of the interstitial water, the particle size, and the temperature are the key parameters in

57 such clay formation and decomposition (Correns, 1963). Weathering of freshly extruded basalts on the sea floor near spreading centers provides a K-rich smectite to the regional sedimentation (Hart, 1973). As regards distribution, montmorillonite is widespread in the Pacific (Griffen et al., 1968), no doubt because of the extensive occurrence of volcanic ash. In the tropical Atlantic it shares a 1 : 1 : 1 relationship with illite and kaolinite, as noted above, and may be related in part to wind-borne supplies from the semi-arid regions. It is also common (even dominant sometimes) in the Gulf of Mexico. In the North Pacific and North Atlantic, it seems to be related to the high-latitude volcanoes. The role of montmorillonite in the North Atlantic is very much masked by the high content of primary (Paleozoic) illite there. The absence of illite from the primary weathering cycle in most of the tropical soils, however, suggests perhaps that its presence in tropical oceans is the result of the breakdown of montmorillonite and kaolinite (see discussion by Van Andel and Postma, 1954). The fact that montmorillonite is widespread in the tropical Pacific without reacting rapidly with the exchangeable ions available is puzzling, but inasmuch as this region is more or less smothered by carbonate oozes, the surface and immediately subsurface pH is kept near 8, which may stabilize the montmorillonite. It would be interesting to know what becomes of the latter at depths of a few cm below the high bacterial layer (where the pH is less, in spite of the calcium-rich environment). It is noteworthy that in more or less land-locked lagoons, where evaporation becomes dominant and gypsum is precipitated, the pH is generally above 8.5 and montmorillonite is replaced by illite (Millot, 1953). In a study of some Permian clays from Kansas, M.F. Norton (personal communication, 1958) found that there was a positive correlation between the abundances of dolomite and the high-Mg montmorillonites, and between those of limestone and illite. Millot remarked that ancient limestones are generally associated with illite. Thus, the pH and Eh of the depositional environments were normally high (pH = 7.5-8.5), and sediments were poor in organic matter. Montmorillonite appears to be favored by the lower Eh (a feature of many dolomites) as shown by the scattered pyrites in the latter (Fairbridge, 1957). Kuolinite is the simple two-layer lattice clay, characteristic of the so-called “high-alumina clays” (20-40% A1203). It does not appreciably expand upon increasing water content, and does not exchange iron or magnesium. The typical group formula is: A1,0, . 2Si0,. 2 H 2 0 , sometimes given as: A1,Si ,05(OH), or as A1,Si 40,,(OH)8 Kaolinite is well known as the stable end-product of laterization or latosols, produced by deep tropical weathering with pH of 5-7; but being well oxygenated, the Eh of environment in which it forms is neutral to high.

58 Its formation is especially favored by aluminous igneous rocks such as granite (Ross and Kerr, 1931). In this process, C a 2 + , M g 2 + , and Fe2+ are leached out. In subtropical regions silica and iron rise to the surface by capillarity to form a ferruginous duricrust, whereas the Ca2+ and Mg2+ tend to pass into the ground water and drainage system and then to the sea. Kaolinite itself tends to be picked up in the arid lands by wind or subject to fluvial erosion only as a result of tectonic and eustatic oscillations of base level. Under such conditions it is transported to the ocean (as during the Quaternary). In the tropics it tends to dominate the interglacial clastics, being replaced in part by feldspar in the glacials (Damuth and Fairbridge, 1970). In the ocean today, kaolinite is, therefore, mainly a transported, nonauthigenic sediment, and is most widespread in the tropical Atlantic where the principal supply is eolian (Biscaye, 1965), though partly fluvial (fed by the Amazon, Orinoco, Congo, and Niger), according to Correns and Von Engelhardt (1938), Millot (1953), and Heezen et al. (1960). It is also widespread in the tropical Pacific (Griffen et al., 1968) and the Indian Ocean (Venkatarathnam et al., 1976). Authigenic kaolinite is found here and there in former marine sediments, but it is evidently epidiagenetic for it is restricted to areas where the interstitial waters are fresh (Shelton, 1964). Grim (1951) pointed out that kaolinite and montmorillonite are much less common than illite in ancient sediments, and thus one may suspect that diagenetic replacement by illite has been operative. The possibility of longterm geochemical changes in sea water, however, should not be forgotten (see Conway, 1945). Millot (1 953) posed a good question: how is kaolinite, which is formed in the acid environment of the continent, preserved from diagenesis in the alkaline environment of the sea floor, with which it seems to be out of equilibrium? Zen (1960) believed, on the other hand, that the intimate association of calcite and kaolinite implies an equilibrium condition with sea water. In Paleozoic rocks, he noted the five-component system calcitedolomite-chlorite-kaolinite-quartz, which seems to be a stable mineral assemblage. Millot made the helpful suggestion that where kaolinite is rapidly accumulating on the ocean floor, it is quickly sealed in by a relatively impermeable layer, and with plentiful bacterial action the pH quickly drops to 6-7 and the H 2 S lowers the Eh potential. To close this section, one may conclude that the limits of clay diagenesis are still rather poorly defined, especially the lower boundary of halmyrolysis.

59 Oxidation and reduction-processes and environments Inasmuch as normal marine basins are in constant circulation, oxygen saturation is generally maintained, and the normal open-sea floor is oxygenated at the sediment-water interface. Euxinic environments, such as the Black Sea, constitute special cases where reducing conditions, even in the liquid medium, exist everywhere below a certain “sill depth” (Goldhaber, in: Fairbridge and Bourgeois, 1978, p. 297; Degens and Ross, 1974). In well-ventilated basins, the oxidation of sedimentary particles at the sediment-water interface is another of the earliest possible diagenetic processes. The slower the sedimentation rate, the more complete is the oxidation. Thus, the deep-water areas far from land, e.g., central Pacific, are the classic sites for the red clay deposition of Murray and Renard (189l). due to the oxidation of iron. This is also the site of manganese nodule formation which occurs around any sort of nucleus from a glacial boulder to a pelecypod shell. This coating of MnO, is a function of the time exposed at the sediinent-water contact. Divalent Mn2+ present in sea water is apparently oxidized to the tetravalent species Mn4+ (Goldberg and Arrhenius, 1958). Large boulders do not contain a coating of MnO, on their deeply buried side. Such boulders were carried out into the Arctic Basin by icebergs during a time of open water conditions (a controversial question, but the present writer suggests about 5000- 10,000 years ago during the “climatic optimum”, Fairbridge, 1961). After cool conditions returned, the basin was covered by pack ice and further sedimentation was greatly retarded. Inasmuch as the thickness of MnO, accumulated is only 2 or 3mm, the accumulation rate is around 0.0003 mm per year. The production of red beds in the geological past is a complex problem, depending in part on source materials and in part on diagenesis (Pettijohn, 1957; Van Houten, 1961; Allen, 1980). Red beds are often continental, but include the marine “red muds” (Twenhofel, 1950, p. 331) of the Orinoco, Amazon and Yangtze; they may extend far out to sea in depths down to 2000m, but the color is steadily lost by reduction. According to the theory of biorhexistasy (Erhart, 1956) long-term stable conditions under subtropical weathering conditions (the norm for most of geological time) leads to a general deep weathering resulting in lateritic soils, accompanied by removal of calcium and SiO,. Such times (biostatic) are marked in the marine realm by limestones and cherts (Erhart, 1973a). A cyclic lowering of the water table which takes place eustatically during periods of cold climate and tectonically lowered marine basins (“bathygenesis” of Termier and Termier, 1963) leads to a mechanical break-up of old soils (rhexistasy), to stream dissection, and to transport of red, iron-rich

60 lateritic debris into continental basins and to the ocean. Such sediments are so rich in iron oxides that even in deltas (cf. the Devonian Catskill-type deltas of the Appalachians), which normally develop reducing environments, the red colors often persist, although green layers and mottled and bleached patches bear witness to reduction by H,S (Moulton, 1926). Buried along with sediment there is always a certain amount of organic matter, which serves as bacterial nutrient, the principal source of geological reducing agents (Irving, 1892). In some areas (as off deltas) and likewise in euxinic basins, where bottom scavengers are inhibited by the H,S poisoning, this amount is very large. Inasmuch as open-sea floors are normally inhabited by both epi- and in-fauna, there is a considerable and rapid consumption of much of this organic nutrient. The proportion of organic material that is ultimately incorporated in the sediment (below burrowing depth) depends on two factors: ( 1 ) the size and vigor of the benthonic population and (2) the rat,e of deposition. As Twenhofel (1942) pointed out, a slow sedimentary rate permits thorough scavenging. He stated (p. 105): “It seems a paradox that the more congenial the conditions on the sea bottom for bottom dwelling forms and the more numerous the colonization by organisms, the less likelihood there is of many fossils in the sediments which finally attain entombment. In other words, an abundant bottom population under conditions of slow deposition produces deposits with few complete shells and more or less complete elimination of all nutrient matter. Accumulation of organic materials is not possible under such conditions”. By the same token, the bacterial population in these low sediment deposition rate areas is also low in the sediments just below the surface and, thus, the opportunities for authigenesis by reduction under these conditions are strictly limited. Two distinct environments are so characterized: (1) continental shelf and slope regions where the sedimentation and subsidence rates are low, and (2) abyssal plains and rises far removed from continental sediment sources. In contrast, shelf regions near deltas or other sediment source, and particularly abyssal plains richly fed by turbidity currents, have their benthonic populations constantly smothered by seasonal, or longer cyclic, invasions of sediment that seal off the organic debris and provide a large reservoir of bacterial nutrients. This is the ideal site for authigenesis by reducfion. In estuaries, Baas Becking and Moore (1959), have found that the aoeruge organic content is 12%, and may exceed 25%. In the same samples the iron, initially in the form of FeO(0H) is found to become completely reduced and anywhere that the organic content of the sediment is above 2% there is an excess of H,S.

61 The principal minerals so formed are the common ferrous sulfides. marcasite and pyrite, FeS,. These are produced by reduction of iron oxides and the various hydrates (first to FeS and later to FeS2) by the action of H,S liberated by sulfate-reducing bacteria and by bacterial breakdown of organic sulfur compounds. It seems that marcasite (the orthorhombic form) is produced under neutral to acid conditions (pH generally less than 7.0). whereas pyrite (the isometric form) is favored by slightly alkaline conditions (Allen et al., 1912; Newhouse, 1927; Tarr, 1927). According to Berner (1970. 1971), the process occurs in steps, through the monosulfide. It is observed that pyrite is the common form associated with impervious clays and shales (Love, 1963), whereas marcasite is most often found in sands, silts, chalks, and limestones, as well as acidic fresh-water swamp deposits such as coal formations. Edwards and Baker (1951) seemed to be under the impression that marcasite was exclusively fresh-water, whereas pyrite was of marine origin. This is not so. however, because bacterial action lowers the pH of fresh marine sediments into the stability range of marcasite. It may be suggested that this is a function of the ‘permeability of the sediment. The clays under reducing conditions rapidly become alkaline and can preserve a high p H for extended periods of time, whereas the pervious sands, calcarenites, etc., favor the aerobic bacteria, which keep the pH low until most of the organic matter is consumed. The Eh, however, will steadily drop as the organic material breaks down. Even after saturation with H,S ( a weak acid), the p H is still low and calcite fossils are generally destroyed (Mosebach, 1952). Considering some of the evidence of the geological past, it must be noted that pyritized fossils, such as typically pyrirized ammonites in the Mesozoic, are often associated with black shales which lack any benthonic fauna. General poisoning of bottom conditions might occur from time to time, even without an actual barring of the basin, for example by seasonal or cyclic invasions of protista or algae (“waterbloom”) (Rutten. 1953), or other sources of mass mortality ( Brongersma-Sanders, 1957). Reducing environments are mainly syndiagenetic and also anadiagenetic. evidently produced during or soon after deposition (in the time range of about 1-1000 years), with Eh ranging from 0 to -400 mV and pH either slightly above or below 7, as appropriate either for marcasite or pyrite formation. Exceptionally, the iron sulfides are also formed at considerable depths (3000-5000 m) as observed in deep oil fields, namely, under anadiagenetic conditions. In such cases it would appear that sulfur bacteria were living on the petroleum hydrocarbons (Bastin, 1926; Ginter, 1938). The fate of the sulfides after the early burial stage is partly illustrated by the 32S/34Sisotope ratio (see introductory notes, in the section on Syndiagenesis). The similar isotope ratios in recent sulfides, and in a Precambrian

62 shale more than 1.8 . lo9 years old (Ault and Kulp, 1959), suggest that there may be little change in the nature of the bacterial reduction during that time-span. It has been pointed out by Holland (1973), however, that rapid fluctuations in the isotopic composition of marine gypsum and anhydrite disclose comparable variations in the sea water itself within time frames of lo7 yr or less, as dictated by global paleogeographic changes. The sulfates below the level of the syndiagenetic phase may be derived either from evaporites, or from the re-oxidation of the sulfides. On emergence into the oxidizing conditions of the epidiagenetic phase, nodules of pyrite and marcasite tend to acquire a coat of limonite, and this in turn is normally externally dehydrated to hematite. The marcasite, in general, seems to be the less stable of the two forms, and the centers of nodules are often found to be ultimately broken down to a grayish powdery form rnelnikovite. An interesting and unusual example of reduction in the epidiagenetic phase (otherwise very generally oxidizing) occurs in the anhydrite caprock of certain salt domes. The anhydrite was probably laid down as gypsum in evaporite deposits that rose up diapirically during the anadiagenetic phase. The anhydrite becomes concentrated in the caprock by differential solution of the more soluble halite and other salts. The heavily fractured domal structures lead to the migration and accumulation of petroleum hydrocarbons, which provide nutrients for sulfate-reducing bacteria ( Desuljovibrio desulfuricans). The H,S gas is then produced, which on oxidizing (assisted by thiobacteria) gives rise to enormous native sulfur deposits. The biologic nature of this reduction process is proven by the low 32S/34Sratio of the sulfur with respect to the sulfate (Feely and Kulp, 1957). Furthermore, the 12C/13Cisotope ratios in the calcite of the caprock are typical of petroleum and not of marine limestones, suggesting that the C0:- was derived likewise from the hydrocarbons. Under certain conditions of oxygenation, generally epidiagenetic, and under conditions of heavy tropical rainfall (high CO, 0, intake), there may even be a liberation of H,SO, from pyrite or marcasite. Some mine waters actually have a negative pH, which can be explained by the following reaction:

+

2 FeS,

+ 2 H,O + 7 0,

+

2 FeSO,

+ 2 H2S04

It may, however, be better represented (Baas Becking, 1959, p.61) by the equation: FeS,

+ 8 H,O

-+

FeSO,

+ HSOT + 15 H + + 14e-

Inasmuch as FeSO, is stable only in an anaerobic environment, it is likely to hydrate to melanterite, FeSO, . 7H,O, and eventually perhaps to

63 coquimbite, Fe,(SO,), 9H20, in the epidiagenetic stage. In the course of epidiagenesis, which is likely to be protracted and frequently episodic, oxidation is followed by desiccation, resulting in the deposition of hematite or limonite in Liesegang rings. These are precipitation patterns that are either planar (in joints) or omnidirectional and roughly concentric (in porous media). They were first documented by R.E. Liesegang in 1896 (Stern, 1954), who recognized them as diffusion phenomena due to a periodic alternation between solution mobility (diffusion) and supersaturation (nucleation and precipitation). In a poorly stratified sandstone they may sometimes be confused with penecontemporaneous sedimentation phenomena, such as slumping, load casts, and concretions. The Liesegang phenomenon is not, however, limited to the epidiagenetic phases of sediments, being reported in the syndiagenetic phases of sediments, both marine (Stetson, 1933) and fresh water (Sugawara, 1934). They are also noticed to be diffused in syndiagenetic cherts (Bissell, 1959). Hydration-dehydration-processes and environments

Inasmuch as one deals with a watery medium in most sedimentary regions, initial hydration or hydrolysis is the rule. The gradual compaction of the basin during anadiagenesis, however, not only raises the confining pressure, but also tends to drive off the interstitial and bound water. This phase is marked by dehydration and dehydrolyzation; in the epidiagenetic phase rehydrolization or revived hydration may be expected.

( a ) Gypsum-anhydrite. The most prominent reactions involve gypsum, CaSO, * 2 H,O, and anhydrite, CaSO, (Berner, 1971). The hydrated sulfate is produced as a primary evaporite deposit as soon as ordinary sea water is concentrated to somewhat less than 50% of its original volume, as in lagoons along the Texas coast or in the Persian Gulf. It has been stated that when the water temperature exceeds 42"C, anhydrite becomes the stable phase (Conley and Bundy, 1958); however, this has not been so observed in nature. Gypsum is also formed in the early burial stage of syndiagenesis by bacterial action, e.g., in the lagoon muds of New Caledonia (Avias, 1953, 1956). I t is also found in certain beachrocks (MacFadyen, 1950). In some deep-sea carbonate-ooze environments, authigenic gypsum is also found at 0.3-4 m below the surface, in water depths of 1000-5000 m (Briskin and Schreiber, 1978). Evidently early syndiagenetic, it is suggested that episodic influxes of C0,-rich Antarctic bottom water during interglacial climatic fluctuations cause carbonate solution, liberating calcium in the near-surface sediment layer. In these pore waters sulfate content is directly proportional to sedimentation rate (Berner, 1980). During glacial conditions,

64 the pack ice reduces the bottom conditions to near-stagnation and carbonate corrosion ceases. This euhedral gypsum is not to be confused with clastic, re-worked gypsum, sometimes carried into the deep sea as turbidites. Petrographically, two forms of this secondary gypsum are recognized (Holliday, 1970); a porphyroblastic and an alabastrine type. The first form may be syndiagenetic, but may recrystallize to the second form, whereas the latter is most common as an epidiagenetic product. Within the littoral sabkhas bf the Persian Gulf, North Africa, and Baja California, the present writer has indeed found anhydrite nodules a few centimeters below the surface and they are clearly of Holocene age. But this is not to demonstrate that they are direct precipitates. Quite the contrary, i t now seems clear that the nodules were formed as gypsum, but were dehydrated to anhydrite under high temperature, i.e., under insolation (West et al., 1979). Experimental growth of primary anhydrite at low temperatures and normal pressure, with salinities resembling those of coastal lagoons and sabkhas, has always failed in the past under abiogenic conditions. Recently, however, in the presence of macromolecular organic compounds, characteristic of the hot, semi-arid littoral belt, Cody and Hull (1980) have now succeeded in growing well-crystallized anhydrite by primary nucleation. A very remarkable process, operating in the Persian Gulf on a Kuwait sabkha has been described by Gunatilaka et al. ( 1 980). Gypsum forms a fine crust on the halophytic plants and rapidly dehydrates to anhydrite, which retains the pseudomorphic texture. Upon the eventual death of the plant, the anhydrite hydrates once more to gypsum, and back again to anhydrite forming a bed up to 2 m thick. From field evidence, the present writer is of the opinion that dehydration from gypsum to anhydrite occurs most commonly upon burial to depths of the order of 100m or so (confining pressures of ca. 50 kg/cm2). Thus, for example, anhydrite has never been observed forming today in the Persian Gulf, but i t is extremely widespread all over the Mesopotamian Basin in Tertiary formations that, near the basin margins, have been buried under little more than l00m (Fig. 2-8). The gross effect of dehydration is to reduce the volume of the formation; from l 0 0 m of gypsum only 62m of anhydrite would remain. MacDonald (1953) believed that it would require an 800-m load of sediment to set this reaction in motion, but Braitsch (1962) and Borchert and Muir (1964) stated that even a very slight dynamic metamorphism (with shear stresses, produced by local faulting and uneven settling of the basin) would greatly reduce this figure. It also seems probable that at certain times in the geologic past under distinctly different paleogeographic conditions and somewhat higher mean

65 Growth of gypsum by displacement within pre-existing sediment Peposition of gypsum in standing body of water,

Meteoric water

Ill,,

replaces anhydrite E PIDIAG E,NES IS

SY NDlAGENESIS

lGypsuml ANADIAGENESIS

I

I 1000 m I

Growth of anhydrite by replacement and as

I

gypsum after anhydrite \

Fig. 2-8. Schematic diagram illustrating the gypsum-anhydrite diagenetic cycle. (Modified after Blatt et al., 1980.)

temperatures primary precipitation of anhydrite may have occurred, so that not all sedimentary anhydrite should be regarded as the result of diagenetic dehydration (Fairbridge, 1967a). Annual layers of gypsum in the Tertiary of Sicily are observed to pass upwards, with increasing grain size, into anhydrite, which apparently represents a high summer peak (Ogniben, 1955). Likewise, in Mesozoic anhydrites from Texas, primary structures, including graded bedding, have been recorded by Riley and Byrne (1961). The question of primary deposition or possible dehydration of gypsum to anhydrite must, therefore, rest on the presence or absence of primary structures. The presence of collapse structures due to the reduction in volume might also be instructive. The reverse process, that is to say, hydration in the epidiagenetic phase, is very well known; the anhydrite expands about 40%, often producing ribbons of intestine-like folds (what Grabau, 1913, called enterolithic structures). These have a curious symmetry and lack of orientation that distinguishes them from slumping or drag folds. Sometimes the expansion is limited to simple domes or rolls. The “tepee structures” of the Guadaloupe Mountains, New Mexico, and elsewhere in the semi-arid southwest may be related to this phenomenon (Newel1 et al., 1953). ( b ) Silica gels. Even more important and certainly more complicated is the dehydration of silica gels, to form authigenic opal, chalcedony and yuurtz.

66 The opal is not completely dehydrated and is, thus, somewhat unstable. It will not normally pass through the anadiagenetic phase without dehydrating to chalcedony or quartz (Fig. 2-9). The formation of quartz from amorphous silica requires something in excess of 10 million years (Mizutani, 1970). Opal is, however, a very common form to reappear in the epidiagenetic phase, wherever a very high pH permits remobilization of silica, especially in the deep ground waters of semi-arid regions. The introduction of silica into the oceanic sediments occurs in four principal ways: ( 1) colloidal form and solution, through river transportation; (2) dust, by eolian transport from deserts; (3) volcanic glass and ash, from eruptions; (4) organic fossils (radiolarians, diatoms, and sponge, alcyonarian and holothurian spicules). The annual increment of SiO, from rivers to the oceans is 4.3 . lOI4g or 0.31 mg/m3 of sea water (see Chilingar, 1 9 5 6 ~and Burton and Liss, 1973), but the ocean is undersaturated with respect to silica (0.1-4 p.p.m. near the surface and 5- 10 p.p.m. near the bottom) mainly due to biogenic removal. Amorphous silica is soluble at 100- 140 p.p.m., the solubility rising mainly with temperature, in both normal sea water and distilled water. Quartz, however, is much less soluble (only 7-14 p.p.m., according to Siever, 1957a). The shells of radiolaria and diatoms appear to have some organic protection from solution (Lewin, 1961), but rapid burial also protects them (Riedel, 1959). There are quantitative problems, however, with the contemporary silica cycle (Erhart, 1973; Drever, 1977). First, the supply from rivers may be doubled if one allows for littoral processes, Antarctic supply, and wind transport, to bring the total to 5-12 lOI4g SiO,/yr. There are strong, Thick- Age of Mineralogical composition of silica ness deposition

(Si02)

2

m.$.BP 0

Quartz

A

C

Fig. 2-9. An example of the amorphous silica (A)-cristobalite (C)-quartz transition from Japan. (After Mizutani, 1970.)

67 climatogenetic cycles, short- and long-term. Pelagic ooze removal at the present time may account for about 2 . lOI4g SiO,/yr and estuarine processes up to 1 * 1014 g/yr. The disparity lends strength to the Sillen’s (1961) idea of a silicon sink to the clay minerals (Mackenzie and Garrels, 1966). In the geologic record, silica concentration is found either as bedded chert or nodular flint (see recent collected papers: McBride, 1979). The origin of the silica and the nature of its petrogenesis is a classic problem. Non-actualistic mechanisms are often invoked. Ronov ( 1964) has demonstrated that while bedded cherts are dominant in the Precambrian, they become less frequent in the Phanerozoic, whereas colloform nodules become more and more common, especially in limestones and chalks. Non-biologic accumulation of silica gel today was first demonstrated in ephemeral lake basins in Australia (Peterson and Van der Borch, 1965), where the pH is raised to over 10 by algal photosynthesis, causing saturation by available silica. During desiccation, the pH falls and a colloform gel precipitates. At the present time, silica gels on the sea floos are distinctly rare (Twenhofel, 1932, 1950; Russell and Russell, 1936); but the present period is not necessarily a favorable one, and it would seem that among the four main sources listed in the previous paragraph there may be considerable quantitative variations from time to time. Ancient cherts have sometimes been correlated closely in time and space with periods and loci of orogenesis and volcanism (Hoss, 1957). But this may not be the only factor. One of the major discoveries of the Deep-sea Drilling Project and, indeed, one of the major developments in sedimentology in recent decades has been the discovery of extensive abyssal cherts (Van der Lingen, 1977). A systematic plot of the horizons where cherts were encountered shows that most of the chert was concentrated in the Eocene, i.e., around 50 m.y. old, and almost none was less than 10 m.y. old. There is a rough correlation of the principal clusters with the seismologists’ “horizon A” (Mid-Eocene) and “horizon B” (Early Cretaceous). Herman (1972) proposed that these accumulation peaks corresponded to episodic sea-floor spreading which, in turn, triggered vigorous oceanic circulation favorable to planktonic life. At the same time submarine volcanic emanations provided abundant silica, augmenting further the oceanic fertility. The question of direct contribution of volcanic material to the chert bands is controversial, although certainly they are very rich in siliceous plankton. The presence in them of smectite and zeolite suggests a duality (see discussions in Van der Lingen, 1977). There is controversy also over the Tethyan radiolarites, now found on land (Folk and McBride, 1978). Some argue for an abyssal environment (e.g., Folk), whereas others favor shallow-water origin (e.g., McBride). Whereas many cherts and flint nodules are the result

68 of replacement, Folk and McBride (1978) offer compelling evidence for cyclic accumulation of radiolarian opal transported by rhythmic currents, with subsequent slow transformation through a 5-25 m.y. time-span. The supply of silica solutions from soil-forming processes may be greater than that from volcanism, and soil weathering rates are conditioned by temperature and crustal stability. Today it is necessary to consider the exceptionally active erosional history of the Quaternary with its rapid eustatic ups and downs. Erhart (1956, 1963) has brought out that the long stable (“biostatic”) periods would favor liberation of soil silica and calcium (see the thalassocratic stage in the diagram by R.W. Fairbridge in: Termier and Termier, 1963, fig. 197, p. 333; see also Fig. 2-7). During the low or oscillatory sea-level phases, there is a break-up of the soil profiles (“rhexistasy”) and mechanical detritus would dominate. As a result, silica would tend to be eclipsed by floods of detritals. Indications of silica gels at times other than the present are numerous. Many instances of evidence for Phanerozoic primary silica gels have been listed by Dangeard and Rioult (1961). In the lowermost Ordovician of Poland, Samsonowicz (1948) found chalcedony forming a cast of ripple marks. Several generations of ripples were similarly preserved in the Jurassic of Normandy (Lemaitre, 1960). A chert containing a worm burrow was noted by Debelmas (1959). A fossil Exogyra, with its muscle tissue preserved in silica, seems to call for genuinely “contemporaneous” replacement (Gidon, 1959). Wetzel (1933) found perfectly preserved pollen grains in the Cretaceous flints of northern Europe and Deflandre (1936) reported in them delicate Flagellata and hystrichosphaeridae with pseudopodia perfectly preserved. From the perfect preservation of fossils, Cayeux (1897) long ago concluded that the dehydration of the silica gel began on the outside, and worked slowly inwards. There are numerous indications of the small-scale coagulation of silica gel around nuclei or centers of lowered p H during the initial stage of syndiagenesis. The roots of mangrove-type trees, for example, are permanently fixed in marine muds, where the pH-Eh conditions show remarkable fluctuations (tidal and diurnal). Cylindrical-shaped chert concentrations result and may be. termed rhizomorphs (Northrop, 1890) or rhizoconcretions (Ters, 1961, p. 172). “Fossil roots” of this sort are equally well preserved also by travertine in coastal eolianites and their paleosols (Fairbridge and Teichert, 1953). Discussing their form and occurrence in the Jurassic of Normandy (“Pierre de Caen”), Dangeard and h o u l t (1961) have made a helpful review of this entire field. Such concentrations were found forming today in the mangrove swamps of New Caledonia (Avias, 1956). The world’s oldest fossils, apparently bacteria and filamentous blue-green algae. contained in 3.0-3.6 billion year-old formations in Australia, South

69 Africa, and Canada are partly found in silicified limestones. Some of the cherts (e.g., Onverwacht, Fig-Tree, Soudan, and Gunflint formations), however, appear to be either primary or very early diagenetic, inasmuch as the extracted hydrocarbons contain curiously heavy ‘3C/12Cisotope ratios that seem to reflect an original condition (Oehler et al., 1972). The question of migration of gels towards nuclei presents further difficulty. Mutual attraction of like molecules in solutions are normally attributed to the so-called Van der Waals’ forces, assumed to exist particularly in liquids. In their absence, the random motions of electrons would appear to disperse the components. The supposed motion of gels (possibly as solutions) through well-packed sediment for distances of a meter or more towards certain nuclei, usually organic, still poses interesting problems. The dehydration of silica gels has not been thoroughly studied, but in general seems to pass through a porous opal stage to tridymite and cristobalite (Jones and Segnit, 1971). The slow crystallization to quartz changes the rock eventually to chert and leads to complex brecciation phenomena (Taliaferro, 1934, 1935). Referring to brecciated chert, CJayeux (1929, p. 371) remarked that here one deals with “a little known subject.. .consolidation of sediments as they are deposited”. Gignoux and Avnimelech ( 1937) noted that ...“on looking closely at the fragments, one often has the impression that they might have been fitted in and cemented together like pieces of a jig-saw puzzle.. .as if the debris had been formed by breaking up, scattered and the space between filled with cement”. Such brecciation, filling and rehardening has happened repeatedly with many flints and cherts, evidently due to the slow desiccation of the exterior, crushing, and refilling from the reservoir of the still plastic gel. In the hills of eastern Judea (between Jerusalem and the Dead Sea, for example) there are Upper Cretaceous chert beds 30 cm- 10 m thick, that have been involved in violent slump folding, 3-10m high. They are interstratified in soft white chalky limestones extending over several hundred km2 (see Figs. 2- 10 and 2-1 1). The adjacent limestones, however, are involved only in very mild undulations (Lees, 1928). The silica gel must have been still essentially plastic (externally brecciated only, like lava flows) and, thus, easily susceptible to sliding before the deposition of the next layers of chalky limestones. The movement was thus penecontemporaneous and corresponds in time perhaps to some of the early movements along the Jordan-Dead Sea Rift system. The gels seem to have been buried by a few meters of soft chalk, which provided support for the folds. One might place the motion in the earliest stage of syndiagenesis, say within some 100,000 years after burial. Penecontemporaneous slump structures, large and small, are features of almost all chert formations. Careful mapping of the Judean and Cretaceous cherts shows that the early

70 WEST

EAST

JUDEAN HILLS (Upper Cretaceous chaiks)

Fig. 2-10. Sketch of slump structures in Upper Cretaceous chert beds in the Judean Desert, east of Jerusalem. Sliding must have taken place during syndiagenesis while the silica gel was still completely plastic. (Sketch by the author.)

(slump) structures are oblique with reference to the principal folding (Steinitz and Kolodny, 1978). Chert-filled dikes transverse to the bedding point to the slow lithification of the silica (Steinitz, 1970). Microstructures of the syndi-

Fig. 2- I I . Slump structure in the Upper Cretaceous cherts, east of Jerusalem. Height of the scction is about 40 m. (Photograph: courtesy of G. Steinitz.)

71 agenetic phase vary according to horizon. Isotopic work discloses evidence of rapidly fluctuating, eustatic sea levels, introducing successively evaporitic and fresh-water conditions (Kolodny et al., 1980), i.e., the concept of “schizohalinity” (Folk and Siedlecka, 1974). Relict evaporite traces have been found to be widespread in the Senonian of Israel (Steinitz, 1977). One can turn now to a second example of a timing indicator. In eastern Denmark, south of Copenhagen at Stevns Klint and MQns Klint, the Upper Maastrichtian chalk is marked by parallel planes of scattered flints, i.e., parallel to the stratification (see Figs. 2-12 and 2-13). There are also vertical joints, and here and there the silica of the flints has flowed up along the joint

Fig. 2-12. Flints near the top of the Senonian chalk at Stevns Klint, Denmark, illustrate the slow dehydration of silica gel. The flints occur as concretions both parallel to bedding planes and merging (continuously) into vertical joint planes. Evidently, the gel was still mobile at the time when the chalk was sufficiently dehydrated to develop joints. (After Fairbridge, 1967b. fig. 6a, p. 62.)

72

Fig. 2- 13. Detail of Fig. 2- 12. Pencil= 15 cm. Apparently silica gel has migrated up and down along a joint plane in the chalk.

planes (so-called “flint curtains”, Rutten, 1957, p.433); there is no sign of brecciation and the flints of horizontal extension appear to be completely continuous with their vertical offshoots. One must conclude that the intrusion and dehydration of the silica gel occurred immediately subsequent to the diaclastic phase. Such jointing must result in loss of gas pressure, release of CO,, CH,, etc., and would accelerate dehydration (Sujkowski, 1958). The joints rise to the top of the chalk which is cut off by unconformable layers (“fish-clay” and Cerithium limestone) of Early Danian age (Rosenkrantz and Rasmussen, 1960). Evidently the jointing and syndiagenesis occurred ufter the general compaction and dewatering of the chalk but before the unconformity occurred. The timing of the flint dehydration may have been of the order of 1 . lo5 years. Complete dehydration of silica gels may be extremely slow, and Sujkowski

73 ( 1958) mentioned some “flints” encountered in deep bores that were still

rather soft. There was also a report of a vein of silica gel encountered during the construction of an Alpine tunnel. Nevertheless, the emplacement of the silica is usually an early phenomenon. Studies of the ‘80/160 isotope ratios in coexisting cherts and limestones suggest early diagenesis under similar marine environments (Degens and Epstein, 1962). In contrast to the above evidence of very late silica gel mobilization, there is the fact that penecontemporaneous flints are found eroded, reworked, and reincorporated in some conglomeratic beds in the chalk, as for example on an old buried anticlinal ridge joint west of Paris. Evidently these flints were already hard and subaerially exposed and, indeed, some show conchoidal fracture from the erosive period. Further there are examples of silicification preceding the penecontemporaneous dolomitization of many midwestern Silurian coral reefs (Dapples, 1959). What process could cause this apparently accelerated dehydration of silica gels to form penecontemporaneous hard flint? Rutten (1957, p. 436) suggested that on the shallow platforms temporary emergence could lead to the desiccation of the gels (which often enclose minutely perfect fossils of sponges, etc.). Cayeux ( 1941) in his “Causes anciennes ...” regretfully came to the conclusion that there were occasionally in geological time conditions which simply cannot be matched today, and chert formation, though formerly penecontemporaneous (as recognized by him in 1897), must be regarded today as non-actualistic. Rutten, though disagreeing in some other respects, concurred that this could be the case with these desiccating silica gels, i.e., that there is simply no comparable shallow carbonate platform today where this phenomenon might reasonably be expected to be operative. But there are some contemporary carbonate platforms-in the Bahamas, for example. West of Andros there is a broad bench of carbonate mud that is slowly drying out, ca. 60 cm above m.s.1. Radiocarbon dates show that the mud is about 2500 years old (the time of the “Abrolhos Submergence” that appears to be a eustatic and thus world-wide high sea level of up to 1.5 m above the present; Fairbridge, 1961, p. 169). Yet, unfortunately, this mud carries no desiccating flints. It is undoubtedly true that brief emergences of shallow platform environments are to be expected from the evidence of the eustatic theory and, indeed, the “hardgrounds”, “corrosion zones” or “discontinuity surfaces”, that are a characteristic feature of many neritic limestones, are now generally accepted as evidence of brief emergence (Weiss, 1958; Jaanussen, 1961 ; Bathurst, 1975; Wilson, 1975). In such cases, the epidiagenetic phase is temporarily applied, the mud becomes rapidly dehydrated, and sometimes even develops a karst crust with a trace of red soil (terra rossa). When

74 reburied, no further reactions are likely and the remaining syndiagenetic phase is locally short-circuited. Even briefer is the formation of tropical beachrock from calcareous sands of coral or molluscan debris. The beachrock may be loose one year and cemented the next. Within a few years it may become sufficiently recemented so as to ring to a hammer blow and yet may be sharply corroded by its intertidal exposure (Revelle and Emery, 1957). Rutten (1957) felt that in due course, somewhere, contemporary flint formation will be discovered, though he admitted that the present (postglacial, post-orogenic) time is abnormal in the light of historical geology. On the other hand, the evidence presented above suggests rather that one can find flint nodules only by coring to some tens of meters or more. It does not follow from the evidence of the reworked flint layers in the Upper Cretaceous that the flints had only just been formed “penecontemporaneously”, as Cayeux (1929) and Rutten (1957) argued. The chalk accumulated slowly, and the eustatic fluctuations or the revival of an old anticline could cause the loss of 10-20m of youthful sediment before the flint layers were reached. The chalk accumulation rate was between 0.1 and 1 mm per year. The level of flint diagenesis may represent a stage 10,000-100,000 years older than the contemporary sedimentation, and possibly as much as lo6 years (see discussion under Timing of Anadiagenesis). ( c ) Authigenic feldspar. The formation of authigenic feldspar is a phenome-

non that has long been known, but little understood, for it is easy to assume that the feldspars are high-temperature silicates. It now appears, however, that in some cases it may be little more than a leaching and dehydration event, probably in the anadiagenetic phase such as outlined by Kastner (1971). If one takes, for example, a hydromica (muscovite, or its degenerated form illite), it is not too difficu4t to forecast the loss of alumina and dehydration to form orthoclase. In simple empirical form it can be presented as follows:

K,O . 3 A1,0,

- 6 SiO, . 2 H,O + K,O . A1,0,

6 SiO,

+ 2 H,O + 2 A1,0,.

Weiss (1954) reported three horizons in the Ordovician shales of Minnesota, which are now predominantly authigenic orthoclase, but still show remnants of hydromica and montmorillonite, apparently due to feldspathization of a bentonitic volcanic ash. Even more common is authigenic albite which seems also to be usually derived from montmorillonite. In other words, the weathering hydrolysis of silicates is reversed. Some of this initial hydrolysis may even take place in the syndiagenetic phase, for the feldspars may not have been subjected to low pH conditions during mechanical weathering and transport in sea water having pH of 8 or more. In this

75 way, the marine realm tends to unify sediments and destroy traces of climatic extremes. Arid zones (extremely cold or hot deserts) furnish unweathered feldspars, but if feldspars come to rest in a rich organic ooze on the ocean floor, the acid weathering and hydrolysis that was denied them on land can be provided. Attention was first drawn to authigenic feldspars in the Cretaceous chalk of the Pans Basin (Cayeux, 1897, 1916), and they have been very generally reported elsewhere: in England (Reynolds, 1929), India (Spencer, 1929, and North America (Reed, 1928; Goldich, 1934; Gruner and Thiel, 1937; Willman, 1942). An exceptional case was observed by Daly (1917) in the Late Precambrian oolite of Glacier National Park, Montana, where up to 40% of the rock was replaced by authigenic orthoclase. Such massive feldspathization, however, is almost unique (Berg, 1952). Generally, the crystals occur isolated (as in chalk and limestone), as overgrowths on existing crystals (usually in sandstones), or finely disseminated (as in shales). They may even replace fossils (Stringham, 1940; Van Straaten, 1948). The most coherent paleogeographic model has been offered by Bryce and Friedman (1979, who showed that the authigenic K-feldspars of the Cambrian-Ordovician in the North Atlantic region were directly associated with the carbonate shelf regions landwards of the island arc, which then stretched from Spitsbergen and Norway through Scotland and Newfoundland through the Appalachians to Georgia. It is proposed that the authigenic feldspars here are all that is left of tephra that were remobilized to form overgrowths on pre-existing clastics. The transport of K-feldspars would have been favored by the high alkalinity of the weathering process in the Early Paleozoic pre-land plant environment. Zeolites were probably present, but, being unstable, could contribute to the feldspathization. The source of ions necessary for an independent authigenic growth of feldspar could come most readily from volcanic glass (Honess and Jeffries, 1940). There seems to be an inverse relationship to zeolites with time, i.e., as the latter decrease in frequency in older periods, the abundance of authigenic feldspars increases (Hay, 1966). Some almost pure K-feldspar beds in Cenozoic lake beds of western North America lack any trace of tuffaceous texture, although certainly they were derived from volcanic ash (Surdam and Parker, 1972; Sheppard and Gude, 1973). An interesting aspect of this dehydration and recrystallization phenomenon is that adsorbed cations, including various trace elements, are shed during the reorganization. Thus rubidium, a key element in certain geochemical studies, is lost in the diagenesis of illite to orthoclase (Horstman, 1957). Other familiar minerals, that are normally regarded as high-temperature forms, are known to be also authigenic. Rutile, brookite and unatase may be derived from biotite (also from ilmenite, etc.; Sun and Allen, 1957). Further,

76 tourmaline and zircon are known in the authigenic form (Boswell, 1933; Pettijohn, 1957), but little seems to have been done about investigating their origin. Other examples (sphene, etc.) were mentioned earlier. Natural chromatography

In a compacting sedimentary basin, there is a steady hydrostatic head which causes interstitial waters to rise upwards. These waters are for the most part connate, and hence, in a marine basin, essentially modified sea water. There may be some admixture of lacustrine or other fresh waters in a mixed marine-continental sequence. In volcanic regions, there is the possibility of some juvenile water being present, but such mixtures are regarded as quantitatively unimportant in most basins. All sedimentary formations are to some extent porous and permeable and additional permeability may be induced by diaclastic action. As the connate waters are progressively squeezed from the more deeply buried horizons they will pass vertically, and to some extent laterally (up-dip), into higher formations in which the chemical equilibrium may be completely different from that of the underlying sequences. Here one must consider the possibility of the operation of selective chemical filtering, that is a “selective adsorption process”. Laboratory use of this as a technique formerly involved the separation of colored substances in a fractionating column. Hence, the term “chromatography” developed which is now an important standard method of chemical analysis. The concept of “natural chromatography” has been raised in connection with the evolution of petroleum, and has been applied to sedimentary basins in general by Nagy (1960). The passage of a liquid, heavily charged with organic and inorganic solutions and colloids, through a porous membrane, is likely to cause a filtering of the larger molecules and ions of opposite charge to that of the membrane. Nagy devised a simple and adequate experiment to demonstrate natural chromatography through a quartz sand. Inasmuch as quartz carries very little charge, the separation is mostly mechanical, i.e., differential capillarity. In the case of clays-and most sediments carry at least some admixture of clays -the filtering can become mostly chemical. This is true especially in the case of montmorillonite which has strong negative charges. Bredehoeft et al. (1963) suggested that at first the passage of the negatively charged anions will be mechanically restricted and then the corresponding cations (Ca2+, M g 2 + ,Na’ and K + ) will be trapped. Deep basin brines may thus achieve a salinity up to six times that of ordinary sea water. In studying these ancient brines, Chave (1960) came to the conclusion that the variability in the nature of the sedimentary membranes was so great that a tremendous variety of

77 brines would result, and that no deductions could be drawn about the salinity of the ancient oceans. Similarly, wide variety is experienced with the petroleum hydrocarbons, which are exposed to every phase of diagenesis (Breger, 1960; Krejci-Graf, 1963a, b). In some basins, the solutions are nevertheless very weak, but the extended time factor may effectively permit reactions not otherwise easy to contemplate (Irving, 1892). The average sedimentary basin may accumulate over a period of 106-108 years, and by periodic revival (as noted in the section on Anadiagenesis) may. obtain a multi-tiered structure. The rising solutions, apart from mechanical filtering, will lead to various reactions which, today, can be deduced in part from fabric studies. The most obvious reactions are expressed by cementation and decementation phenomena. The cements are mainly CaCO, or S O , with Fe,O, playing a role in some special environments. As outlined earlier, the interplay between high and neutral pH (say 9-7) in the connate waters of normal (alkaline) marine sediments results in calcite or quartz cements. The acid waters required for the formation of Fe,O, are present in certain near-surfaGe bacterial environments in the initial stage of syndiagenesis. The acid oxygenated water can also be introduced in the epidiagenetic stage. Cementation, involving complete silicification, dolomitization, and dedolomitization introduces the special problem of sedimentary metasomatism. Diagenetic metasomatism The term metasomatism, coined by Naumann ( 1850), was applied essentially to the formation of pseudomorphs, either of the individual minerals or of whole rocks, involving a chemical replacement, atom by atom, but without change in form or volume, and obeying the “volume law” of W. Lindgren (see discussion by Holser, 1947). This process can be defined as a low-temperature enrichment of the sediment by new components “from without”, the original ions or molecules being removed in whole or in part. Essentially it is due to change in the chemistry of the enclosed or passing solutions. Metasomatism thus included several processes discussed already, such as oxidation-reduction, hydration-dehydration, etc., but it is convenient to consider under this heading the major ion exchange and replacement reactions. These normally obey the law of mass action and follow the entry of stronger solutions into the sediments. The usual site of such progressively increasing solution strengths is in compacting basins and, thus, normally occurs during the anadiagenetic phase, though sometimes in the epidiagenetic phase (e.g., phosphatization of limestone). Hypogene metasomatism may occur at high temperatures, as in some ore genesis, so that non-hydrothermal metasomatism should be designated as “diagenetic metasomatism”.

78 Some examples of diagenetic metasomatism may be considered briefly as follows. Dolomitization Even after countless research investigations, the understanding of the dolomitization phenomenon is still beset with problems (see Chapter 7 of Vol. I, entitled “Dolomites and Dolomitization”, by Chilingar et al.). Individual cases are often difficult to interpret because it occurs in all three phases of diagenesis. ( a ) Syndiagenesis. As pointed out by several authorities, for example, Carozzi (1975, p. 354), the history of the dolomite problem has gone through two stages of activity. The early studies, in the middle and late 19th century up till the 1920’s, disclosed the complexity of the problem (see review by Fairbridge, 1957). First, there was the non-actualistic development of dolomites: the “uniformitarian dilemma”, whereby their frequency increases in going back in time (Zenger, 1972a, b). Secondly, they have proved singularly difficult to synthesize in the laboratory under reasonable earthsurface conditions. The second phase of the dolomite problem’s history began in the mid-20th century with introduction of the X-ray technique and other sophisticated laboratory equipment, coupled with extensive field work on actualistic situations. The first critical discovery was that of “protodolomite”, a disordered calcium-rich dolomite, synthesized under “reasonable” earthsurface conditions by Graf and Goldsmith (1956). High-magnesium calcites of organogenic origin had been found in shallow-marine situations by Chave (1954); like aragonite, they are metastable over relatively brief periods of geologic time. Chave (1954) and Chilingar (1953, 1962) showed that the higher the environmental temperature, the higher the content of biogenic magnesium in the calcites. Here was the potential for syndiagenetic dolomitization. Then, in the 1960’s came a succession of spectacular discoveries in shallow-water lagoons and intertidal mudflats (“sabkhas” in Arabic) of the warm subtropics, as predicted by Chilingar (1956d). Wade (191 1) had reported dolomite rhombs in the Holocene lagoonal muds on the Red Sea coast of Egypt. Intertidal magnesium-enrichment in tropical seas (notably beachrock) was identified by Reuling ( 1934) as “Gezeiten-dolomi tisierung” (intertidal dolomitization). Sir Douglas Mawson (1929) found them in the Coorong lagoon of South Australia, where he later showed them to the present writer. Here the study was taken up again by Alderman and Skinner (1957) and further by Skinner (1963). A comparable lagoonal situation was shown to the writer by R.A. Bramkamp in 1959, in the Persian Gulf south of

79 Dhahran, where gypsum was observed actively precipitating with what proved to be protodolomite. Another lagoon setting, at Bonaire in the Dutch West Indies, was recognized by Deffeyes et al. (1965). On the supratidal mudflats (sabkhas) of the Persian Gulf, the seasonal reversal of wind systems alternately wets and dries the littoral zone. A much wider zone has been alternately flooded and dried by small eustatic oscillations (+/- 3m) during the Holocene in 500-1000-yr cycles (Fairbridge, 1961). A revolutionary study by Illing et a]. (1965) described the tiny rhombs of protodolomite that replace the original aragonite and gypsum when exposed to magnesium-rich brines. Under high temperature (> 40°C) and capillary rise, the pore water salinity rises to 6-8 times that of sea water and the Mg/Ca ratio rises to over 10, whereas the pH remains low (ca. 6.7) because of biogenic activity. Radiocarbon dating of sediments at depths of 20-30 cm gives 2000-3000 B.P. ages. Valuable reviews of the Persian Gulf carbonate environments have been presented by Bathurst (1975) and Wilson (1975); a special volume has been edited by Purser (1973), and the “Benchmark” papers have been collected by Grkland and Evans (1973). A somewhat comparable setting was established in the intertidal muds west of Andros Island in the Bahamas (Shinn et al., 1965). It seems that 1965 was the “vintage year” for protodolomite. Whereas most investigators believe that the Holocene protodolomites are syndiagenetic, the development of penecontemporaneous dolomites in ancient formations tends to be controversial. Sedimentational evidence often suggests that the dolomites are indeed penecontemporaneous, but the petrologic evidence often indicates complex subsequent histories of recrystallization or metasomatism. The geochemical setting whereby dolomite could develop from sea water during gypsum formation (i.e., in partial isolation) due to the rise in Mg/Ca ratio was recognized as a theoretical model already by Hunt (1859), but experiments with mixed carbonates required artificial heating to generate dolomite. With progressive concentration, the solubility curves of CaCO, and MgCO, intersect, and at this point CaSO, also reaches saturation (Teodorovich, 1955), but still this does not help to explain non-evaporative situations. In spite of much experimental work, however, it was not until Chave (1954) showed the biogenic origin and metastable nature of the high-Mg calcites that the dilemma was overcome. The principal Mg-calcite secreting organisms are the calcareous algae, which because of their photosynthetic requirements and preference for warm water, are characteristic of shallow tropical seas and tidal flats. As pointed out by the present writer (Fairbridge, 1957, p. 154), there is abundant evidence associating many ancient dolomites with warm, shallow seas. The diachronism of the Mississippian Greenbrier Dolomite (Rit-

80 TOP

BED N O I

EXPLANATION MgCO,

0CaCo3 INSOL

Fig. 2-14. Vertical section, with MgCO, and CaCO, analyses, reduced to beds of equal proportions, of a sequence of Lower Ordovician rocks west of Harrisburg, Pa. (After Fairbridge, 1957.) There is rhythmic alternation of dolomitized and nondolomitic limestones, the dolomitization coinciding always with the beds containing maximum terrigenous (insoluble) material, and therefore presumably closer to the original coastline. (Analytic data by Lesley, 1879.)

tenhouse, 1949) also demonstrated “a very precise relationship to the shoreline”. Most dramatic are the 1879 analyses by Leslie of an alternating series of Lower Ordovician limestones and dolomites from near Harrisburg, Pa., apparently a eustatic cycle, which show that the terrigenous components are always highest in the dolomitic beds (Fig. 2-14). Sarin (1962) has proposed a cyclic killing of algae in the limestone facies by supersaline magnesium-rich currents. In any case, the dolomite here is a near-shore indicator. An interesting, consequence of these conclusions was the reversal of some contemporary oil exploration strategy: previously, the dolomites were often assumed to be the deep-water indicators. Others who had earlier pointed out the warm, shallow-water characteristics of various dolomites included: Skeats (South Pacific, 1903; Tyrol, 1905), Dixon (Britain, 1907), Van Tuyl (Iowa, 1916, 1918), Steidtmann (Wisconsin, 1917), and Fondeur et al. (France, 1954). In the widespread epicontinental platform seas of the late Precambrian

81 and Paleozoic, there seem to have been times of almost universal dolomite production. It has been customary to refer to such rocks as “primary dolomites” on the assumption that a primary crystallization would be favored by a marine hydrology that was somewhat different from that of today. There seems little doubt that the atmospheric oxygen level was low, and the p C 0 , was appreciably higher than today (Chilingar, 1956~;Strakhov, 1967, 1969; Fairbridge, 1964, 1967a). Experimental work by Baron ( 1960) demonstrated the possibility of direct precipitation of dolomite. A forceful argument, however, against a primary dolomitization environment in the Late Precambrian is the fact that there were also widespread limestones that were quite low in magnesium. Clearly there were regional patterns. An important paleogeographic factor is that throughout the period in question the equator transected North America, where much of the evidence is found. Whether or not these ancient dolomites underwent a rapid syndiagenetic transition, is still an unsolved problem (see, for example, Carozzi, 1975, p. 357). What is certain, however, is that they are commonly associated with evidence of photosynthetic algae, so that vast amounts of high-magnesium calcites could be generated that would then be susceptible to penecontemporaneous dolomitization, i.e., syndiagenesis. Isotopic data appear to point to this interpretation (Degens and Epstein, 1962). In certain varieties of calcareous algae, the MgCO, in calcite reaches a maximum of 25-30%. Inasmuch as this is a metastable condition, within a period of time (10’-104 years) it is likely to invert to pure calcite, whereas in the presence of high Mg2+ concentration dolomite will develop. Probably this is true metasomatic replacement of Ca2+ by Mg2+ (see, for example, photographs of calcareous algae from an Eniwetok Atoll bore, with euhedral dolomite rhombs beginning to form in the middle, in Schlanger, 1957). Numerous syndiagenetic limestones and dolomites are so tightly crystallized that they still retain the fetid odor of decaying organic material, which suggests rather rapid diagenesis (Lucas, 1952). The presence of specks of iron oxides, pyrite, marcasite, or ankerite has often been noticed in dolomites; in fact, the Fe/Mg ratio is virtually constant in dolomite (Cooper, 1954). This iron concentration seems to call for a reducing subsurface condition and reasonably vigorous benthonic biota (Moretti, 1957), not likely to be found under evaporite conditions. Syndiagenetic dolomite, as isolated rhombs, occurs also in the deep-sea today, in depths of 3000-4000m (Boggild, 1912, 1916; Correns, 1939; Zen, 1960). It has been suggested by Weynschenk (195 la, 195Ib) that massive and extensive deep-sea dolomitization was possible, but the D.S.D.P. “Glomar Challenger”-drilling has not disclosed any examples. For the scattered dolomite rhombs it would seem that in some cases transported algal nuclei of

82

high-Mg calcite may provide the source. Bonatti (1966) envisioned an igneous association, whereas Davies and Supko ( 1973) discussed other suggestions. Dolomite is also forming today in the Atlantis Deep of the Red Sea, at a depth of over 2000m, but in this instance it is associated with hot metal-rich brine seepages which are properly classified as hydrothermal. ( b ) Anadiagenetic dolomite. Anadiagenetic dolomite is by definition a secondary dolomite, but belonging to a later stage than those listed above. In any metasomatic replacement of this sort, a high concentration of Mg2+ is a prerequisite. It has been suggested by several workers that a refluxing of dense brines from the saline lagoons would lead to an ideal setting (Illing, 1959; Adams and Rhodes, 1960). The hydrodynamic motion can be lateral as well as vertical. In the former, the flux could be from a saline lagoon bounded by a barrier reef; in the Paleozoic sections the barrier reef limestones are almost always dolomitized. Mennig and Vatan (1959) found a similar barrier reef correlation in the Devonian of the Ardennes; they spoke of “epigenesis”, but the present writer suspects that this,is merely a different use of the term. It seems likely that this metasomatism is not achieved without some elevation of pressure. Progressive dolomitization results in ca. 12% increase in porosity due to reduction in volume (see detailed discussion in Chapter 5 of VOI. I). Whereas it is observed on a massive scale below a depth of 196m in Funafuti Atoll corresponding to 20 atm (Schmalz, 1956; Fairbridge, 1957, p. 148), such metasomatism is absent from modern stable shelf coral limestones, such as those of the Great Barrier Reef where the base does not exceed 7 atm pressure today (Fairbridge, 1950a). Pressure is not a controlling factor, however, for the Bikini-Eniwetok bores are only partially affected. It is suspected that it is the primary distribution of the high-magnesian calcite nuclei that plays the vital role (Fig. 2-15). Where the results of tectonism have been added to those of simple compaction, there are extensive fault zones and joints which permit more thorough access by rising waters. On either side of a fault zone there is often a “Christmas-tree” effect produced by the rise and lateral spread (along more porous zones) of dolomitizing waters. Across the Paleozoic of the Midwest it is often noticed that dolomitization is more prominent as one approaches the tectonic belts. In the great evaporite basins of the world, there is often an interstratification of anhydrite-dolomite beds with the salts (halite, etc.). A clear distinction is usually made between these evaporite basins and the littoral evaporites such as those of the lagoons and sabkhas of the Persian Gulf and North Africa. With the progressive loading, with/or without some tectonic defor-

83 WATER PRESSURE DEPTH TEMP 28*C AM

s100-

4-

3200-

3002,

I I

l

I

l

I

Fig. 2-15. Diagram to illustrate, in an idealized atoll, the suspected relationship of primary Mg-rich calcite (largely fixed by algae) to contemporary dolomitization, which may take place beneath the lagoon floor at a depth of 600-700 ft, where there is a pressure of about 20 atm. a temperature of at least 10°C (possibly raised by organic activity), a free access of Mg2+ ions from sea water, but a high alkalinity (pH about 9-10) due to reducing conditions brought about by organic activity both in the atoll walls and in the sediments themselves. Below the level of contemporary dolomitization, the atoll core may have been dolomitized discontinuously, the process being affected by changes in depth (and pressure) due to eustatic oscillations of sea level and varying rates of atoll subsidence. Periods of stability permit complete dolomitization to be achieved, while rapid subsidence permits a quick buildup of sediment, partly sealing off the former zone of dolomitization, so that completion is not achieved (e.g., 5-20% MgCO, in some of the Funafuti analyses). Funafuti atoll stands as the prototype for this scheme; many other Tertiary dolomitized atolls in the Pacific and Indian oceans have been partially elevated, so that the critical top 600 ft have been removed by erosion. Nondolomitized atolls, such as Bikini and Eniwetok may, for some reason or other, have failed to provide the necessary physicochemical requirements, such as the reducing conditions (owing to freer circulation) or may have subsided too rapidly. (After Fairbridge, 1957.)

mation, there ensues within the halite-potash type salts what is known as a “geothermal metamorphism” (Borchert and Muir, 1964). This takes place in fact within the zone of normal anadiagenesis, but is somewhat more dramatic in its changes because of the low stability range of the various salt minerals. Whereas some of the dolomitization is recognizably syndiagenetic, resulting from downward percolation of brines from the lagoon floor, there is some replacement of dolomite by anhydrite (Stewart, 1965). DuriPg deeper burial, there is further dolomitization, now anadiagenetic and associated with ascending solutions, and with rising temperature and pressure. Lateral flux of saline solutions in massive anadiagenetic dolomitization has been mentioned by Hite (1970), in addition to the reflux seepage model of Adams and Rhodes ( 1960).

84 ( c ) Epidiugenetic dolomite. Epidiagenetic dolomite is the superficial phase of

secondary dolomitic metasomatism. Such strictly epigene dolomites are not nearly so common as the anadiagenetic ones, for the simple reason that descending waters are generally weak solutions and require some sort of enrichment. This may be achieved near the lower boundary of the soil profile (and here epidiagenesis impinges on weathering), resulting in the formation of an unconformable, dolomitized hard-pan horizon. Such a horizon has developed, for example, in the Eifel district, Germany, where folded Devonian dolomites and limestones are now capped by this epidiagenetic dolomite, without any respect for what was the original lithology (Fairbridge, 1957, fig. 13; see Fig. 2-16). I t is thus sometimes known as “subsequent dolomitization” (Hatch et al., 1938, p. 193). Dolomitization by ground water (i.e., of meteoric origin) was considered by Steidtmann (1911) and some other early workers, but not taken very seriously. In most situations, insufficient volumes of magnesium solutions

Fig. 2-16. Four generations of dolomite in the Eifel district of West Germany (adapted by Fairbridge, 1957, from work of Udluft, Reuling, and others). Idealized section showing the Devonian sediments, consisting of noncalcareous Lower Devonian followed by the Middle Devonian Couvinian (low-magnesian) limestone and Givetian (penecontemporaneous) dolomite ( 1 ); these were broadly folded in Pennsylvanian-Permian times and peneplaned. During Permian soil formation, magnesium-saturated ground water (exposed to weathering of the Givetian dolomite) led to dolomitization of the bevelled surface layer of the Couvinian limestones, as a continental “hardpan” ( 2 ) . Uplift led to some dissection of the peneplain, and locally the hardpan was stripped to expose nondolomitized Couvinian limestones. With the progressive invasion of the Triassic sea, sands (Bunter sandstone) were provided with the Middle Devonian limestone boulders (now red from terra rossa soil formation), and Couvinian dolomite boulders ( S ) , thus by the two earlier generations of dolomite. Penecontemporaneous and subsequent faulting, associated with further uplift, allowed hydrothermal solutions in places to introduce a fourth generation of dolomite ( 4 ) . locally accompanied by ore formation.

85 would be available to cause large-scale dolomitization. This objection is overcome at the fresh-water-salt-water mixing zone in porous littoral formations, such as exist in Florida and Yucatan (Hanshaw et al., 1971), in Jamaica (Land, 1973), in Israel (Magaritz et al., 1980), and elsewhere. Although Holocene nodules and hardpans have been found in such settings, no massive epidiagenetic dolomitization has yet been identified. Nevertheless, an attempt to explain the Middle Ordovician dolomites of Wisconsin in terms of a fresh-water-salt-water mixing zone has been made by Badiozamani ( 1973). This so-called “Dorag” model postulates an early interruption of syndiagenesis by epidiagenesis, which can be instituted in a cratonic setting of this sort by eustatic fluctuations. Dedolomitization Dedolomitization is generally a metasomatic replacement of dolomite by calcite; for this, Smit and Swett (1969) prefer to use the term “calcitization”. I t may also result simply from leaching, with a molecule by molecule solution followed by a selective removal of the MgCO,.‘The result is a loose sandy-textured calcite. It seems likely that this phenomenon goes hand in hand with the production of the epidiagenetic dolomite noted above. Where dolomite and gypsum have been interbedded, the dolomite is often fractured and etched (mechanical collapse and solution). Von Morlot (1847), who first applied the term dedolomitization, suggested the following reaction: CaMg(CO, ),

+ CaSO, . 2 H ,O * 2 CaCO, + MgSO, + 2 H 2O

Sander (1951) has stated that in the Northern Limestone Alps the calcitization of dolomites is really more important today than the original dolomitization. Shearman et al. (1961) have described the dedolomitization textures in the French Jura. Hydrothermal solutions may also play a role in dedolomitization (Faust, 1949). A second reaction, possible where MgSO, dominates the brines, may be given as: CaMg(CO,),

+ MgSO, * CaSO, + 2 MgCO,

In this reaction anhydrite is the remaining solid phase; crystals of anhydrite have been observed, in the Permo-Triassic evaporites of England, “eating” into dolomite (Stewart, 1949). Multiple reversals are favored in the dolomitization-dedolomitization of a region subject to a complex geologic history. Walker (1962) pointed out the reversible nature of chert-carbonate replacements. In the Ordovician of Wisconsin he discovered chert nodules replaced by dolomite, but at a later stage the process was reversed as evidenced by chert pseudomorphs of dolomite rhombs.

86 Silicification Inasmuch as silica is relatively insoluble at normal temperatures in waters having pH of less than 9 (Correns, 1949; Krauskopf, 1959), it is rarely mobilized after dehydration of the primary gels in the syn- and early anadiagenetic phases (see Fig. 2-7). Even low concentrations of dissolved SiO, will be precipitated at the low pH (3-6) existing in some lakes, leading to silicification of wood, as CO, is liberated by bacterial action (Correns, 1949), but such conditions can rarely be achieved in oceanic sediments. With progressive compaction and concentration of the connate waters during advanced anadiagenesis, however, sharp contrasts in pH may be obtained. According to calculations by Siever (1957b), equilibrium with the ZaCO, is reached at pH of 9.8 at 25OC, particularly in a thin film of water containing originally atmospheric CO, between grains of quartz and calcite. At this point silica dissolves to the limit of the very restricted water film. Pitting (etching) of quartz grains is often observed (Dapples, 1959). In quartz sands the usual products of anadiagenesis are eventually quartz overgrowths and quartz cements, which result in quartzose sandstones (or “orthoquartzites”). According to Krauskopf ( 1959), the first dissociation constant of H,SiO, occurs at pH of 9.8 and, therefore, the often quoted solubility curve of C.W. Correns (in: Barth et al., 1939, p. 129) may be misleading. Siever (1959) pointed out, however, that at a pH of 9.8 partial dissociation of H,SiO, produces silicate ions and essentially doubles the quartz solubility at lower pH values (Fig. 2-17). An additional factor is the increased polymerization 12

r

Fe3++ I(FeOHIz4

PH

Fig. 2-17. Solubility of silica gel, aluminum hydroxide and ferric iron hydroxide in relationship to pH. (After Correns, 1939.) Solid line indicates a modified SiO, solubility, following the work of Krauskopf (1959). Broken line indicates original curve of Correns.

87 rate produced by dissolved electrolytes, as found in many deep connate waters. Precipitation, therefore, varies with the ionic strength of the ground water, and solution is particularly favored by rising temperatures (Siever, 1959), as well as by conditions which raise the pH. Hydrothermal waters may assist such temperature rise, and silicification is a marked feature of deep-seated faults. Silicification involves the cementation, or complete replacement of sedimentary particles or fossil remains, by silica, usually in the sequence opalchalcedony-quartz, ‘the end-point requiring lo7 yr or more. The process is often included in the term “induration”, which implies general hardening of the rock. During weathering under seasonally contrasting climates, the capillary transport of silica to the surface of exposed rocks results in reprecipitation, producing a siliceous rind, or “case-hardening”. Far-reaching silicification is even known in certain coal formations. Whereas normally the low pH associated with the syndiagenesis of coal would inhibit such silicification, a late anadiagenesis with rising waters of high pH would favor it. In heavily folded parts of the Ruhr Carboniferous, nodules, lenses, and massive replacements by silica are observed (Hoehne, 1957). When the silica is removed molecule by molecule by solutions of high pH, it will tend to be replaced by the least soluble components in such solutions at a given pH. These are generally the carbonates. This progressive metasomatism of chert is described by Walker (1962), who has further demonstrated multiple replacement reversals. This silicification-desilicification sequence seems to be explicable only by proposing an alternating passage of waters of higher pH (over 9) and lower pH (under 9). Such conditions can be visualized during the gradual dewatering of a compacting basin during anadiagenesis, when successively new connate water sources are liberated from their cement traps by jointing and faulting as the compactional and diastrophic evolution progresses. The same phenomenon may continue during epidiagenesis, but with additional sources of water of lower pH. It is under these conditions that pyrite and marcasite are most liable to become oxidized, sharply dropping the pH and raising the temperature. Silica solubility rises with higher temperature, whereas the reverse is true of calcite (Okamotu et al., 1957). Quartz-calcite intergrowths occur often (Dapples, 1959; Walker, 1962). As epidiagenesis proceeds, there is a gradual tendency towards stabilization in the intermediate pH range, where calcite is still soluble and silica is completely stable. Thus, on old land masses, stable tectonically for a long period of time, there is a widespread silicification of limestones. Where, as in Africa, Australia, Peninsular India, and the Brazilian Shield, the climatic record of the last lo8 years or so has been tropical or subtropical, there have

88 been immense and almost continuous supplies of SiO, passing into the vadose water system by leaching down from lateritized soils. Under such conditions, the epidiagenetic stage is one of massive silicification. Solid limestones often become completely silicified down to some hundreds of meters below the weathering zone (Fairbridge, 1950b). It is curious that the Collenia type calcareous algae (stromatolites) of the Precambrian in those regions long went unrecognized, because they were found only in what appeared to be primary quartzites. A widespread phenomenon of epidiagenetic silicification of sands is the development of silcrete, the siliceous duricrust, or hardpan, of the semi-arid tropics and subtropics (Dury, 1974; Langford-Smith, 1978). A comparable argillaceous rock in this same way is converted to a porcellanite (also spelled “porcelanite”), which has the conchoidal fracture of unglazed porcelain. The environmental requirement for epidiagenetic silicification is a highly contrasting, seasonal climate fluctuation, whereby the silica is mobilized during the (warm) wet season and then reprecipitated during capillary lift during the hot dry season (Fairbridge, 1975). It is curious that, whereas the silcrete crusts of the present tropics and subtropics of the Southern Hemisphere are usually only 1-10 m thick, the massive silicification of many porous sandstone formations in the Northern Hemisphere may extend to a depth of 200-500m. Thus, for example, most of the Paleozoic quartz sandstone and conglomerate formations of the Appalachians are porous at depth (acting as oil and gas reservoirs); nevertheless, in outcrop, they are massively silicified, the induration having blocked all porosity. The explanation appears to lie in the fact that in plate tectonic history the Northern Hemisphere continents have shifted polewards since the Mesozoic, progressively traversing tropical to temperate climate belts and undergoing extensive differential uplifts. This amplified the entry of meteoric waters and promoted deep-working epidiagenesis. An interesting phenomenon that has received notice is what in the Paris Basin is called “meulerization” from the French word for millstone (Termier and Termier, 1963, p. 345). It is believed to be analogous to the formation of siliceous hardpan or “silcrete” and is observed in progress in South Africa and the northern Sahara today (Alimen, 1958). Carbonate sands or limestones are locally cemented and partly replaced by opaline or chalcedonic silica; exceptionally, gypsum may also be replaced. The geomorphologic history of the Paris Basin discloses a Late Tertiary subtropical peneplanation, during which small ephemeral-lake depressions appear to have determined the sites of the “meulerization”. An opposite reaction is sometimes observed during the epidiagenetic phase in temperate latitudes. This has been called “Fontainebleau sandstone crystallization”, to describe the local formation of a calcite cement in a

89

Fig. 2-18. Desilicification of quartz sand grains. aided by the base exchange of intercalated clays (after E. Thomson. 1959). in four stages: A. Nature of quartz and clay shortly after deposition. Clay is present between some grains and not others. B. Base exchange of K by Ca2+ and Mg’+ begins along edges of sheets. C. Exchange proceeds rapidly along cleavage surfaces. K,CO, is formed and a higher pH develops in clay-rich regions. D. As pressures increase. Si4+ dissolved in regions of high pH migrates to regions of low p H and precipitates.

desilicified quartz sandstone (or uncemented quartz sand). The cement is continuous and fracture surfaces of the rock show (by their reflection) that the calcite is in crystallographic continuity. Like meulerization the phenomenon was first described from the Paris Basin. One may suggest that this has been a late diagenetic phenomenon dating from the late periglacial (cold-wet) phases when acid ground waters were generated by podzolization. Striking examples of desilicification may be observed in quartz-sand-clay mixtures, such as often found in graywackes (see Fig. 2-18). Thin sections may disclose the growth of newly-formed illite, eating into the quartz grains. The same phenomenon is associated with stylolite formation. This is unlikely to be simply pressure solution as suggested by many authors (e.g., Fairbairn, 1950; Heald, 19551, but is greatly facilitated by a minor clay fraction (Thomson, 1959). Stylolites in sandstones or quartzites have long been an enigma, but Thomson showed that a clay layer between sand grains would liberate K,CO, if subjected to C0,-rich waters. K,CO, is a strong alkali, which would mobilize SiO, at the clay contact, only to be reprecipitated nearby in the generally acid solution.

90 Phosphatization Another example of diagenetic metasomatism that has provided geologists with considerable problems has been the phenomenon of phosphorization. Generally, phosphatic acids replace the carbonates in limestones. The problems associated with phosphatization are quite as complex as those connected with dolomitization. Phosphorites (a collective name for phosphatic rocks) appear to be of both biogenic and inorganic origin. Most biogenic phosphorus 'is part of a food chain that begins with unicellular marine organisms, which pass to fish and then sea-birds. Droppings of the latter accumulate as guano. Leaching of this guano into coral or other porous carbonate substrate generates an epidiagenetic phosphorite. Phosphorus is usually introduced (in reactable form) into the sedimentary cycle at the present time by organic agencies. It is present in nucleic acids formed in all living matter. It is contained in many proteins (phosphoproteins), many lipids, and many carbohydrates. Its abundance is rather low in sea water (about 0.07 mg/l), so that the limiting factor in biogenic phosphatization is the local concentration by organic metabolism. Organisms employ inorganic phosphate to synthesize ADP (the diphosphate) or ATP (triphosphate), which in turn provide fundamental organic energy sources. The present-day inorganic phosphorites are mainly limited to nodular concretions of carbonate fluorapatite or a phosphormicrite, which show discontinuous but penecontemporaneous accumulations. Some are inherited (lag) deposits. Deposits are associated with shelf margins from 30- 1000m depth, off California and Mexico, Peru, Chile, northwestern and eastern Australia, northwest Africa, southwest Africa (Namibia), South Africa, Oman, and Somalia. The phosphatization of benthic foraminifera was reported by Manheim et al. (1975). From almost all periods of geological time, the sedimentary phosphates are of marine origin. Although the secondary terrestrial concentration often occurs, the initial segregation is marine. Certain periods seem to be more favorable than others, e.g., Cambrian, Permian, Upper Jurassic, Lower Cretaceous, Upper Cretaceous, and Tertiary (Gimmelfarb, 1956); these are essentially transgressive, thalassocratic stages. Such stages are associated with the expansion of broad epeiric shelves, which are the preferred environments for photosynthetic algae. Among these, the Chlorophyceae (green algae) are the principal organisms to accumulate calcium and phosphorus (Demolon and Boischot, 1948). The open-shelf phosphatites are generally nodular or concretionary, and are often associated with coprolites and glauconite, the geochemical evolution of which gives a clue to the phosphatizing environment (Visse, 1953; Riviere and Visse, 1954). In subsiding basins marked by a higher accumulation rate and probably by somewhat greater depths and poor circulation, the

91 phosphatites tend to be bedded and dark (even black) in color. Here they are commonly accompanied by pyrite, reflecting the reducing environment of the syndiagenetic phase. The pyrite, in turn, is often oxidized to gypsum (anadiagenetic). Besides continental margins, some estuaries carry up to 1-ppm phosphate, so that a site parallel to an old, deeply eroded orogenic belt, such as the Appalachian Piedmont, could also be favorable. To explain the nodular and pelletal phosphorites of the geologic record, Cook (1976) suggests the following criteria: (a) influx of phosphate-rich water to a warm shallow marine shelf basin or margin relatively free from important terrigenous sedimentation; (b) development of a rich neritic biota; (c) existence of an anoxic benthic environment, which inhibits predation and bottom scavengers and, thus, permits the accumulation of a, phosphatic mud (during the early syndiagenesis phase there is a loss of C, N and H). In this way formation of an organogenic ooze is postulated; the possibility of purely inorganic precipitation, perhaps around organic nuclei, remains open. The next step is phosphatization under the low pH-high alkalinity conditions of early burial in the syndiagenetic phase. Leaching of the phosphaterich solutions leads to metasomatism of carbonate fossils, fecal pellets, bones, coprolites, shark teeth, and oolites by apatite. Even a siliceous ooze can be phosphatized. Bottom currents and wave action lead to periodic winnowing of the sediments, concentrating phosphatic nodules and pellets as lag deposits. The dilemma of phosphates, in general, is their apparently non-actualistic development in the geologic record. At certain times vast concentrations occurred, but the Present is not one of them. In the Late Cenozoic, sedimentary phosphates appear to be restricted to marginal-shelf regions, marked by strong upwelling, mainly in the trade-wind belts between 40°N and 40"s. This theory was proposed by the Soviet worker A.V. Kazakov in 1937 (see review by Blatt et al., 1980). For earlier periods, plate reconstructions and paleomagnetically established paleolatitudes suggest similar distributions (Cook and McElhinny, 1979). The saturation of PO:- in sea water is greatest near the shelf margin (0.3-0.8 ppm) in upwelling areas of the trade-wind belts. At shallower depths it is reduced by algal metabolism, whereas at greater depths it is reduced by the higher pC0,. Massive phosphate precipitation, however, would also require a reduction of the combined nitrogen level due to denitrification by bacteria, according to Piper and Codispoti (1975). This could have occurred during warmer times of the geologic past when there was expansion of the oxygen minimum layer (0, content of less than 0.1 ml/liter), which today is generally located below 200 m, at depths marked by higher pC0,. The Quaternary, in general, is an unfavorable period for this scenario.

92 Thus the stage for large-scale phosphatization is preset in geological time by certain geotectonic-paleogeographic events, that do not prevail at the present stage in history. The extreme and large-scaled oscillations of sea level during the Quaternary, however, favor a special epidiagenetic environment that is found in isolated atolls. These offer sites for sea-bird sanctuaries, free from predators until recently introduced by man. Deposits of phosphate-rich guano accumulated here. During the high sea-level stages, the islands were small and leachng of the soluble organic phosphates by rain water brought them quickly back into the sea. During the low sea-level episodes, however, the atolls became emerged and resembled mediaeval castles with limestone walls and floors. This was an ideal setting for leaching of phosphoric acids into freshly formed limestone (often composed of porous metastable aragonitic corals), to form, as a rule, collophane (collophanite), Ca,P20, . H 2 0 , or dahllite, Ca,(PO,), - CaCO, . $ H20. On small volcanic islands with a barrier or fringing reef, the incomplete atolls (e.g., Navassa Island, West Indies), or along semiarid mainland coasts (e.g., southwestern Africa, northern Chile), the same thing may happen. In such cases, however, the igneous weathering products may interact during the diagenesis, resulting in vivianite, Fe,P20, * 8 H 2 0 , or wavellite, 4 AIPO, 2 Al(OH), * 9 H 2 0 . The geochemistry of phosphatization of atolls suggests a low pH in a fresh-water-saturated environment, without any unusual temperature or pressure requirements. Teodorovich ( 1954) believed that weakly reducing or neutral environments favor phosphatization. The widespread stratified phosphatites of the past, being so largely marine, might suggest a different setting. It seems likely that it is the marine condition that is essential only for the organic segregation, which, if followed by a small negative eustatic swing over a shallow shelf region, would lay it bare for the phosphorus mobilization under rain-water leaching. Such shelves are equally favorable for the accumulation of marine carbonate sediments, which provide the necessary “host” rocks. On warm shelves, the bulk of the host material is likely to be in the form of aragonite, which is particularly soluble in fresh water of only slightly reduced pH (6.5-7.5). Such processes on exposed atolls or shelves would both be examples of epidiagenesis, where the syn- and anadiagenetic stages have been rapidly bypassed by a sudden drop of sea level. The duration of such negative eustatic stages during geological history may have been of the order of 5000-50,000 years, which is the upper limit of the time required for phosphatization of quite thick formations. In some atolls, dolomitization has preceded phosphatization, probably during an earlier cycle involving subsidence and re-emergence. Whereas some islands have become phosphatized under eustatic control alone, the

93 majority show tectonic movement as well. This seems essential for deep phosphatization. A paragenetic relationship between phosphatization and the accumulation of concretionary silica in the Permian Phosphorite Formation of western U.S.A. has been suggested by McKelvey et al. (1953). Studies in Morocco, however, show that while bedded cherts there commonly succeed phosphatic sands in the sequence, this does not imply an interrelationship (Salvan, 1955). Sideritization Sideritization is a fifth type of sedimentary metasomatic diagenesis. Primary precipitation as siderite, FeCO,, probably does not occur in open marine environments today. In swamps and other restricted basins, however, the ferric oxide hydrosols of river waters would be reduced to the ferrous state, and removal of CO, by photosynthesis of plants would dissociate the bicarbonate ions to cause the direct precipitation of FeCO,. Some sideritic iron ores are thus associated with some coal deposits. Teodorovich (1949, 1961) claimed also that there were at times distinct “siderite facies” of marine environments of the past, marked by reducing conditions and by strong fluctuations of pH and Eh. Much more usual, however, seems to be the condition where marine limestones are metasomatically replaced by siderite during syndiagenesis (Cayeux, 1916; Hatch et al., 1938, p. 135). The CaCO, is often in a highly porous state, such as oolites or organogenic calcarenites (often crinoidal), and may also be aragonitic. Thus a pervious and metastable “host” is provided, just as is necessary for dolomitization and phosphatization. In the same way that calcitic fossils (brachiopods, bryozoans and certain molluscs) in a matrix of carbonate material, believed to have been originally aragonitic or, equally well, metastable organogenic high-magnesian calcite, are often “spared” by the dolomitization, so too they are often found preserved in a matrix totally replaced by siderite. This fact alone is strong evidence for a syndiagenetic origin of such siderites, although they are commonly stated to be epigenetic (e.g., Twenhofel, 1950, p. 431). On entering the ocean, ferric oxide hydrosols are electrolytically flocculated and thus deposition is likely to occur on continental shelves rather than in the deep sea (Moore and Maynard, 1929, p. 507). On semi-tropical shelves one may thus expect the ideal environment. Iron is readily leached out of the laterite soils on the land and brought down to the oceans by rivers. In the warm ocean there is steady concentration of organogenic carbonates, largely in metastable form. It is not surprising, therefore, that the great sedimentary iron ores are, like the phosphates, normally concentrated in the more transgressive marine stages (Ordovician, Silurian, Jurassic, Lower Cretaceous). The liberation of large quantities of iron in the periods immediately

94 preceding orogenic episodes was stressed by Cayeux, but the role of biorhexistatic soil cycles in concentrating the iron should not be forgotten (Erhart, 1956, 1973a, b). Siderite is also found as a scattered authigenic mineral in some formations, but this is rather unusual. The sideritic limestones are often associated with hematite, Fe,O,, and with chamosite, 3 F e 0 . A1,0,. 2Si0, * H,O. These seem to be mostly primary, though possibly non-marine. In the Wabana (Newfoundland) deposits, the hematite-chamosite oolites appear to be cut through by algal borings, whereas the siderite clearly resulted from diagenetic replacement (Hayes, 1915). Equally striking is the syndiagenetic pyritization of the Cleveland chamosite-oolite (Lower Jurassic of England), and the incorporation of penecontemporaneous pebbles of pyritized ironstone in the same formation (Hatch et al., 1938, p. 135). In the Middle Jurassic Northampton “ironstones” of England and in the contemporaneous Minette iron ores of Alsace-Lorraine, the iron was clearly derived from the tropical weathering of nearby land. The sedimentary environment was a shallow continental shelf where oolitic carbonates, mudstones, siltstones, and sandstones were accumulating in rhythmic sequences. The ores are now oolitic limonite-chamosite-siderite rocks, but the question of replacement is controversial (Kimberley, 1974). Others argue for a primary chamosite mud (Taylor, 1949; Knox, 1970). Eustatic fluctuations provided periodic exposure (Hallam and Bradshaw, 1979). It constitutes yet another non-actualistic problem area (Taylor, 1949). The situation in the case of the Silurian Clinton (Llandovery) iron ores of the eastern United States (New York State to Alabama) appears to be similar, but here the oolites and carbonate fossil debris have been replaced only by hematite (Hunter, 1970). Desideritization is observed in some instances in the Jurassic iron ores. Just as in the case of dedolomitization, siderites may be replaced by calcite forming very similar textures (Taylor, 1949). Certainly, the most remarkable of all sedimentary iron ore formations are the so-called “BIF”, the banded iron formations of the Precambrian (Trendall, 1968; James and Sims, 1973). The oldest are dated around 3.75 billion years, but the bulk are 2.6- 1.8 billion years old. Characteristically, they consist of finely interbedded chert and magnetite- or hematite-rich layers, in some cases with siderite, pyrite, or other minerals. In a few instances, there are oolitic textures. From the evidence of ancient stromatolites and other traces of photosynthetic life, it is likely that free oxygen was being generated at an early stage, so that ferrous iron liberated by weathering could be oxidized to hematite (Dimroth and Kimberley, 1976). Nevertheless, it is argued by Cloud (1973) that all, or almost all, of the oxygen was rapidly locked up in these massive oxides and so its seasonal supply may well have played a limiting boundary condition to the BIF-layering.

95 Moore and Maynard ( 1929) have explained, and demonstrated experimentally, how ferric oxide and silica can be alternately precipitated from sea water. It is puzzling, however, that the great banded iron ores (Fe20,-SO,) are only Precambrian. They are known from all over the world, but are present only in the older rocks. Hough (1958) has suggested that this alternation is a seasonal phenomenon in giant lakes, and it seems not improbable that the Precambrian sea had a salinity somewhat lower than the present (and a lower pH) so that the ocean of that time might be compared geochemically with a brackish lake (Fairbridge, 1967a). Siderites are also known in these deposits. It seems unlikely that such a lacustrine condition would favor sideritization of aragonitic sediments. Instead, the primary precipitation of cherty siderite probably occurred under local photosynthetic removal of CO, and elevation of pH. The deposits were modified in some cases by later hydrothermal action. Bauxitization Bauxite is the rock name for a mixture of alumina minerals, amorphous or crystalline hydrous aluminium oxides and aluminium hydroxides, mainly gibbsite, often with boehmite and diaspore. In many places it is an aluminous laterite, i.e., an indurated residual soil or paleosol. “Bauxitization” is believed to be usually the epidiagenetic alteration of the parent material, a kaolinite, by desilicification (Erhart, 1973b). In some cases, the parent material may be feldspars in the underlying bedrock or in volcanic ash carried in subaerially. The hydrolysis of kaolinite may be expressed as follows:

A120,. 2 SiO, . 2 H,O (kaolinite)

+ H 2 0+Al,O,.

+

n H 2 0 2 SiO,. 2 H,O (water) (bauxite) (silicic acid)

Contemporary large-scale bauxitization does not appear to have been demonstrated, so that yet another non-actualistic process must be assumed (Davidson, 1964). Bauxitization appears to be relatively restricted in time and space, occurring only in tropical latitudes in the littoral belt or at relatively low elevations. Quite exceptionally, a lacustrine site is known. The most likely sites include tropical mangrove swamps, because the most important deposits in South America (Surinam, Guyana) are associated with a Miocene shoreline and contain fossil mangrove roots (Valeton, 1972). Related deposits are colluvial spreads apparently derived from them, but distributed during the Pliocene and Quaternary phases of lowered base level. Other deposits do not fit this scenario and appear to be related to phases of lowered sea level in otherwise marine limestone sequences. Although it is widely assumed that the bauxites are derived from the clays liberated by

weathering of immense thicknesses of carbonate deposits, the volumetric problem is unsurmountable. Thus it seems likely that in most cases the clays are derived from volcanic ash weathering. Again, coastal swamps with warm temperatures and low p H may provide the necessary biogenically controlled epidiagenetic setting, to be followed by desiccation with partial dehydration.

GLOSSARY Anadiagenesis (adj. -etic): Lithification or other modification of sediments during deep burial, marked by expulsion and upward migration of connate waters and other fluids (petroleum, etc.), often marked by high pH and low Eh (Fairbridge, 1967b). Anamorphism (adj. -ic): Metamorphism at depth, forming more complex minerals (Van Hise, 1898; modified to exclude low-temperature alteration, i.e., by diagenesis). Authigenesis (adj. -ic, -om): Formation of new sedimentary minerals in situ, within the enclosing sediment, during or after deposition (e.g., Pettijohn, 1957). Diagenesis (adj. -etic): Physical and chemical changes which a sediment undergoes after deposition and during lithification, without introduction of heat (6ver ca. 300°C) or great pressure (ca. 1000 bars) (Von Guembel, 1868; modified slightly by Walther, 1894; Fairbridge, 1967b). Epidiagenesis (adj. -etic): Lithification or other modification of sediments during and after uplift or emergence, characterized by infiltration of meteoric water and downward migration, usually marked by low pH and high Eh. Near the surface merges with the zone of weathering. ( Fairbridge, 1967b). Epigene (adj.): As a general term-all processes or phenomena produced at or near the earth’s surface (Geikie, 1879); specifically for mineral deposits formed later than the enclosing rocks or by secondary alteration. Epigenesis (adj. -etic): Changes in the mineral character of a rock due to external influences (A.G.I. Glossary). Also applied to mineral deposits, as epigene. One may have both epigeneric supergene (with descending waters) and epigenetic hypogene (ascending waters). Halmyrolysis (adj. -1ytic): Geochemical modification of sediments during deposition, due to reactions with sea water (ionic transfer), originally called “submarine weathering” by Hummel (1922); but applies also to ionic rearrangement and replacement (Pettijohn, 1957). Hypogene (adj.): Minerals, or changes in rocks, related to ascending waters of magmatic origin, specifically applied to mineral deposits, but formerly to any deep-seated endogenic processes, involving magmatism and metamorphism (Lyell, 1833; Geikie, 1879). Ka;amorphism (adj. -ic): Alteration of rocks, particularly solution and breakdown at or near the earth’s surface, due to either supergene or hypogene waters, forming simple minerals from complex (Van Hise, 1898). Thus it includes both weathering (upper zone) and near-surface cementation (lower zone), but Leith and Mead (1915) excluded the latter. Lithification: The complex of processes that converts an unconsolidated sediment into a hard rock, including compaction, dehydration, cementation and induration (e.g., Pettijohn, 1957). Lithogenesis (adj. -etic): Synonymous with petrogenesis, relating to the origin of a rock (A.G.I. Glossary; Bates and Jackson, 1980). Supergene (adj.): Applied to mineral deposits or enrichment related to descending waters (A.G.I. Glossary). One may have both epigenetic supergene deposits and syngeneric supergene deposits (as in manganese nodules).

97 Syndiagenesis (adj. -etic): Modification of sediments during and immediately following deposition, often by biochemical influences, marked by extreme variations in pH and Eh. (.Bissell, 1959; Fairbridge, 1967b). Syngenesis (adj. -etic): Formation of mineral deposits more or less contemporaneously with the deposition of the enclosing rocks, i.e., the opposite of epigenesis (A.G.I. Glossary). Specifically refers to the time of geochemical changes, i.e., penecontemporaneous (Fersman, 1922).

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113 Weaver, C.E., 1958, Geologic interpretation of argillaceous sediments. BUN. A m . Assoc. Pet. Geol., 42: 272-309. Weaver, C.E., 1960. Possible uses of clay minerals in search for oil. Bull. A m . Assoc. Pet. Geol., 44: 1505-1518. Weiss, M.P., 1954. Feldspathized shales from Minnesota. J . Sediment. Petrol., 24: 270-274. Weiss, M.P., 1958. Corrosion zones: a modified hypothesis of their origin. J . Sediment. Petrol., 28: 486-489. Wells, A.J., 1962. Recent dolomite in the Persian Gulf. Nature, 194: 274-275. Wermund Jr., E.G., 1961. Glauconite in early Tertiary sediments of Gulf coastal province. Bull. A m . Assoc. Pet..Geol., 45: 1667-1696. West, I.A., Ali, Y.A. and Hilmy, M.E., 1979. Primary gypsum nodules in a modern sabkha on the Mediterranean coast of Egypt. Geology, 7: 354-358. Wetzel, O., 1933. Die in organischer Substanz erhaltenen Mikrofossilien des baltischen Kreidefeuersteins. Paleontographica, 47: 146- 188. Wetzel, W., 1955. Seltene Metallverbindungen in Sedimenten. Geol. Rundsch., 43(2): 464-469. Weynschenk, R., 1951a. The problem of dolomite formation considered in the light of research on dolomites in the Sonnwend Mountains (Tirol). J . Sediment. Petrol., 21: 28-31. Weynschenk, R., 1951b. A new sedimentary petrological interpretation of the results reported by the Swedish Albatross Deep-sea Expedition 1947-48. J . Seqiment. Petroi., 21: 82-84. White, D.E., 1957. Magmatic, connate and metamorphic waters. Bull. Geol. SOC.A m . , 68: 1659-1682. Whitehouse, U.G. and McCarter, R.W., 1958. Diagenetic modification of clay mineral types in artificial sea water. In: Clays and Cloy Minerals. Natl. Acad. Sci./NRC, Publ., 566: 81-119. Wickman, F.E., 1944. Some notes on the geochemistry of the elements in sedimentary rocks. Ark. Kemi, Mineral. Geol., 19(B): 1. Williams, D., 1960. Genesis of sulphide ores. Proc. Geol. Assoc. (Engl.), 71: 245-284. Willman, H.B., 1942. Feldspar in Illinois sands. Ill. State Geol. Sum., Rep. Inoest., 79: 87 pp. Wilson, J.L., 1975. Carbonate Facies in Geologic History. Springer, Berlin, 471 pp. Zen, E-An, 1960. Carbonate equilibria in the open ocean and their bearing on the interpretation of ancient carbonate rocks. Geochim. Cosmochim. Acta, 18: 57-71. Zenger, D.H., 1972a. Significance of supratidal dolomitization in the geologic record. Bull. Geol. SOC.Am., 83: 1-12. Zenger, D.H., 1972b. Dolomitization and uniformitarianism. J . Geol. Educ., 20: 107- 124. ZoBell, C.E., 1942. Changes produced by micro-organisms in sediments after deposition. J . Sediment. Petrol., 12: 127-136. ZoBell, C.E., 1946. Studies on redox potential of marine sediments. Bull. A m . Assoc. Pet. Geol., 30: 477-513.

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115 Chapter 3

DIAGENESIS IN ARGILLACEOUS SEDIMENTS ARIEH SINGER and GERMAN MULLER

INTRODUCTION AND DEFINITIONS

The writers prefer to apply the term “diagenesis” to all changes which take place in a freshly deposited sediment until it reaches the stage of metamorphism. According to Read and Watson (1962): “Diagenesis comprises all those changes that take place in a sediment near the earth’s surface at low temperature and pressure and without crustal movement being directly involved. It continues the history of the sediment immediately after its deposition and with increasing temperature and pressure it passes into metamorphism”. In physicochemical terms, diagenesis can be regarded as the process of equilibration of mineral (mostly detrital) phases with aquatic environments of differing salinities, under conditions of increasing temperature and pressure. According to Friedman and Sanders (1978), diagenesis involves, among other things: (1) compaction; (2) addition of new material; (3) removal of material and transformation of material by: (4) change of mineral phase, or (5) replacement of one mineral phase by another. The changes undergone by particles from the time of arrival into the aqueous basin up to their lithification under conditions of high pressure and temperature in the nearly total absence of liquid water can be seen as a continuum. The accurate delimitation of those processes to be designated as “diagenetic”, presents, therefore, some difficulties. According to most European authorities, such as, for example, Millot ( 1970), diagenesis starts only after deposition of the particles, that is, processes occurring during their transport or suspension are excluded. Diagenesis passes into iow-grade metamorphism when recrystallization of sheet-silicates occurs at about 200°C (Dunoyer de Segonzac, 1970). The term “weathering” covers the destruction of rocks and minerals near to or on the land surface, which is not constantly under water, by exogenic processes (e.g., due to insolation, frost, water, atmosphere, and. organisms) leading to the formation of soils. Whereas weathering commonly involves the loss of metallic cations, compensated by the gain of water, the inverse is the case with diagenesis. Taylor (1964) compared diagenesis with weathering in the following way: “For the most part diagenetic changes involve increasing lithification, weathering the reverse. In a sense, weathering may be regarded as retrograde diagenesis.”

116 The largest part of the products of weathering, together with a certain amount of unweathered rock materials, is carried by rivers into the sedimentary basins (mainly into the sea). After some time they are deposited there and become part of the sediment. During this stage of transportation, the rock and mineral particles undergo further mechanical and chemical alteration. This applies in particular to the most finely grained material because of its relatively large surface area. These particles are particularly strongly reactive. In addition, they can remain floating in the marine sedimentary basins for centuries because of the low settling velocity. In the transition from the fluviatile to the marine environment, the physicochemical conditions change radically. The still floating minerals go through the same process of adjustment as the particles which have already settled on the surface of the sediment. Inasmuch as many physicochemical and chemical processes are similar or even identical during the subaqueous transportation-alteration and the earliest stage of diagenesis, which takes place in the upper layer of the sediment and is largely controlled by the chemistry of the subjacent water, it seems advisable to deal with these processes together and to give them the same name. A possible choice for an overall term for the processes occurring in a marine (or saline) environment is the term “halmyrolysis” introduced by Hummel (1922) (halmyros = salty; lysein = to dissolve). Thus, halmyrolysis can be defined as: all chemical and physicochemical processes which occur during the marine transportation-alteration and the marine pre-burial stage of diagenesis. Packham and Crook (1960) have already used the term halmyrolysis in this sense. Because an overall term for the physicochemical processes occurring in the fresh-water environment is lacking, Muller ( 1967) proposed the term “ a q ~ a t o l y s i s(aqua ~ ~ = water; lysein = to dissolve) for all processes which take place in fresh water during transportation-alteration and in the earliest stage of diagenesis. Thus, aquatolysis may be defined as: all chemical and physicochemical processes which occur during transportation-alteration and pre-burial stage of diagenesis in fresh-water environment (Table 3-1). The various stages of diagenesis can be divided into: ( 1 ) pre-burial stage, (2) shallow-burial stage, and (3) deep-burial stage. The pre-burial and shallow-burial stages correspond to the “early diagenesis” of Dunoyer de Segonzac (1970), whereas the deep-burial stage is equivalent to his “middle” and “late” diagenetic stages. In argillaceous sediments, the depth at which the sediments are buried under younger deposits is more essential for the diagenetic evolution than the length of time which passes after the deposition of the sediments. This is especially true of the physical changes which take place in the sediment in the course of diagenesis. This becomes particularly clear by a comparison of a sediment core from Lake Zurich with

117 TABLE 3-1

Relationships between parent rock, weathering, diagenesis and metamorphism SOIL

SEDIMENT

Fresh - water 0000000 Marine

Submorine basement

~0000000000000

1

' /

AQUATOLYSIS

1

+++++++

P A R E N T ROCK

HALMYROLYSIS

1

Fresh-water P r e - b u r i o l M a r i n e DIAGE N E SIS

1

METAMORPHISM

those from the Santa Barbara Basin and the Black Sea (Table 3-11). Although the sediments in the Santa Barbara Basin and the Black Sea, covered by about 5 m of younger deposits, are at least 10 to 100 times older than those in Lake Zurich, the decrease in porosity with increasing depth is almost identical. Obviously, the time during which a sediment is situated in a certain physicochemical environment is important as far as the chemical and mineralogical processes are concerned. The pre-burial stage is easiest to define. It comprises the physicochemical processes (halmyrolysis and aquatolysis) which take place in the upper (youngest layers of the sediment) in the presence of oxygen. To differentiate between the shallow-burial and the deep-burial stages is much more difficult. In the case of argillaceous sediments, the line can be drawn where a soft clay mud becomes an indurated, firm and coherent mudstone (mudstone fissility = shale). This consolidation can be brought about by compaction alone; however, there is often a combination of compaction and cementation. The shallow-burial and deep-burial stages of diagenesis thus correspond to the two principal stages in the lithification of argillaceous sediments proposed by Lomtadze (1955) on the basis of experimental work: (1) the conversion of argillaceous mud to mudstone (or shale), and (2) the conversion of mudstone to argillite (employing the term argillite to mean a non-metamorphosed rock). Slate is a metamorphic rock. In terms of porosity, the point at which a clay mud is converted into mudstone lies at

+

118 TABLE 3-11 Decrease of porosity with depth in subaqueous argillaceous sediments (After Rukhin, 1958. Emery and Rittenberg, 1952, and Ziillig, 1956, cited by Miiller, 1967, p. 163.) Depth (m)

Porosity (%) Black Sea

0.00 0.20 0.50 1.oo

2.00 3.OO 4.00 5.00 6.00 7.00 8.00 I

' Santa Barbara Basin, Calif. (U.S.A.)

Lake Zurich (Switzerland) 88

-

79 73

72 71 70 65

82 81 79 77 75

74 73

78 77 75 73 71 68 66 64 62 60

Porosities in percent calculated from water content

about 30%. At this porosity, the initial bulk volume of the clay mud decreases by about 50%, with loss of at least 50% of the connate water. Hamilton ( 1959), who studied the relationship between overburden pressure and porosity, observed that no matter how porous the original material, when the pressure reaches about 100 kg/cm2, the porosity usually decreases to about 29% for various clay-rich sediments. In the Tertiary clay sediments of the Po Basin and of Venezuela, the 30% porosity limit is reached when the sediment cover is about 500 m thick. For subaqueous argillaceous sediments, Strakhov ( 1956) assumed that lithification is almost complete at a depth of about 250-300 m below the depositional interface. As compared to the near-surface sediment layer the increase in pressure can amount up to 80 atm and the increase in temperature up to an average of 9°C. At the boundary between shallow- and deep-burial zones, the depth and pressure are around 500 m and 50- 100 kg/cm2, respectively. Laboratory experiments on clay-water systems showed that at pressures above 50 kg/cm2 the influence of the electrolyte content and the exchangeable ions is insignificant. There are essential differences between marine and fresh-water sediments, because the chemical-mineralogical and partly also the physical changes in a sediment during diagenesis depend mainly on the chemistry of the water (subjacent and interstitial) in contact with the mineral particles. This applies at least to the first two stages of diagenesis. During the third stage of

119 diagenesis it can be assumed that the pore solutions in most cases have already become similar in composition. Inasmuch as detrital clays have commonly been formed during continental weathering under the impact of meteoric, non-saline water, diagenetic effects are much more pronounced in the highly saline marine environment. Diagenetic effects, however, can also be obtained in non-marine sedimentation basins of comparable solute concentration, such as saline lakes. In this chapter, “argillaceous sediments” are defined as fine-grained sediments (with an average grain size in the range of about 1-10 pm) which mainly consist of silicate clay minerals (chiefly layer-silicates). These types of sediments are usually to be found in the deeper offshore regions of marine and lacustrine basins. INITIAL COMPOSITION OF RECENT ARGILLACEOUS SEDIMENTS

In order to understand the diagenetic changes in inorganic and organic matter, it is essential to be familiar with the initial composition of a clay sediment . Among the most important allogenic components which come from outside the sedimentary basin are: (1) clay minerals including gibbsite, (2) quartz, (3) feldspar, (4) carbonates, ( 5 ) amorphous silica and alumina, (6) pyroclastic material, and (7) organic matter. In addition, biogenic carbonates, biogenic amorphous silica, and organic matter are formed in the basin itself. Clay minerals and quartz are most stable against changes in the physicochemical environment, whereas pyroclastic materials are the least stable. The important allogenic and biogenic components of modern clay sediments are described below. Clay minerals (including gibbsite)

The distribution of clay minerals in Recent ocean sediments has been shown by Biscaye (1969, Griffin et .al. (1968) and Rateev et al. ( 1969) to be chiefly detrital and related to general source areas on adjacent continents. In the tropics, where leaching and chemical weathering are intense, there is a conspicuous abundance of the kaolin group minerals and gibbsite near continental masses (Fig. 3-1). The kaolinite content in the sediments varies from 40 to 60% of the total clay mineral content. Areas of its maximum content occur in the equatorial parts of the Indian and Pacific oceans, with a gradual decrease in abundance towards the poles. Gibbsite is localized in the tropical and subtropical latitudinal belts. The zone of maximum concentration of smectite (40-60%) coincides in general with the zone of high kaolinite content, although it is somewhat wider. This general distribution pattern is modified by superposition of

120

121

patches of azonal concentrations associated with volcanogenic ash material. Thus, in the Atlantic, the occurrence of southern and northern maxima is due to the development of areas of Recent and Quaternary volcanism, such as near Iceland (Rateev et a]., 1969). In these areas, the clay mineral distribution appears to be determined by the ratio of detrital input to in-situ neoformation. Smectite concentrations are greater in the Southern Hemisphere oceans, where there is a larger input of volcanic material relative to detrital clays, because the former lends itself to alteration to smectite. The highest concentrations of smectite occur in the South Pacific, associated with phillipsite and volcanic glass shards. Maximum concentrations of chlorite (30%)are located in northern-southern zones, whereas equatorial zones are nearly devoid of this mineral. The distribution pattern of chlorite is thus distinctly zonal. For that reason, in recent deep-sea sediments of the North and South Atlantic the distribution of chlorite has been observed to be nearly reciprocal to that of kaolinite (Biscaye, 1965; Zimmermann, 1977). More illite occurs in the Northern Hemisphere ocean sediments than in the southern oceans, reflecting the impact of continental input in the former. Mite concentrations in marine sediments bear close relation to river-borne sediments, whereby the highest illite concentrations are found in the North Atlantic, in association with chlorite. The clearest latitudinal zonation in the distribution of illite is observed in the sediments of the Indian, and to a lesser extent, of the Pacific oceans. While maximum contents (60-80%) i n the Atlantic Ocean were observed in its northern parts, a near-equatorial zone of minimal contents is missing. Small areas of very high illite contents occur at the southern end of the African continent, probably representing the clay weathering products of the arid climates prevailing in parts of this continental mass. This review of clay mineral distribution in Recent marine argillaceous sediments has been limited to the most common occurrence patterns. Although local accumulations of relatively rare layer-silicates do exist (palygorskite, talc, pyrophyllite, serpentine, etc.), in such cases their occurrence Fig. 3-1. Map showing the distribution of kaolinite in the fine-pelitic fraction of the bottom sediments of the world oceans. I , 2 , 3 and 4 =kaolinite contents expressed as percentages of the total clay mineral contents. Stations where observations have been made, are marked with solid circles ( 5 ) in each ocean basin. On the continents, the distribution of the main weathering zones is shown: I =zone of moderately humid climate; I I =zone of tropical humid climate; I I I = tectonically active areas lacking well-formed weathering crusts; I V = for each continent, the total river load discharged, given in million of tons per year; V=thc amount of river load (in the same units) and the direction of discharge, at each of the main discharge points. The kaolinite percentages are as follows: I =40-60; 2 =20-40; 3 = 10-20: 4 = 10; 5=stations. (After Rateev et al., 1969, p. 28, fig. I.)

122 always depends greatly on the presence of these minerals on the adjacent continent. For example, there are considerable amounts of pyrophyllite in the basins adjacent to Central America and the northern part of South America, pointing to source areas on the northern coasts of Columbia, Venezuela, French and British Guiana, Surinam, and Brazil, and the southern parts of Cuba, Haiti, and the Dominican Republic (Biscaye, 1964). The clay mineral composition of lake sediments is determined to an even greater extent by the hinterland. The most important clay minerals here are also illite, kaolinite, chlorite and montmorillonite, as well as random mixed-layer illi te-montmorillonite. Quurtz and feldspars

Quartz and feldspars (particularly plagioclase) are the most important allogenic minor constituents of argillaceous sediments. The abundance and distribution of quartz in Recent ocean sediments is directed primarily by aeolian input from adjacent desert regions (Rex and Goldberg, 1958; Windom, 1975; Windom and Chamberlain, 1978). Quartz in the Holocene sediments of the Indian Ocean is derived mainly from continental sources and has been transported in the form of atmospheric dusts from Arabian and Australian deserts into the adjacent ocean (Kolla and Biscaye, 1977). It appears that during the last glacial period the aeolian influx of quartz into the Indian Ocean was higher. An aeolian origin for feldspar in North Atlantic sediments was shown by Windom and Chamberlain (1978). The increase in quartz content near Antarctica in the Northern Atlantic probably stems from glacial outwash (Biscaye, 1964); a similar tendency for feldspars in these areas has also been generally noted and explained in the study of Biscaye (1964). The feldspar distribution in South Pacific pelagic sediments was investigated by Peterson and Goldberg (1962). Most of the feldspars are of volcanic origin, and several source areas could be established (especially for basic and acidic groups). The composition of the plagioclases ranges between oligoclase and anorthite. Alkali feldspars (sanidine, orthoclase) also show a wide distribution. Curhonutes

Allogenic carbonates (such as calcite and dolomite) from outside of the sedimentary basin can be assumed to play only a very minor role in argillaceous sediments as compared to the biogenic carbonates (mainly calcite and aragonite). Between the argillaceous muds with no or little carbonate content (red clay) and pure biogenic oozes (Gfobigerina ooze covers enormous areas in the Atlantic, Pacific and Indian oceans) all transitions are possible.

123 Amorphous silica and alumina

In argillaceous marine deposits, amorphous silica (and alumina in some cases) may become very abundant. In the sediments of the Atlantic Ocean, amorphous silica content ranges from 1 to 56% (Biscaye, 1964). Most of this silica is biogenic and was produced by planktonic diatoms, radiolarians and silicoflagellates, and some benthonic sponges. An alumina content of 0-3% found in the same sediments by Biscaye (1964) may be accounted for by the varying gibbsite contents of the samples. Gibbsite is strongly attacked as a result of the leaching procedure for the determination of the amorphous silica. The presence of amorphous alumina, therefore, cannot be definitely confirmed. Sediments with high biogenic opaline silica are found in the Subarctic Convergence, the Equatorial Divergence, and the divergences along the west coasts of the continents (Arrhenius, 1963). Accumulations of non-biogenic amorphous silica and alumina (up to 60%) were reported by Moberley (1963) from sediments adjacent to Hawaii, which are derived from tropically weathered basalt. In fresh-water lakes, small amounts of biogenic opaline silica are mainly produced from diatoms. Yellow-brown algae and siliceous fresh-water sponges are of minor importance. During the Ice Age, diatoms played a much more important role in sedimentary processes in lakes. Pyroclastic material

The main location of recent volcanism is the area around the Pacific Ocean (circum-Pacific Circle). More than two thirds of all active volcanoes are situated here. According to Sapper (cf. Brinkmann, 1961), between 1500 and 1914 about 18 km’ of lava and 330 km7 of pyroclastic material were erupted from the 339 active volcanoes in the circum-Pacific Belt. During the eruption (in 1883) of the Krakatau in the Sunda Straits, about 18 km7 of pyroclastic material were hurled up as high as 50 km into the atmosphere. Extremely fine-grained volcanic ash circled the earth several times. In addition to this terrestrial volcanism, mainly limited to the rims of the continents, there is also a submarine volcanism the magnitude (dimensions) of which can hardly be estimated. It may be assumed that submarine volcanism has played an important role, especially in the Pacific area. Huge amounts of palagonite, which spread over enormous basin areas associated with the seamounts and guyots, are a definite indication for submarine volcanism. The tuffs of terrestrial volcanoes, mainly consisting of volcanic glass, either land directly in the sea during eruption or are carried into the sea from

124 already existing tuffs on the continent. Acid porous glass (pumice) because of its low specific weight due to the high porosity can remain drifting in the sea for a relatively long period of time. In the case of submarine eruptions, the conditions of genesis of the partially devitrified glass (palagonite) are completely different, because it is not the high gas content of the magma but the interaction of hot basic lavas with cold sea water which is decisive. In the Pacific, the palagonitk distribution generally coincides with the area of phillipsite occurrence. In the recently formed sediments of the Atlantic area, volcanic glass occurs much less frequently. Accumulations of glass should occur particularly in the immediate neighbourhood of volcanic areas (Mediterranean, Cape Verde Islands, Canaries, Azores, Central American volcanic provinces, etc.). For example, the Recent sediments in the Gulf of Naples are almost exclusively formed from volcanic glass and its alteration products (Muller, 1961). Orgunic mutter

The contribution of Degens and Mopper to the companion Volume of this book (Chapter 4) covers the subject of organic matter composition and diagenesis in argillaceous sediments.

INITIAL (PRE-BURIAL) POROSITY A N D STRUCTURE OF ARGILLACEOIJS SEDIMENTS

Initial porosity

The initial porosity and thus the water content' of argillaceous muds is very much higher than that of sands. Clay muds from Recent sea and lake bottoms as a rule have a porosity of 70-9056, corresponding to a water content of about 50-80%. In sands, porosity is only 30-50% which corresponds to 20-30% water content. Figure 3-2 shows the dependence of the water content of Recent clay muds on the amount of the clay fraction (< 2 pm) in different sedimentary basins: with increasing clay content the water content and porosity increase. Sands essentially consist of more or less rounded particles of quartz and feldspar, the geometrical arrangement of which can be compared with

'

In this chapter the water content is expressed as a percentage of the wet weight.

125

90

80-

-r 60--

80

s

;

-m U

Y

c

C

’O

:40.

60

al

s;

c

v)

U L

$

.

I11

20.

V

0

/

50

/

al c

a

/

J 20

40 % Fraction

c

60 2pm

80

Fig. 3-2. Correlation between water content and percentage of clay fraction (<2 pm) in different environments: I = Lake Constance; II=Zuiderzee; I I I = Mississippi delta; I V= Rockport area, Texas; V= California Basin. Environments: I = fresh water; 11-I V= brackish; V=marine. (After Miiller, 1967, p. 136, fig. 1.)

packings of spheres. With decreasing grain size, the amount of layer-silicates increases. Over the last 15 years, a large number of laboratory and in-situ measurements have been made of bulk density, porosity, grain density and other properties of surficial (0-30 cm) sediments of the sea-floor. Data in Table 3-111, given by Hamilton (1976), show that average densities of pelagic (“red”) clays vary from 1.347 g/cm’ for clayey silt textures to 1.414 g/cm3 TABLE 3-111 Abyssal-hill environment: sediment densities, porosities, sound velocities, and velocity ratios (after Hamilton, 1976) Sediment type: deep-sea (“red”) pelagic clay Clayey silt Silty clay Clay

Density, (g/cxn’)

Porosity. (%)

Velocity, (m/sec)

Velocity ratio

avg.

SE

avg.

SE

avg.

SE

avg.

SE

1.347 1.344 1.414

0.020

81.3 81.2 77.7

0.95 0.60 0.64

1.522 1.508 1.493

3 2 1

0.995 0.986 0.976

0.002 0.001 0.001

0.011 0.012

126 for clay textures, whereas corresponding average porosities vary between 81.3% and 77.7%. Although these measurements were made in sediments from the Pacific and Indian oceans and adjacent seas, the listed values are applicable to the same sediment types elsewhere. From the purely geometrical point of view, clay minerals could aggregate to a compact sediment with only a very low porosity. The fact that this is not the case with argillaceous sediments shows that their porosity cannot be understood only on the basis of simple geometrical models as is qualitatively possible with sands. The special behavior of layer-silicates is due to their extremely large surface areas. The sediments of the different sedimentary basins with the same amount of clay-size fraction (or the same average grain size), have different water contents. According to Meade (1964), this may be due to the following important factors: ( 1) influence of clay-mineral composition, (2) influence of interstitial electrolyte solutions, (3) influence of exchangeable cations, and (4)rate of sedimentation. In addition, alkalinity or aqidity and associated organic matter can also play a role. These influences, however, have not yet been studied in detail. Influence of clay mineral composition If pure clay minerals are allowed to settle in water, it becomes evident that in the resulting sediments montmorillonite retains more water than illite, and illite in turn retains more water than kaolinite. These differences are mainly due to the different grain-size distributions and to the differences in the specific surface areas. According to Meade (1964), the specific surface area of montmorillonite is 800-600, of illite 100-65, and of kaolinite 30-5 m2/g. The effect of specific surface area on water sorption is that a montmorillonite particle, which is extremely fine-grained, adsorbs a water envelope of the same thickness as that on a relatively large kaolinite particle. In comparison to the mass, therefore, more water is adsorbed by montmorillonite than by kaolinite. Influence of electrolyte solutions In natural and artificially produced sediments there is a relation between the water content of a freshly deposited sediment and the electrolyte content of the depositional medium. With increasing electrolyte content, water content decreases. Clay-rich fresh-water sediments, therefore, as a rule contain more water than comparable marine sediments This behavior can

’

According to Bolt (1956), the effect of electrolyte concentration seems to change with particle size. In very fine-grained clays (
127 be explained by the different forces (van der Waals’ attractive forces between the clay plates and repulsive forces due to the presence of diffuse electrical double layers on the charged clay plates-“osmotic swelling”), which have a mutual effect on the neighboring clay particles. For details on this subject, the reader is referred to the work of Van Olphen (1963a. b), Meade (1964) and Yariv and Cross (1979). High electrolyte contents bring about a coagulation of the particles. This is important for the behavior of clay minerals in the transition from fresh water to salt water. The coagulation effect of the ions increases rapidly with increasing valency (Schulze-Hardy valency rule). The concentrations of uni-, di- and trivalent cations, which are necessary for flocculation are in the ratio of 500/10/1. The sequence of the flocculation intensity is: A1 > Ca > Mg > K > Na. Influence of exchangeable cations Experiments carried out by Samuels ( 1950) on relatively coarse-grained kaolinite showed that the effect of exchangeable cations on the initial water content of the sediment is only slight. Kaolinites saturated with A1 have larger water contents than Ca- and Na-saturated kaolinites, but the differences between the latter two clays are extremely small. With relatively fine-grained montmorillonite the differences are much greater and the situation is reversed: Na-saturated clays contain more water than Ca-montmorillonites and these, in turn, contain more water than Al-montmorillonites. For further discussions on this subject, see Meade (1 964). Rate of sedimentation Up to now no comprehensive studies were made on the dependence of initial porosity on the annual rate of sedimentation. Observations made by Fuchtbauer and Reineck (1963) in the southern part of the North Sea seem to indicate that with a high rate of sedimentation the porosity is greater than with a low rate. Clay muds from a Recent bay (rate of sedimentation up to 50 cm/year) had a porosity of 83% on the sediment’s surface, whereas clays from the Wadden Sea and the foreshore had a porosity of only about 70%. Also the relatively very high porosities in the uppermost few centimeters of fresh-water sediments in lakes (Table 3-IV) with an estimated sedimentation rate of about 1-5 mm/year, seem to indicate that with this high rate of sedimentation an abnormal porosity exists only for a short period of time. As the porosities of the lower layers show, they are reduced to the normal values after a very short period of time.

128 TABLE 3-IV Water content of young lake sediments (after Ziillig, 1956, cited by Miiller, 1967, p. 161, table IV) Zuger See

Lake Constance

depth of burial (cm)

water content

calculated porosity

(48)

(48)

0 1.2 2.4 3.6

83.6 74.0 74.2 70.6

92 87 87 85

.

depth of burial fcm)

water content

0 -0.5 0.5- 1 .O 1.2-1.5 2.3-2.8 4.0-4.6

73.8 72.2 68.0 61.7 60.3

calculated porosity

(%I 87 86 84 80 79

The combined effects of different factors in natural sediments

In natural argillaceous sediments, several clay minerals &cur together, as a rule, and each clay mineral can occur in different grain-size classes; furthermore, varying amounts of organic substances and non-clay minerals can be present with the clay minerals. Consequently, the combined effects of several single factors described above and their interrelationships are very complicated and cannot be investigated without exact knowledge of the sediments in question, which in most cases is not available. As shown in Fig. 3-2, in comparison to the marine sediments the fresh-water sediments of Lake Constance, with a clay content of less than 40%, show the expected high porosity. With a clay content of more than 40%, however, this situation is reversed. This behavior could be explained if both sediments had a high content of particularly fine-grained clay minerals, as explained earlier (effect of electrolytes). One cannot ascertain, however, whether this is the case, because only the fraction 2 p m was determined, without making a further subdivision.

-=

initial structure

The dependence of primary porosity on particle size, type of clay mineral, electrolyte of the depositional medium, and exchangeable cations, as shown in the previous section of this chapter, determines the structure of the newly formed clay sediments. Van Olphen (1963b, 1964) suggested that the interaction between plate-like clay-particles may result in three different modes of association: (1) adhesion between the flat oxygen faces of two parallel particles (face-to-face, FF); (2) adhesion between broken bond surfaces of neighboring particles (edge-to-edge, EE); (3) adhesion of broken bond

129 surfaces to flat oxygen faces (edge-to-face, EF). The EE and EF types of association lead to three dimensional voluminous ”card-house’’ structures and to aggregates that can be classed as flocs. The FF type of association, which is the slowest, leads to thicker and larger flocs, to which the terms “oriented aggregates”, “books”, or “domains” have been applied. Figure 3-3 shows this fabric in electronmicrographs made by Rosenquist (1962) of several marine clays. In the case of sedimentation in non-saline river or lake environments, if the electrolyte concentration is below the critical coagulation concentration of the FF mode of particle association, EE and EF modes are obtained, forming the open arrangement referred to as “card-house’’ (Yariv and Cross, 1979). This structure may later, under compaction, give rise to an oriented type of structure. In fresh-water clays, Rosenquist (1962) observed a greater degree of parallel orientation between clay particles. At electrolyte concentrations above the critical coagulation concentration,

Fig. 3-3. Electronmicrograph showing mutual arrangement OF minerals in the blue Oslo clay. (After Rosenquist, 1962, p. 18, fig. 5.)

130

Fig. 3-4. Two-dimensional model of the aggregate structure. (After Von Engelhardt and Gaida, 1963, p. 927, fig. 1 1 . )

face-to-face (FF) coagulation becomes dominant. Therefore, sedimentation in a marine environment, where salt concentration is above the critical coagulation level and the FF mode of particle association becomes more stable than either the EE or EF mode, individual platelets aggregate together in “books” or “domains” and the “books” then form an open arrangement referred to as “book-house” (Yariv and Cross, 1979). From the behavior of permeability during the compression of clays with different electrolyte contents in their pore solutions, Von Engelhardt and Gaida (1963) concluded that clays which settled from solutions rich in electrolytes have a heterogeneous or aggregate structure (Fig. 3-4). Porosity is due to the internal pore spaces between aggregates. Clays deposited from solutions poor in electrolytes will consist of small aggregates or of free primary particles. Orientation, therefore, will be good and permeability low because the channels for fluid flow are very small. With higher electrolyte concentration, aggregates will be large, orientation poor, and permeability high. CHANGES IN CHEMISTRY AND MINERALOGY DURING DIAGENESIS

Changes during transportation and pre-burial stage (aquatolysis and halmyrolysis)

A quatolysis

The clay minerals carried by rivers and deposited in lakes and oceans are

131 derived from soils, or directly from the outcropping sedimentary, igneous and metamorphic rocks in the hinterland (see Table 3-1). The minerals of soils and, to a large extent, also those of the sedimentary rocks are largely adjusted to the conditions on the earth surface. This, however, does not always apply to those minerals of igneous and metamorphic rocks which have not undergone the soil-forming weathering processes. Analyses for the K-content of clay-grade mica minerals, which are suspended in the Alpenrhein (Rhine river before it enters Lake Constance) and deposited in Lake Constance and which may have been derived from micas of outcropping igneous and metamorphic rocks in the Alps, show a remarkable loss of potassium, reaching up to 20% of the initial K-content. There is a corresponding increase in the water content (hydroxonium) (G. Muller, unpublished data). In this case, there is a very rapid process of the formation of clay minerals from micas during the subaqueous transportation. The mechanical breakdown and aquatolysis result in the formation of (1) illite (dioctahedral) from muscovite, and (2) ledikite (trioctahedral) from biotite. I t is probable that after deposition these processes also c,ontinue during the pre-burial stage. Possibly, aquatolysis also occurs in the case of K-rich feldspars. Pedro et al. ( 1978) described diagenetic peloidal nontronite in surface sediments of Lake Chad. They showed that nontronite is obtained by the transformation of oolitic sand-sized Fe-oxide grains, in a process that begins in the interior of the grain and proceeds outward. They suggested that low silica activity in solution, having near-neutral pH and near-zero Eh, favor the process that is analogous to the ane responsible for chamosite and glauconite formation from peloidal material. The coprecipitation of silica from lake waters with hydrated oxides of iron and manganese has been demonstrated by Kato (1969), thus explaining the correlation of some lacustrine silica levels with the redox cycle in the lakes. Nontronite neoformation in Recent sediments of Lake Malawi has been explained by Muller and Forstner (1973) as resulting from the percolation of geothermal solutions rich in SiO, through the sedimentary fill of the basin, and their reaction with ferrous iron. Under reducing conditions and a pH lower than 7, iron and manganese are leached into the lake water. In more highly aerated areas of the lake, iron hydroxides and opal precipitate simultanously. Experimental synthesis of iron layer-silicates led Harder ( 1978) to suggest that low-temperature formation of iron layer-silicates is feasible under reducing conditions within relatively short spans of time in the presence of soluble iron and silica. In the clay minerals already derived from soils or sediments, or other layer-silicates of the metamorphic rocks, no considerable change is to be expected to occur during the subaqueous transportation.

132 Halmyrolysis When terrigenous clay minerals are transported to the sea, probably the first change to result from the fresh-water-saline-water transition is ion exchange. Russell (1970) suggested that ion exchange is the only rapid (time scale of a year or less) reaction occurring when land-derived clays encounter sea water. Exchangeable Ca2+ on the clays is exchanged for M g 2 + , K + and Naf from the sea water. Equilibration of the clay adsorption complex with the sea water appears to take place fairly rapidly. Experimental data obtained by Roberson (1974) show that after a few hours, and at most a few days of immersion in sea water, four expanding clays had exchanged most of their original interlayer cation Ca2+ for M g 2 + , K + and Na+ . Thereafter, there appears to be very little, if any, additional reaction. The ratio of adsorbed Mg2+ to adsorbed K varies between about 1 : 1 and 5 : 1 for these samples (Fig. 3-5). A large proportion (up to 40%) of the interlayer cations become non-exchangeable after prolonged contact with sea water. Results

a Poitras Sornple

401

30 20

L

-0>

~

2

m

, , , ;, 4 6 8 1 0 1 2 Time, w k s

%

2

4 6 8 Time, wks

1012

rn

0

0

0

0

i ,aP \

\

m

E c

Y)

; 70 D 0

9

Mg

0

0

V

60

V

D

50-

b Oquoga Somple

V

c:

d Loidig Somple

zziy,;o

40-

10

2

4 6 8 Time, wks

1012 Time, wks

Fig. 3-5. Cation adsorption in sea water as a function of time. (After Roberson, 1974, p. 443, fig. I . )

133 obtained by Russell (1970) indicate that over two thirds of the loss in exchange capacity can be accounted for by the fixation of K + . Also, some of the Na’ and Ca2+ cations become fixed. The decrease in exchange capacity is associated by Roberson (1974) with the formation of a poorly crystalline mixed-layer phase, consisting of illite-like expansible layers. Possibly. this phase can be regarded as a precursor for discrete illite layers. The decrease in expansibility of clays upon contact with sea water may be more pronounced with “degraded illite”, that is a potassium-depleted variety of illite, in which a large amount of K + had been replaced during weathering by exchangeable cations. Cation-exchange processes continue after deposition. A recent study of Atlantic Ocean sediments indicates that the reactions controlling the fluxes of most components across the water-sediment interface occur almost exclusively in the upper 100 cm of sediment (Sayles, 1979). Diagenesis has led to the uptake of Mg2+ and K + and the release of Ca2+, HCO, and Na+ by the solid phases. Contributions of M g 2 + , Ca2+, K + and HCO; from below 100 cm amount to less than 15% of the calculated fluxes across the interface. Reactions in the upper 30 cm account for 70-90% of the fluxes of the components across the interface. Only reactions involving Na may have a deeper source. Contrary to earlier views, it appears highly unlikely that much dissolved SiO, is sorbed on clay-mineral surfaces in the shallow areas of the continental shelf (Siever and Woodford, 1973). In some of the deeper waters of the oceans, however, or in the interstitial water of deposited sediments in areas of diatom ooze formation, dissolved silica levels might be high enough for sorption on clay surfaces. Can the newly acquired cation penetrate the silicate lattice of the clay mineral? Drever (1971a) has shown that in the Recent sediments of h o Ameca Basin, Mexico, the non-exchangeable Mg-content of the clay fraction is higher and the Fe-content is lower in sediments from strongly reducing environments than in similar sediments from less reducing environments. The mineralogy of the clay fraction did not show any parallel changes. According to Drever, in a strongly reducing environment, Fe leaves the smectite structure, to form a sulfide, and Mg enters the same sites from the water, so that gross clay-mineralogy remains unchanged. Commenting on this proposed process, Perry et al. (1976) have shown that addition of Mg by ion exchange can account for at most 50% of the dissolved Mg flux from rivers and ground water. They, therefore, proposed an additional mechanism of Mg removal from solution as a result of basalt alteration to smectite (see below). In an additional comment, Heller-Kallai and Rozenson ( 1978) indicated that direct Mg + Fe exchange in clays seems improbable, but that depletion of Mg in interstitial waters of anoxic sediments may be due to reaction of Mg-containing solutions with partially disintegrated clay.

134 Dunoyer de Segonzac (1970), following the concept of aggradation developed by Millot (1964), described the process whereby clay minerals that have been deprived of some of their cationic components during continental weathering (“degradation”) take up the same or similar cations on contact with sea water. Occurrences reviewed by Millot include mainly the formation of illite and chlorite from disordered mixed-layer structures. The major process involved is the adsorption of K + and Mg2+ into the interlayers. As early as 1956, Whitehouse and McCarter have shown by experimental studies that smectites exposed to artificial sea water for prolonged periods, were altered to yield illite-like and chlorite-like clay minerals. N o alterations whatsoever were produced in the original kaolinitic and illitic clay structures. From a consideration of the large amounts of Al, Fe and Si carried in solutions by fresh water into brackish or marine water, Jeans (1971) postulated neoformation or diagenesis of clay minerals in the very earliest stages of deposition and even before burial. Pre-burial diagenesis of clay minerals is also suggested by the comparison of adjacent deltaic and marine environments of deposition. The formation of metabentonites (K-bentonite) is relegated by Dunoyer de Segonzac (1970), at least partially, to the pre-burial stage of diagenesis. He concedes, however, that a considerable amount of evidence indicates burial diagenesis also for this type of clay mineral. The absence of extensive structural changes in clay minerals during halmyrolysis is possibly due to inhibition by dissolved organics which block interlayer sites otherwise available for uptake of cations (Berner, 1971, p. 184). Summing up, the relative importance of pre-burial diagenesis in the sense of clay-mineral lattice transformations, resulting from contact with sea water of normal salinity, appears to be only very minor. On the other hand, many significant clay-mineral transformations that can be attributed to early diagenesis have very probably resulted from the effect of hypersaline (and possibly warm) marine solutions on detrital clay minerals. Halmyrolysis in the supersaline marine environment During evaporation of sea water, the already high salt content is increased further. Increases in the salt content of sea water can also occur as a result of other phenomena such as hydrothermal activity. When a mineral passes from fresh into supersaline water, the radical change in environment could be expected to result in more significant mineral alterations than the freshwater-sea-water transition. The frequent occurrences of several clay-mineral types in salt clays appear to confirm this hypothesis. Corrensite, a regular mixed-layer chlorite-smectite is most common in the salt-rich Permo-Triassic deposits and is generally believed to be an intermediate step in the evolution of illites towards chlorite by fixation of magnesium. While Lucas and Ataman ( 1968) suggested that this evolution

135 occurred in the Triassic of the French Jura mainly during transport of the particles and before their deposition and burial, evidence summarized by Dunoyer de Segonzac (1970) indicates that this aggradation process could most probably be attributed to the action of highly saline solutions obtained by the gradual confinement of the Triassic basin. The aggradation process may have started during deposition and continued during shallow burial. From its appearance in sediments associated with hard beds, chlorite in Triassic sediments from southern Israel was proposed to have formed diagenetically through the sustained contact of detrital illite with constantly replenished saline water, that may have been of a higher salinity, percolating through the permeable rocks, perhaps in the inter- or supratidal zone (Heller-Kallai et al., 1973). Following diagenetic stages for this occurrence are proposed: dioctahedral illite dioctahedral illite-smectite dioctahedral smectite-di- and/or trioctahedral vermiculite-chlorite di- and/or trioctahedral chlorite. Chlorite, which is relatively abundant in some Messinian sediments from the Mediterranean area is considered by Chamley et al. (1978) to be of diagenetic origin for the following reasons: (a) proximity of the chlorite to sub-regular mixed-layer chlorite-smectite (corrensite) suggesting the initial stages of a “chloritization” of detrital minerals: (b) association with Mg-rich sediments such as authigenic dolomite: and (c) association with authigenic palygorskite. Chlorite and illite are the most abundant clay minerals in Paleozoic evaporite rocks (Droste, 1963). The origin of this association is controversial. It could represent not only primarily detrital accumulations, but also transformation products of detrital clay assemblages. Chlorite and illite are also the major clay minerals in silicate assemblages from an Upper Silurian rock salt bed in New York, U.S.A. (Bodine and Standaert, 1977). Textural features and Br-content of the salt indicate precipitation from shallow-marine brine. A relatively uniform chemical composition of the chlorite, that is distinctly different from that of normal chlorites, suggests an authigenic origin in the marine-evaporite environment (Fig. 3-6). The illite is clearly less degraded than normal shale illite, suggesting that some recrystallization occurred in the hypersaline environment. Postdepositional (and/or postburial) diagenesis resulted in improved crystallinity that involved isochemical recrystallization of the bulk silicate assemblage. The term “hyperhalmyrolysis” is introduced by Bodine and Standaert to denote mineral reactions which occur in the marine hypersaline environment. +

- -

Halmyrolysis in the supersaline lake environment Zeolites are the best known diagenetic products in saline lakes. Their occurrences have recently been summarized by Eugster and Hardie (1978). and Surdam and Sheppard (1978). They form through reaction of volcanic

136 1.0

++ N

2 0.4

N

d

0.2

00 20

22

24

26

20

30

32

3.4

36

Si (atoms per four tetrahedral sites)

Fig. 3-6. Compositional classification of the chlorite minerals with heavy boundaries outlining the probable limits of most “shale chlorite” compositions as determined ftrom X-ray diffraction data. o= stoichiometric end-member Mg-clinochlore. Chlorites associated with marine of four clinochlore samples from Retsof, N.Y.; =chlorites from evaporites: .=average other marine evaporite localities: I = clinochlore from the German Zechstein, KonigshallHindenburg; 2 Zpenninite from the German Zechstein, (?) Werra; 3 =clinochlore from an altered melaphyre, Austrian Haselgebirge, Hallstatt. (After Bodine and Standaert, 1977, p. 64, fig. 4.)

glass, deposited within the lake basins, with the lake brine. A zonal arrangement of the various diagenetic products reflects lateral salinity gradient. The peripheral areas are characterized by unaltered glass. They border on a zeolite zone, to be followed by analcime, and, in the very center, by potassium feldspar (Eugster and Hardie, 1978). This zonation can be observed in the Pleistocene Lake Tecopa, California (Fig. 3-7). In Lake Magadi, studied by Surdam and Eugster (1976), the Na-zeolite erionite is the most common diagenetic product. Depending on the nature of the dissolved salts, other zeolites, such as clinoptilolite or phillipsite have been observed to form in other lake sediments. The conversion of erionite into analcime has been represented by Surdam and Eugster as follows:

3 H,O Nao.5Ko,5Si3.509.

+ 0.5 Na+

+

Magadi-erionite

NaAlSi,O,. H,O

+ 0.5 K t + 1.5 SiO, + 2 H,O

molcimr

Analcime, that represents an environment of higher salinity, may also

137 116' 20'

116"lO'

Fig. 3-7. Mineral zonation in the tuffs of Pleistocene Lake Tecopa: dots ( I )=fresh glass; circles (2)=zeolites; lines (j)-K-feldspar. (After Eugster and Hardie, 1978, p. 254, fig. 10.)

form directly from volcanic glass, or from Na-Al-silicate gels. In the most saline environments, K-feldspars replace analcime as the common diagenetic product. Other common diagenetic products of saline alkaline lakes are bedded cherts. These appear to have formed from a sodium-silicate precursor, magadiite [NaSi,O,,(OH), - 3 H,O], first described by Eugster (1967) from Lake Magadi, Kenya. Some evidence appears to suggest that claymineral diagenesis may also take place in saline lake environments. Oolites composed principally of stevensite in the Eocene Green River Formation from Central Utah are believed to have formed diagenetically by precipitation from solution in the ancient Green River Lake (Tettenhorst and Moore, 1978). Earlier, diagenetic trioctahedral smectite has been reported to occur in the same formation (Dyni, 1976). The Mg-rich clay minerals such as Mg-montmorillonite, hectorite, vermiculite, and mixedlayers ( 14M-14c, 14,,,-14$,, and 14v-14c) are attributed to diagenetic neoformation in clayey Neogene sediments deposited in saline lacustrine environments in Turkey, in association with borates, silicates, and carbonates (Ataman and Baysal, 1978). The possibility of a smectite + illite conversion in the Tilton Shale Member of the Eocene Green Rtver Formation has been considered by Tank (1969). The vertical clay mineral distribution pattern in two sediment cores taken from two East African lakes distinctly suggests that illite may have formed diagenetically by the reaction of a smectite precursor with highly saline lake brines (Stoffers and Holdship, 1975; Stoffers and Singer, 1979; Singer and Stoffers, 1980). Existence of saline paleobrines is evidenced by

138 the presence of zeolites and protodolomite in sediments. In one lake sediment, illitization of smectite appears to have been favored by an unusually high K/N ratio in the lake water. The diagenetic illitization of smectite in the second lake sediment may have resulted from the parallel process of analcime formation, with the necessary K supplied by the volcanic glass + K, Na-zeolite + analcime conversion process (Fig. 3-8). Diagenetic palygorskite and sepiolite Whereas many palygorskite and sepiolite occurrences in marine and lake sediments can be traced back to continental origins, some are beyond doubt of diagenetic origin. The diagenetic formation of palygorskite in marine sediments has been documented extensively by: (1) Weaver and Beck (1977) for the Miocene deposits of the southeastern United States; (2) Couture ( 1977) and Church and Velde (1979) for Pacific Ocean occurrences; and (3) Timofeev et al. (1977) for the Atlantic Ocean sediments. The subject has been reviewed by Singer ( 1979). Smectite and volcanic glass are frequently mentioned as precursor minerals for palygorskite and sepiolite formation. In other occurrences the minerals appear to have formed by precipitation from solution. Requirements for their diagenetic formation appear to include high (alkaline) pH, high Si and Mg activity and low A1 activity. These requirements possibly are met in specific situations, such as: (a) near sites of hydrothermal activity; (b) in peri-marine, shallow-water environments, close to land-masses undergoing intensive desilicification by weathering; or (c) in response to fluctuations in ocean-water temperature that affect solubility levels of limiting chemicals, such as Si. The formation of palygorskite and sepiolite probably extends from the pre-burial stage well into the shallow-burial stage. The formation of glauconite “La glauconie est caracteristique du milieu marin: ceci appartient a SMECTITE GLASS

SMECTITE INTERLAYERS ILLITE

ILLITE

SOLUTION CHABAZITE ERlONlTE

ANALCIME

b

INCREASING SALINITY

AND

ALKALINITY

Fig. 3-8. Diagrammatic representation of the processes leading to the diagenetic illitization of srnectite and the formation of analcime from volcanic glass, with increase in water paleosalinity and paleoalkalinity of Lake Manyare, East Africa. (After Singer and Stoffers, 1980.)

139 l’alphabet de tout geologue” (Millot, 1964). The most familiar and characteristic product of halmyrolysis is glauconite. It forms during the pre-burial stage of diagenesis and can be found today in many oceans at a depth of about 20-700 m in areas with a decelerated rate of sedimentation, such as, for instance, on the outer edge of the shelf in the Gulf of Mexico and of Trinidad. According to Seibold (1964), the green muds off Guinea at a depth of about l00m can contain up to 50% glauconite. Since the Paleozoic period, glauconite is found in almost all formations. The major mineral in glauconite pellets is an iron-rich, mixed-layer illite-smectite, analogous to aluminous illite-smectite. The nature of the interlayering varies with the properties of the layer types, i.e., from randomly interstratified to highly ordered. Glauconite differs from aluminous illitesmectite in that glauconite contains considerably less potassium per illite layer than does aluminous illite-smectite with the same proportion of illite layers, except near the pure illite composition (Thompson and Hower, 1975). Glauconite pellets frequently contain two textural forms of glauconite (aggregate and oriented), which differ in both mineralogy and chemistry. The aggregate glauconite composes the bulk of most pellets and consists of crystals in a nearly random arrangement. The oriented glauconite occurs primarily as rims on the periphery of pellets and has a honeycomb-like structure (Odom, 1976). Glauconite is formed diagenetically in marine sediments in reducing environments, primarily in fossil-rich carbonate sediments. Many processes have been proposed for the formation of the various mineralogical forms of glauconite. The most widely accepted is the “layer-lattice” theory developed by Hower (1961) after Burst (1958). The process of glauconite formation involves the absorption of K and Fe by any degraded detrital layer-silicate structure under suitable chemical and physical conditions. During the “glauconitization” process, the number of expandable layers is reduced and after considerable time the mineral glauconite is produced. According to Birch (1979) and Birch et al. (1976), illite clay-minerals are the precursors proto-glauconite") for glauconite pellets from the continental margins off the west coast of South Africa. Their data suggest that Fe is emplaced into the clay structure very early in the glauconitization process, possibly by a mechanism which is independent of, and prior to, the fixation of K . Odom (1976) suggested that many forms of glauconite pellets grow by the development of smectite in successive stages or layers, which then become oriented glauconite and later aggregate glauconite. The initial growth of smectite is probably related to decaying organic material in an environment with favourable pH, Eh, and physical conditions. The chemical composition

140 of the crystallized smectite and the rate of its development during successive stages of growth might be somewhat variable in different environments. Kohler and Koster (1976) cited the following proposed precursors for diagenetically formed glauconite: (a) amorphous gels, (b) clastic rock and mineral relics, (c) biotite and other micaceous minerals, and (d) montmorillonite. The detailed chemical and mineralogical investigation of glauconite from ten Cretaceous sediments leads the authors to conclude that the most likely precursors for glauconite formation are metal hydroxides in the form of gels, in addition to amorphous silica of primarily organic origin. Formation of glauconite from clastic material is not likely, because of the very minor amounts of these materials commonly associated with glauconite. Micaceous minerals or montmorilloni te would have to undergo drastic crystallographic transformations in order to change into glauconite and are, therefore, also ruled out as glauconite precursors. Possibly, the nature of the initial precursor materials, either iron-rich or iron-poor, are not of great importance for the “glauconitization” process. According to Velde and Odin (1975) and Odin (1978), pelletal glauconites are the result of the imposition of a chemical gradient upon sediments in a distinct spatial localization. The evolution of sedimentary environment towards that which is favorable for the formation of glauconite necessitates high iron contents and increasing K contents as more illite-like phases are produced. As a result, physicochemical properties of the pelletal aggregate, such as Eh and porosity, may constitute important factors. On the other hand, no distinct “parental” relations exist between the detrital precursor minerals and the diagenetic glauconite mineral. Formation of marine zeolites Bonatti’s (1963) studies in the Pacific show that phillipsite “is one of the most abundant mineral species of the upper layers of the earth’s crust”. In extensive areas of the Pacific Ocean, phillipsite concentrations are greater than 50%. Zeolites are among the most important diagenetic minerals in fine-grained pelagic sediments. Phillipsite and clinoptilolite are two common zeolites in deep-sea sediments. Another zeolite which occurs in significant amounts in deep-sea sediments, is analcime. Phillipsite is associated with argillaceous, volcanic, and siliceous sediments. It occurs in areas and sediments having slow sedimentation rates, at shallow depths in the sediment, and in very young sediments; it is most commonly associated with smectite (Stonecipher, 1976; Kastner and Stonecipher, 1978; Houghton et al., 1979). Deep-sea phillipsite from several Pacific core sediments appears to form near the sediment-water interface, incorporating elements from adjacent

141 mineral particles and interstitial sea water. Concentrations of Na, K, Rb, Ca and Sr that are quite uniform with sediment depth, suggest a common origin from the deep-sea alteration of volcanic material by pore water that contains excess amounts of these soluble cations (Bernat and Church, 1978). Phillipsite represents one of the principal products of the halmyrolytic alteration of basaltic glass, accompanying smectite (Honnorez, 1978). The initial stage of palagonitization of basaltic glass is characterized by the crystallization of an intergranular Na, K-phillipsite with saponite low in Ca and, probably, K-rich. During the mature stage of palagonitization the palagonitized glass granules are replaced in situ by intergranular phillipsite and, possibly, very minor amounts of Fe-rich saponite or Mg-bearing nontronite. At the final stage of palagonitization, the hyaloclastite has been completely replaced by an intimate mixture of authigenic K, Na-phillipsite, with almost no Ca, smectite, and Fe-Mn oxides. While phllipsite is dominant in younger deep-sea sediments, clinoptilolite is most abundant in Eocene and Cretaceous samples (Boles and Wise, 1978). Over 80% of the clinoptilolite occurrences are reported in brown clays and microfossil-rich sediments rather than in the ash beds, suggesting that volcanic ash is not a prerequisite for its deep-sea formation. Possibly, phillipsite forms as a metastable, silica-deficient phase in marine pore fluids and is eventually replaced by clinoptilolite. Time appears to have a more important control on this reaction than burial depth of silica concentration in pore fluids. Some clinoptilolite may also have formed from the dissolution of siliceous microfossils and clay minerals in the absence of volcanic glass. Houghton et al. (1979) noted the frequent association of clinoptilolite with alkali feldspar in the western North Atlantic. Nathan and Flexer (1977) compiled DSDP data on clinoptilolite occurrences and came to the conclusion that the mineral is more abundant in certain stratigraphic periods, namely Late Cretaceous, Eocene, and Miocene, that it is rarer in earlier and later periods, and that half of all recorded occurrences are concentrated between Late Cretaceous and Eocene (Fig. 3-9). This distribution is worldwide. Clinoptilolite occurs in deep-sea sediments as well as in shallow-water sediments and reflects, according to these authors, a warmer climate during this time period, which raises the soluble Si levels in the sea water. Clinoptilolite appears to be able to form from both volcanic and nonvolcanic precursors. In South Atlantic deep-sea sediments (cores), clinoptilolite occurs in association with three dominant sediment components, i.e., volcanic deposits, biosiliceous deposits, and turbidi te clays, each of which may provide a source of silica for diagenetic formation (McCoy et al., 1977). Volcanic material as precursor for clinoptilolite formation is dominant in the Cretaceous deposits. Diagenesis of volcanic material proceeds through de-

142

n

5

z

40

30 20 10

n

Fig. 3-9. Age distribution of clinoptilolite in D.S.D.P. Legs I to XXXIII (excluding Leg XXXII). (After Nathan and Flexer, 1977, p. 851, fig. 5.)

vitrification of volcanic glass. In Eocene time, biosiliceous material gained in importance. Biosiliceous materials act as precipitation cores and are being partially altered into euhedral zeolite crystals. Although the terrigenous association became more important during Late Tertiary time, it does not appear to be very significant. The Si for the formation of zeolites is probably supplied through degradation of clays. According to Petzing and Chester ( I979), the dominant factor controlling zeolite formation is the Si/A1 ratio of the precursor. The low-silica zeolite, phillipsite, is formed by the rapid breakdown of basic alkaline glasses and, therefore, is more abundant in young ocean sediments, whereas the high-silica zeolite, clinoptilolite, is formed by the slow breakdown of acidic glasses and is, therefore, more abundant in older sediments. There is a good correlation between the spatial and temporal distribution pattern of clinoptilolite and phillipsite and those of subaereal acidic and basic volcanicity. In addition, the dissolution of siliceous organisms in sediments may result in local silica enrichments in sea water, which will increase the Si/AI ratios and, thus, promote the formation of clinoptilolite. In the Recent sediments of the Gulf of Naples, Italy, analcime was found by Miiller (1961) together with newly formed opal, quartz, and clay minerals, which owe their origin to the halmyrolitic transformation of the sediment mainly composed of volcanic glass. Analcime had also been observed by Norin (1953) in sediment cores of the central Tyrrhenian Sea, which contained ash layers. In comparison to the predominantly basaltic composition

143 of the Pacific pyroclasts, these glasses have a trachytic-leucitic chemistry. This could possibly be the reason for the formation of analcime rather than phillipsite. A frequently encountered vertical zonation of zeolites in marine sediments is explained by Iijima (1978) by burial diagenesis. The zones are distinguished mineralogically on the basis of the reaction series of silicic glass + alkali zeolites albite (Fig. 3-10). Zone I is characterized by the presence of silicic glass which is partly altered to montmorillonite and opal-A or opal-CT. Zone I1 is characterized by the reactidn of silicic glass with interstitial water to form alkali clinoptilolite, alkali mordenite, opal-CT, and montmorillonite. The transformation of clinoptilolite and mordenite into analcime characterizes zone 111. In zone IV analcime transforms into albite. This downward succession of authigenic minerals in general represents decreasing hydration with depth. The precursor zeolites commonly persist as relics, especially in younger Tertiary sediments. +

Alteration of basalt and pyroclastics Iron-rich saponite, nontronite, montmorillonite, celadonite, mixed-layered

ILLITE PREHNITE

I

PWPELLY I TE

I

Fig. 3- 10. Schematic diagram showing zonal distribution of authigenic zeolites and silicates in silicic volcanic sediments in a thick column of marine deposits due to burial diagenesis. The zoning is based on the reaction series: silicic glass- alkali clinoptilolite and mordeniteanalcime-albite. Zone I11 is subdivided into two subzones on the basis of the reaction: heulandite- laumontite. Prehnite and pumpellyite may be in the metamorphic regime. (After Iijima, 1978, p. 179, fig. 3.)

144

chlorite-smectite and chlorite are among the more common alteration minerals produced during the halmyrolytic alteration of volcanic material. They are often localized in veins and vesicles. In addition to clays, alteration products often include zeolites (clinoptilolite and phillipsite), iron and manganese oxides, calcite, and minor amounts of K-feldspar. One of the common but ill-defined products of the halmyrolytic alteration of basaltic glass is palagonite, which essentially consists of a devitrified, hydrated glass in incipient stages of alteration into layer-silicate structures and zeolites. Phillipsite and smectite are the most common diagenetic minerals of the palagonitization process (Honnorez, 1978). From the study of DSDP Leg 34 basalts, Bass (1976) suggested two distinct diagenetic environments: ( 1) Non-oxidative diagenesis is characterized by relatively low oxidation rates, which produce assemblages of saponite, chlorite-smectite, talc and minor celadonite. Most of the Fe3+ is incorporated into silicate lattices and only a small amount is available for discrete oxide phases. (2) Oxidative diagenesis is characterized by the abundance of ferric oxides (goethite, limonite, and hematite) and celadonite; smectite and chlorite are less important and are frequently stained by iron oxides. Oxidative zones commonly occur in the upper few centimeters of each cooling unit, apparently due to direct contact with superjacent bottom waters. Oxidized phases can also be found, however, in lower portions of individual flows, filling veins and vesicles. An iron-rich saponite and accompanying celadonite are the dominant sheet-silicates in veins within the basalt cored during DSDP in the Nazca plate (Seyfried et al., 1978). Oxygen isotopic analyses of these clays and associated calcite indicate a formation temperature of 25 "C. Celadonite is commonly associated with goethite and hematite, suggesting that this phase was formed by precipitation within a dominantly oxygenated environment of components leached from basalt and also those provided by sea water. In contrast, iron-rich saponite, containing significant A1203, appears to have precipitated from a non-oxidizing, distinctly alkaline fluid containing a high Na/K ratio relative to unmodified sea water. Sea-water-basalt interaction at low temperatures, resulting in the formation of celadonite and smectite, may explain chemical gradients observed in the interstitial waters of sediments overlying basalts. On the basis of observed oxygen-isotope, C a z + , M g 2 + , and K + gradients in interstitial waters as a function of depth, Kastner (1976) proposed the following alteration sequence for basalt from DSDP Site 322: (1) At shallow burial depth, basalt releases C a 2 + , Mg2+ and iron. Iron precipitates locally as iron hydroxide and a residual amorphous phase enriched in silica develops. (2) In a second phase, smectite starts to form and consumes the Mg2+ that is released during pyroxene weathering and, possibly, some additional Mg2+ from sea water. Thus, depletion in Mg2+

145 and Ca2+-enrichment in the interstitial water results. Phillipsite formation is most probably responsible for the observed K depletion in the interstitial waters. A downhole decrease in I8O, Mg2+ and K + , an increase in Ca2+ content, and a low 87Sr/86Srratio of 0.7067 in the pore fluids of DSDP Site 323 were interpreted by Lawrence et al. (1979) as having been caused principally by the alteration of volcanic material into smectite, potassium feldspar, clinoptilolite, and calcite. A significant portion of the alteration of ash in the basal sequence must have'occurred before the deposition of the thick sequence of upper sediments. Mass balance considerations and the low 6 l 8 0 values of most of the alteration products suggest that much of the later alteration occurred progressively over the last 13 m.y. Iron- and manganese-enriched clays and even discrete metal oxide phases have probably been produced during the submarine alteration of volcanics without the intervention of hydrothermal processes (Natland, 1973; Jenkyns and Hardy, 1976). The most common halmyrolytic alteration products of pyroclastics are smectites, frequently accompanied by zeolites. Late Cenozoic ash deposits cored in Leg 19 of the Deep Sea Drilling Project in the far northwest Pacific and in the Bering Sea have altered to bentonite beds (Hein and Scholl, 1978). The bentonite beds are composed of authigenic smectite and minor amounts of clinoptilolite. A significant part of the Neogene volcanic history of land areas adjacent to the North Pacific is represented by these diagenetic deposits. Inasmuch as the ash beds become more highly altered with increasing depth of burial, formation of bentonite reduces the number of ash layers in the older deep-sea sediments. The amount of illite layers in the illitesmectite mixed-layer phase is used to distinguish diagenetic from detrital smectites. Diagenetic smectite has less than 15% interlayered illite, whereas detrital smectite contains 45%-55% illite interlayers. Apparently, silica is conserved when silicic ash transforms into smectite and minor amounts of clinoptilolite. The chemistry of bentonite beds reflects the chemistry of their parent ash. The diagenesis of silica in argillaceous sediments The alteration of biogenic opaline silica or opal A (primarily diatoms, radiolaria, sponge spicules and silica-flagellata) and the formation of opal-CT and quartz deposits are among the major diagenetic processes occurring in marine sediments. The studies of Jones and Segnit (1975) on opaline silica have revealed at least three distinct categories of opal, among which opal-CT is the most abundant mineral in common opal and in deep-sea chert. Studies of the structure of opal-CT and its relationship to the other silica minerals show that opal-CT has low cristobalite content and is unidimensionally disordered by interstratified stacking of tridymite layers.

146 According to the “maturation hypothesis” applied to marine cherts in pelagic sediments by Wise and Weaver (1974), disordered cristobalite is the first form of silica which is inorganically precipitated. Within periods of up to 100 m.y., this silica is then converted to a-quartz. Alternatively, the “quartz precipitation” hypothesis proposed by Lancelot ( 1973), states that the mineralogy and porosity of the host sediment is critical for the evolutionary path. A high initial content of impurities, such as exchangeable cations, will tend to retard the internal ordering of disordered cristobalite and hence delay the final timing of conversion to quartz. Consequently, the conversion to quartz will take place first in porous, impurity-free calcareous or siliceous oozes and only later in densely packed fine-grained argillaceous sediments. The presence of carbonates may also serve to chelate inhibiting cations such as iron and magnesium and, thus, to enhance the rate of solution of siliceous tests. Evidence from deep-sea sediments supports the following diagenetic maturation sequence: opal-A (siliceous ooze) + opal-CT (porcellanite) chalcedony or cryptocrystalline quartz (chert). The transformation of opal-A to opal-CT and then to quartz is primarily controlled by temperature, time, specific surface area of sediment, pore water chemistry (pH, types and concentrations of ions in solution and degree of saturation with respect to silica), sediment permeability, and the nature of the host sediment, whether primarily calcareous, argillaceous, or siliceous. Whereas most investigators agree that the opal-A to opal-CT transformation is a solution-reprecipitation reaction, opinions are divided as to whether the opal-CT to quartz transformation is also due to solution-reprecipitation or to solid inversion (Hein et al., 1978). Experiments carried out by Kastner et al. (1977) show that the transformation rate of opal-A to opal-CT is much higher in carbonate than in clay-rich sediments and that opal-CT lepisphere formation is aided by the precipitation of nuclei with magnesium hydroxide as important component. In carbonate-rich sediments, the dissolution of carbonate provides the necessary alkalinity, whereas the sea water supplies the magnesium cation necessary for the precipitation of magnesium hydroxide in the nuclei. In contrast, in clay-rich sediments the clay minerals impede the opal-CT formation by competing for the available alkalinity from sea water. As a result, the clays are enriched in magnesium and the rate of opal-CT formation is strongly reduced. This may serve as an explanation why in Mesozoic clayey sediments opal-CT predominates, whereas in carbonate sediments quartz is most common. The Troodos Massif of Cyprus is overlain by a variety of cherts in pelagic chalks, volcanogenic sediments, radiolarites, and radiolarian mudstones, all of Campanian to Upper Eocene age (Robertson, 1977). Most of the +

147 chalcedonic quartz is derived by recrystallization of previously inorganically precipitated disordered cristobalite rather than by direct precipitation. With time, internal solid-state reorganization of the disordered cristobalite is accompanied by gradual expulsion of impurities until its dissolution which is followed by quartz precipitation. Complete conversion of disordered cristobalite to quartz takes place first in porous calcareous sediments free of impurities, as in Cyprus calciturbidites. In fine-grained clay-rich sediments, such as Cyprus radiolarian mudstone, disordered cristobalite persists much longer. In continental-margin sediments off northwest Africa, no siliceous skeletons are preserved as opal-A in the pre-middle Miocene sediments studied by Von Rad et al. (1977). Siliceous organisms have been either dissolved or converted into opal-CT or quartz (Fig. 3-1 1). Eocene opal-A radiolarian skeletons have been replaced in situ by opal-CT crystallites, whereas diatoms and sponge spicules are always replaced by quartz. Calcite tests of foraminifera and mosaic cement in their chambers are replaced by quartz and not by opal-CT. These transformations represent the early diagenetic stage. During intermediate and late diagenesis, silicification fronts develop, involving the replacement of opal-CT by quartz. A progressive diagenetic maturation of the opal-CT structure within the porcellanite stage is suggested by the positive correlation of the opal-CT 4.1 A spacing with burial depth. This strongly suggests that maturation with time and increasing temperature is the major driving force for the chertification in the clayey and carbonate environments. As indicated by porcellanite relicts, all typical quartz cherts were developed by gradual maturation of porcellanites and not by direct precipitation, without the opal-CT precursor. The rate of opal-CT quartz transformation is much slower in the clayey sediments than in a carbonate environment. In the southern part of the Bering Sea, diagenesis of diatom frustules follows a series of changes that are related primarily to temperature, which is controlled by the depth of burial and local geothermal gradient (Hein et al., 1978). During the first 300-400 m of burial, frustules are fragmented and undergo mild dissolution. Upon reaching a depth of 600m, dissolution of opal-A (biogenic silica) becomes widespread. Silica reprecipitates abundantly as inorganic opal-A at depths between 600 and 700m, which is rapidly transformed by crystal growth to opal-CT. This results in the formation of silica-cemented mudstone and porcellanite beds. Opal-A is transformed to opal-CT at temperatures ranging from 35°C to 50°C. Deposition of at least 500m of diatomaceous sediment was required before the temperature at the base of this section was appropriate for silica diagenesis to occur. Reexamination of the results obtained experimentally by Mizutani ( 1977) showed that amorphous silica changes to quartz through an intermediate +

Silica Sources

.

I

I

I

Weakly

miqrotion of solutions

silicif.

Sediments

replacement of caIci1ic tests 1oIso in pores?

precipitation of dense op-CT inmatrix and of l e p i s p h e r e s in pores

in-situ replacement

[Qtzl I time -dependant ordering of opal-CT l a t t i c e

Porce -

2

lanites

I

I

calcar envir

clayey

envir op-CT latticeof ' metostoble l e g by cotions E low permeab 1

replaced op-CT lepisph

T

q t z transformation by dissol Iprecipil opal-CTin microenvironment

Opal- CT q t z nuclei

in situ re-

re1 fast

Cherts

Fig. 3- 1 1. Schematic diagram showing transformations and diagenetic processes for opal-A, opal-CT, and quartz. The vertical position of the silica phases within the diagram is correlated with their diagenetic evolution (right-hand column). A direct formation of quartz cherts without a porcellanite precursor was not presented. (After Von Rad et al., 1977, p. 894, fig. 6.)

phase of opal-CT, and that the (101) spacing of cristobalite progressively decreases from 4.10 A to 4.05 A. The rate of spacing-decrease increases with increasing reaction temperature. This spacing change represents ordering of opal-CT with the passage of time. Accordingly, it is suggested that stratigraphic boundaries should be parallel to isopleths of d (101) spacings, except in cases where folding had occurred. During the transformation of volcanic glass and the resulting formation of zeolites and clay minerals, large quantities of silica must be liberated, because the alteration products are always poorer in silica than the initial material. If the transformation of pyroclasts in sea water takes place very

149 quickly, all the silica liberated cannot be dissolved in the sea water. Consequently, there is at least a partial fixation of SiO, in the sediment. The sediments in the Gulf of Naples (Muller, 1961) can be mentioned as an example of recent halmyrolitic authigenesis of SiO, -minerals from finegrained glassy pyroclastic particles. The latter are mainly derived from the volcanic tuffs surrounding the coasts of the Gulf. Mineralogic and X-ray examinations showed that in the volcanic glass particles, with a decreasing grain size the content of newly formed quartz and chalcedony increases. In the 0.02-0.002-mm fraction of one sample it amounted to 16.2%. The highest content in the whole sample was 10%. The content of opal present could not be quantitatively determined; however, it is probably considerable. The grain size of the sediments decreases more or less continuously from the coast to the middle of the Gulf of Naples (Muller, 1961, 1967); on the other hand, the quartz and chalcedony contents increase (Fig. 3-12). Inasmuch as large quantities of SiO, are also released during the transformation of feldspars into clay-minerals, SiO, -minerals might also be formed in a similar way. It is probable, however, that this process only plays an important role during the later stages of diagenesis and is important for cementation.

. . .. .. .. . . . ..'.... . .:.. ..." ::'

.:. :. . .: . :. . . : ... .. .I ','

. ..

'

'

Fig. 3-12. Quartz content in Recent sediments in the Gulf of Naples, Italy. (After Muller, 196 1 , p. 7, fig. 2.)

150

Formation of manganese nodules Among the most unusual neoformations on the sea floor, are the concentrically structured manganese nodules, which cover enormous deep-sea areas in the Atlantic, Indian, and Pacific oceans. An extremely low sedimentation rate exists in these areas. All the facts indicate that these are not concretions which have been uncovered by water action from older strata, but are nodules which have grown on the surface of the sediment. The mineralogical composition (at least three different manganese minerals) and the chemistry are very complicated and differ even locally. The major mechanisms for their formation that have been proposed, include: ( 1) inorganic precipitation from normal sea water, (2) precipitation from volcanic hydrothermal solutions, and (3) diagenetic concentration at the sea water-sediment interface. Enrichment of manganese in surfaces of marginal sea floors is caused by the upward migration of Mn’+, due to a diffusive gradient, from anoxic sediment layers, and its precipitation in the oxic surface sediments (Elderfield, 1977a). A comprehensive treatment of the subject appeared in a book edited by Glasby (1977). Changes during the shallow-burial stage Changes in the chemistry of interstitial water As soon as a sediment layer at the depositional interface is covered by a younger layer, the physicochemical conditions in the interstitial water change. In particular, hydrogen ion concentration and redox potential are influenced. Emery and Rittenberg (1952) in their study of the sediments from the basins off the southern California coast found that the topmost sediment layer had an average pH of 7.59 compared to an average of 7.52 for the bottom water. This slight increase commonly continued downwards with pH values ranging u p to 8.5 at a depth of about 8 m; most of the values were between 7.5 and 8.0. This downward increase of pH was not observed by Siever et al. (1965), who determined the pH in several hundred samples of modern oceanic sediments in 22 cores from six general areas in the Atlantic and Pacific oceans. The p H values of core samples were uniformly lower than that of surface sea water, ranging from 7.00 to 7.85, with most values clustering between 7.2 and 7.7. Siever et al. (1965, p.65) state: “The most likely explanation for the p H being lower than sea water is an increase in CO, pressure in the sediment over that in equilibrium with the atmosphere. Corroborating this hypothesis are results of pH measurements on squeezed waters brought into equilibrium with the atmosphere by aeration; uniformly these measurements gave values 8.1-8.2, normal for sea water.” The increase in CO, pressure in the sediments is most likely the result of bacterial

151 oxidation of organic matter. The indications are that this effect is still operative at a depth of 10 m below the ocean bottom. Siever et al. did not determine the sulfate content in their cores. In the sediments of the California basins, Emery and httenberg (1952, p. 789) found a decrease in the sulfate content with increasing depth. In core 5 the sulfate/chloride ratio decreased from 0.13 1 at 0-6 inches to 0.030 at about 70 inches below the surface. At the bottom of the core, 83-87 inches below the surface, no sulfate could be detected. In two other cores changes were not as pronounced.'In most cases, cores from the western Gulf of Mexico (Chave, 1960) also showed a downward decrease in the sulfate/chloride ratio. The decrease of the sulfate content is brought about by sulfate-reducing bacteria in anaerobic environment. If the sediment contains more organic matter than can be decomposed by aerobic processes in the water-containing oxygen (positive Eh values), a depletion of oxygen and, finally, a complete lack of oxygen can be observed. Thus, a reducing environment with anaerobic conditions and negative Eh value would result. This applies both to marine and non-marine environments. The transition from aerobic to anaerobic milieu can take place at some depth in the sediment; however, frequently it is already observed in the uppermost few centimeters of the sediments. In the Black Sea, anaerobic conditions predominate not only in the topmost layer of the sediment but also in the water layers over the sediment. Within the sequence of sediment layers, anaerobic zones can be intercalated among aerobic ones depending on the content of organic matter. Some of the changes in the chemistry of interstitial water can be regarded as a continuation of the processes occurring during halmyrolysis and, therefore, have been discussed in that context. Other changes can be better understood when relating to changes due to deep burial and are discussed later. Only little information is available on the changes of the chemistry of pore solutions of fresh-water sediments in relation to the burial depth. Inasmuch as from the very beginning only very low electrolyte contents were present in these interstitial solutions, the relative changes are probably much greater here. Formation of sulfides During the activity of sulfate-reducing bacteria, H,S is produced which in the presence of dissolved iron or Fe-hydroxides is transformed into black hydrotroilite (FeS . n H,O). This explains the typical black coloring of these sediments. After a very short time, stable iron sulfides (predominantly pyrite) are formed from the unstable hydrotroilite in clayey sediments (Love, 1964).

152 The formation of hydrotroilite is not limited to marine sediments; many sediments in the fresh-water lakes (for instance, Lake Constance) locally contain several percents of this sulfide. In the Jurassic Oxford clay of central England, calcite-rich concretions include pyrite (Hudson, 1978). Concretion growth in pelleted anaerobic mud proceeded concurrently with bacterial sulfate reduction and pyrite precipitation. During post-compactional growth, 6I3C increased and pyrite content decreased, indicating a decrease in organic influence. Framboidal pyrite in clayey Holocene sediments and in Jurassic black shale in the northwestern part of the Barents Sea had formed diagenetically (Elverhafi, 1977). Pyrite formation may also be associated with clay-mineral diagenesis. The close textural association between chlorite and pyrite in two representative shales suggests a common diagenetic origin (Sever and Kastner, 1972). The geochemistry of this diagenetic pair indicates a Fe-Mg exchange in the clay mineral associated with sulfate reduction. In addition to pyrite and marcasite, the formation of which is favored by lower pH, very small quantities of other heavy-metal sulfides (mainly sphalerite, galena, and chalcopyrite) occur in clayey sediments of older formations. In pyrite-rich, bituminous, dark clay-marls of different Mesozoic strata, Miiller ( 1955) and Haussiihl and Miiller ( 1963) observed numerous idiomorphic wurtzite crystals which were formed during the early diagenetic stage. The conditions necessary for the formation of such heavy-metal sulfides have not yet been completely clarified. In all probability, however, the heavy metals are derived from the surrounding sediment itself and were brought to the formation site in ionic solutions. Changes in clay -mineral composition (marine enuironment) It is to be expected that most of the processes starting in the pre-burial stage also continue during the shallow-burial stage, even if greatly decelerated, because a single mineral grain is now in contact only with a limited amount of water. For example, the cation substitutions and the neoformation of chlorite-sudoite (and probably illite) from montmorillonite continue in the marine environment. The early diagenetic formation of illite in clay-rich sediments, containing volcanic ash, was shown to occur in the Mediterranean Sea by Norin (1953) and Miiller (1961). The nature of the partly worm-like crystals completely excludes the possibility of transportation and they cannot be alteration products of already existing minerals. Their development from ionic solutions seems to be very probable. Changes in clay-mineral composition (continental enuironment) Within the continental environment, argillaceous sediments appear to

153 undergo diagenetic changes due to burial by organic material, such as peat. Among the more significant changes is the relative enrichment in kaolinite. Many underclays and tonsteins are rich in kaolinite. It seems that at least part of that kaolinite had formed diagenetically due to leaching by organic acid solutions originating from the organic matter overlying the argillaceous sediments. Thus, Staub and Cohen (1978) found that the kaolinite/smectite ratio in clay strata covered by some peat deposits in South Carolina is more than 2: 1, whereas in clays below this zone this ratio is below 1. Similar processes may have contributed to the formation of underclays below some Paleozoic coals. Due to occurrence beneath acidic marsh, illite clays in Pennsylvanian strata from Iowa-Missouri, U.S.A., were degraded rapidly to mixed-layer clays with maximum content of expansible layers (Brown et al., 1977). On the other hand, poor ordering of kaolinite in the underclays of Carboniferous sediments in Scotland is attributed to degradation of detrital kaolinite in the acidic swamp environment (Wilson et al., 1972). The enrichment in kaolinite of continental argillaceous sediments has also been ascribed to upward leaching by groundwater (dialysis), as in the case of the Kirkwood Miocene Formation in the New Jersey, U.S.A. (Isphording, 1970) or the Triassic flint clay in the Sydney Basin, Australia (Loughnan, 1970).

Carbonate concretions Mudstones and shales poor in carbonates commonly contain concretionary bodies, mainly consisting of calcite and, to a lesser extent, also of siderite. It seems probable that most of the concretions started to form in the shallow-burial (and early) stage of diagenesis, because the enveloped relics of organisms are commonly not deformed. The concretions frequently contain laminae which may continue into the surrounding shale. In the concretion, however, the laminae are several times thicker, which is indicative of an earlier stage of compaction. Inasmuch as relics of organisms are frequently found in the center of the concretions, it seems likely that organic matter played a role in the formation of concretions. Lippmann (1955) explained the genesis of concretions as follows: Ammonia resulting from the decomposition of organisms or amines gives rise to strongly alkaline environments in the vicinity of the animal (or plant) embedded in the sediment. As the solubility of the carbonates decreases with increasing pH, they are precipitated on the fossil from the interstitial solutions, which have been saturated with carbonates by dissolving the disseminated calcareous material (also present in predominantly argillaceous sediments). As the carbonate concentration of the pore solution decreases in comparison to the surrounding environment, and because of the difference

154 in the concentration, more carbonate is constantly diffused toward the fossil. This process, accompanied by a constant growth of the concretion, continues until the production of ammonia stops, or until there are no more dissolved carbonates available in the vicinity. Petrographic, chemical and carbon isotope data indicate a dominantly organic source for the carbon in calcite-rich concretions from the Jurassic Oxford clay of Central England (Hudson, 1978). Concretion growth in pelleted, anaerobic mud proceeded concurrently with bacterial sulphate reduction and pyrite precipitation. During postcompactional growth, 8 I3C increased and pyrite content decreased, showing waning organic influence. The mineralogical and chemical changes are explained by a replacement of originally marine pore waters, modified by bacterial activity, by water of ultimately meteoric origin. As reducing conditions can predominate in the immediate vicinity of a decomposing organism, siderite is also a possible concretion-forming material. The precipitation of the carbonates takes place in the water-filled pore spaces without the latter changing in size to any marked degree (e.g., by the clay minerals being forced apart). Thus, the water content (and consequently also the porosity) at the time of the formation of the concretion can be calculated from the volume ratio of carbonate to non-carbonate minerals in the concretions. For instance, Lippmann (1955) and Seibold (1964) made such calculations on Lower Cretaceous and Liassic concretions of argillaceous sediments from Hoheneggelsen (northwestern Germany) and WutachSchlucht (southwestern Germany). According to Lippmann, the calculated water content is about 55% corresponding to a porosity of about 75%. Seibold’s values are slightly lower: about 70% porosity. From this, both authors concluded that the concretions must have formed in the uppermost few meters of the sediment. Chunges during the deep-burial stage of diugenesis According to Taylor (1964), with increasing depth of burial hydrogen ion concentration and oxidation-reduction potential decrease in importance in controlling the diagenesis. Effects of temperature and overburden pressure, both on mineral grains and on pore fluids, together with the partial pressures of the main components of these fluids, are likely to be the significant factors during diagenesis. Changes in the chemistry of interstitial water Large amounts of pore fluids are expelled during the compaction and dewatering process of continental and marine sediments. The chemistry of these pore fluids undergoes significant changes during this compaction

155

process. It seems that most of the salts present in the waters, which are trapped during sedimentation, are squeezed out during the initial stages of compaction. According to many investigations, reviewed by Rieke and Chilingarian ( 1974), the salinity of squeezed-out solutions progressively decreases with increasing overburden pressure. Results obtained by Chilingarian et al. (1973) indicate that the concentrations of solutions squeezed out of montmorillonite saturated with sea water go through a maximum, or at least remain constant, before starting to decrease with increasing overburden pressure (Fig. 3-13). The conclusion of Chilingarian et al. (1973) that the concentration of expelled solutions during the initial stages of compaction is slightly higher than the initial interstitial fluid was not confirmed by Rosenbaum (1976) who examined the variation in chemistry of pore fluid expelled during compaction of a montmorillonite material saturated with distilled water. The initial concentration of all the analyzed ions (K+ , Na' , Ca2+, Mg2+, SO:-, Cl-) decreased rapidly with increasing stress during

38,400 0

I

1

20 40 Amount of extruded solution ( m l )

3

Fig. 3- 13. Variation in the total dissolved solids content with increasing compaction pressure in subsequent fractions of extruded solutions from montmorillonite clay saturated in sea water. (After Chilingarian et al., 1973, p. 397, fig. 4.)

156 the initial loading and thereafter the rate of decrease declined markedly (Figs. 3-14a, b). This indicates that the decrease in concentration of dissolved solids in the pore fluid is related exponentially to the overburden pressure. The observed concentration decreases can be accounted for on the basis of the double-layer theory for highly charged clays, such as smectite. For these clays, the double-layer theory predicts that the electrolyte content of the expelled interstitial waters should decrease when the interaction between the diffuse layers begins to occur. This negative “adsorption”, taking place already during the initial stages of compaction, is caused by the preferential expulsion of ions having the same sign of charge as the colloids, when compaction brings the latter closer to each other. A refinement of this theory, using a model based on Donnan equilibrium, was recently offered by Appelo ( 1977). Not only the total electrolyte content of pore fluids undergoes changes during compaction and burial, but also the electrolyte composition. DSDP samples of interstitial solutions of deeply buried marine sediments throughout the world oceans have shown that in all but the most slowly deposited sediments, pore fluids exhibit changes in composition upon burial. In pelagic clays and slowly deposited (1 cm/103 yr) biogenic sediments these changes are relatively less pronounced than in the carbonate and biogenic sediments deposited at somewhat greater rates. Compositional changes become significant only at great depth. The major elements, except for SO:-, show little variation down to a depth of 300m. The solids and solutions of these sediments must represent a close approach to equilibrium. In partly consolidated sediment sections approaching igneous basement contact, on the other hand, interstitial Ca2+-enrichment, accompanied by depletion in N a f , Si4+ and CO, were observed (Sayles and Manheim, 1975). These changes are attributed to exchanges of Na for Ca in silicate minerals forming from submarine weathering of igneous rocks such as basalts. Water is also consumed in these regions, accounting for minor increases in total interstitial water salinity. It has been shown that the depth to which diffusional communications with the overlying sea water is possible, is small in rapidly accumulating sediments (50-500 cm/m.y.) on continental shelves, but can be large in slowly accumulating sediments ( < 2 cm/m.y.) (Gieskes, 1975). Any concentration-depth profiles, particularly when they are smooth and continuous, may originate at levels that are deeper than that which was sampled. Gieskes ( 1975, p. 450) concluded: “Therefore, in slowly accumulating sediments ( < 2 cm/m.y.) concentration-depth profiles of Ca, Mg, Na, SO, and other major sea-water constituents can be understood only from information derived from drillholes, preferably ones that have been drilled to the oceanic basement. Then, using information on possible changes with depth of the

157

..

350-

300

Potassium V Sodium A Calcium 8

8

-

.

A

-

5 E

'

A

c

v v

-2500

U

n.

200

A

150

.L

I

I

10

1

100

1

1000

Log e f f e c t i v e a x i a l pressure, kg/cm2

Fig. 3-14a. Variation of K, Na and Ca concentrations in expelled pore fluid with effective axial pressure at 40°C. (After Rosenbaum, 1976, p. 120, fig. 3.)

Magnesium 0 Chloride SulDhote 0

*Ol

8 0

0

l150.

O

15.

0

0

OO 650 51

@,@

e:.'o;sr@

808

400-

0

Fig. 3-14b. Variation of Mg”, C 1 and ~ SO,‘+ concentrations in expelled pore fluid with effective axial pressures at 40°C. (After Rosenbaum, 1976, p. 120, fig. 4.)

158 diffusional transport properties of the sediments, an analysis can be made of the probable reaction sites in the sediments and the elements involved in such reactions. Authigenic material in the sediments does not necessarily result from the reactions suggested by observed gradients and, conversely, suggested reactions must be verified by actual identifications in the sediments. Little information is available on the actual equilibrium state between interstitial waters and the solid sediments. Smooth concentration gradients, however, suggest that such an equilibrium state may not exist and that even the alteration of deeply buried volcanic rocks makes a contribution of dissolved material to the ocean, at least for Ca.” This alteration process may also form a sink for Mg2+ and Na+ , and possibly K + and CO,. On the basis of discrepancies between predicted and observed interstitial water profiles, Kastner and Gieskes (1976) identified two major sites of reactions at DSDP site 323 in the Bellinghausen Abyssal Plain of the southern Pacific Ocean (Fig. 3-15): (a) In the silicification zone, dissolution of plagioclase and a few coccoliths and formation of Mg-rich smectites and K-feldspars are responsible for the observed increases in Ca2+ and decreases in K + and Mg” contents. (b) Below the silicification zone, the observed interstitial water concentration gradients are to be accounted for by the alteration of the basal basalt minerals (pyroxene and plagioclase) and formation of celadonite, goethite, calcite, and smectite. DH

1 ’

meqll 2 3 4 “\

\f

f”

4.’

.

- , 10

u

600 il moles/l NHq 30 m m o Il e s i l S O 4

400 20 ,

t

rnoiesll

I

.

400

1

600

t

Alkalinity

800

Sulfate

I

Ammonia

8001

Silica

400

400

600

200

C

5

600

600

Fig. 3-15. Interstitial water chemistry, site 323, DSDP, Bellinghausen Abyssal Plain. I =clay, silt, and sand; I1 =diatomite, claystone, and porcellanite below a depth of 400 m; I I I claystone: I V = calcareous plankton and Fe-Mn sediments; V =zeolitic clays; W =basalt. (After Kastner and Gieskes, 1976, p. 14, fig. 3.)

159 The differences in the extent of diagenesis associated with basalt weathering at different sites could be due to various factors such as differences in age, the thickness of sediment cover, sedimentation rate, and possible interference of hydrothermal activity. Somewhat similar conclusions were reached by Perry et al. (1976) for DSDP Site 149 in the Caribbean Sea. With increasing depth, the sediment pore-water exhibits a strong depletion in Mg2+ and a corresponding enrichment in Ca2+ content, whereas the alkalinity remains relatively constant. Dissolved SiO, content is nearly constant (6 ppm) in the upper 100 m of the sediment, but is much higher (60 ppm) in the deepest pore waters. A depletion in K content and a decrease in 180/160 ratio with increasing depth were also recorded. The submarine alteration of volcanics to a smectite clay could account for the pore-water gradients (Fig. 3-16). By the alteration process, Ca2+ and Mg2+ are released to the pore water, but Mg2+ is retained by smectite formation. The net reaction allows for the observed relationship between Ca2+ and Mg2+ gradients of interstitial water, with little net change in alkalinity. Another process proposed for the removal of Mg2+ from the interstitial water, is the reaction with clay minerals in anoxic environments according to the following scheme:

3 Mg(S0l"tlo") + 2 Fe(clay) + 4 S

+

3 Mg(clay)+ 2 FeS~(pyrltc)

where Fe(clay) refers to the structural Fe3+ in the clay and Mg(c,av)to the Mg2+ added to the clay structure in the course of the reaction. According to Drever (1971a, b), Mg is derived from interstitial waters supplemented by diffusion from the overlying water body and sulfide is formed from dissolved sulphates by the action of reducing bacteria. From experimental data, Heller-Kallai and Rozenson (1978) recently concluded that direct Mg e Fe exchange in clays, as proposed by Drever and more recently by Elderfield (1977b), seems improbable, but that depletion of Mg in interstitial waters of anoxic sediments may be due to reaction of Mg-containing solutions with partially disintegrated clay. Also another mechanism, proposed by Sholkovitz (1973), which postulates addition of Mg to the exchange sites of clay minerals previously blocked by precipitated ferric hydroxides, appears improbable in view of the findings of Sayles and Mangelsdorf ( 1977) that the proportion of adsorbed Mg on marine clay had been grossly overestimated in the past. Changes in cia-v mineral composition-illite diagenesis The changes in the clay mineral composition during the pre-burial stage and the upper shallow-burial stage can become apparent to a certain extent by means of a comparison of material delivered from a known hinterland

160 SlOZ p M O L E S / L I T E R

K

mMOLES/LITER

rnMOLES/LITER

Fig. 3-16. Vertical concentration profiles of (a) dissolved SiO,, (b) K + , and (c) Ca2+ and Mg2+ in the sediment pore-waters of Hole 149, DSDP. Crosses represent cold sediment squeezes, dots: warm squeezes. (After Perry et al., 1976, p. 415, figs. 3a, b, 4.)

161 with material from the sedimentation basin. In older sediments at greater depths, .this comparison is not possible and, thus, the unknown factors correspondingly increase.

Mineralogy and chemistry. The major mineralogical process occurring during the deep burial of argillaceous sediments is the progressive conversion with increasing burial depth of smectite into illite or into a mixed-layer illitesmectite, with a high proportion of illite layers. This major mineralogical process has been suggested as occurring in association with and being related to other minor changes, such as decomposition of mica and K-feldspar, decomposition of kaolin, and formation of chlorite or chlorite interlayers. In the Eocene to Pleistocene shales of the U.S. Gulf Coast the effects of burial on argillaceous sediments have been studied most extensively (Burst, 1969; Perry and Hower, 1970; Hower et al., 1976). When drawing conclusions from these studies, it must be kept in mind that mineralogical variation may simply reflect changes through time of the mineralogical detritus supplied to the sedimentary basin: For the interpretation of data obtained, this factor must, therefore, be carefully considered. A detailed mineralogical and chemical investigation was made by Hower et al. (1976) of shales from a well in Oligocene-Miocene sediments of the U.S.A. Gulf Coast. Major mineralogical changes with depth take place over the interval of 2000-3500m, after which no significant changes are detectable. The most abundant mineral, illite-smectite, undergoes a conversion from less than 20% to about 80% illite layers over this interval, after which the proportion of illite layers remains constant (Fig. 3-17). Over the same interval, calcite decreases from about 20% (bulk sample) to almost zero, disappearing from progressively larger size fractions with increasing depth: potassium feldspar (but not albite) content decreases to zero, whereas chlorite content appears to increase (Figs. 3-18 and 3-19). Variations in the bulk chemical composition of the shale with depth show only minor changes, except for a marked decrease in CaO, concomitant with the decrease in the calcite content. By contrast, the < 2 p m fraction, consisting of pure illitesmectite, shows a large increase in K,O and A1,0, and a decrease in SiO, content. The atomic proportions closely approximate the reaction: smectite +A13+ K+ = illite Si4+. A similar increase in illite layers with increasing burial depth was observed by Perry (1974) in the Gulf Coast sediments of Louisiana (Fig. 3-20). The measured K-Ar ages of the mixed-layer illite-smectite become progressively younger with increasing depth of burial because of the diagenetic addition of potassium, caused by conversion of smectite layers to illite layers. Smectite decreased in abundance and content of mixed-layer clays increased with increasing depth of shales in the Rhinegraben, southwestern Germany (Heling, 1974).

+

+

162

Percent illite layers

Fig. 3-17. Proportion of illite layers in illite-smectite as a function of depth in the finest and coarsest clay fractions of shales from a well in Oligocene-Miocene sediments of the U.S.A. Gulf Coast. (After Hower et al., 1976, p. 730, fig. 3 . )

The conversion of smectite to illite-smectite during burial metamorphism is reported from Middle Cambrian to Middle Ordovician sedimentary rocks of western Newfoundland (Suchecki et al., 1977). Potassium was derived from K-feldspar and/or K-mica in the detrital silt of the shale. Expandable chlorite and corrensite occurring in association with the illite-smectite incorporate Mg probably released during the smectite -+ illite conversion. Clay minerals from the Lower Cretaceous shale outcrops of northeastern British Columbia were investigated by Foscolos and Kodama ( 1974) in order to assess their degree of diagenesis and their oil-generating potential. Crystallinity index (the width of the illite 001 peak at half-height expressed in mm), sharpness ratio (the ratio of the height of the illite 001 peak at lO.0A to the height of the base line at 10.5 A), content (in W ) of 2M illite, polymorph, and presence of discrete minerals were used as indices of diagenesis. Data on the < 2-pm fraction show that the crystallinity index decreases, whereas the sharpness ratio and the 2M illite polymorph content increase with burial depth (Table 3-V). Results on the (0.08-pm size fraction reveal the ex-

163

Percent potassium feldspar

Fig. 3-18. Potassium feldspar content of > 10-pmand 2-10-pm fraction of shales from a well in the Oligocene-Miocene sediments of the U.S.A. Gulf Coast. (After Hower el al., 1976, p. 73 1, fig. 4.)

istence of a three-component interstratified clay mineral. In addition, Fourier transform calculations and chemical and physicochemical analysis indicate that both the ratio of the amounts of nonhydrated clays (illite) to hydrated clays and the K ,O content of clays increase with burial depth; cation exchange capacity and specific surface area decrease with burial depth.

Conversion process. Burst (1969) viewed the conversion process of smectite into illite as essentially a simple dehydration of smectite under the higher pressure and temperature conditions of deep burial. He distinguished three stages : (1) Expulsion of water from the pores at a depth of around 2500 m; the remaining water (about 30%) is present in between the layers where it forms a bimolecular layer. (2) At a depth range of 2500-4500 m, expulsion of one of the water layers occurs from the interlayer. (3) Below a burial depth of 4500m, there is an expulsion of the last layer

1

I

I

9

Percent

chlorite

Fig. 3-19. Chlorite content of the >2-pm and 0.1-2-pm fractions of shales from a well in the Oligocene-Miocene sediments of the U S A . Gulf Coast. (After Hower et al., 1976, p. 731, fig. 5.)

68000 I-

* c. 8,000

0 f0n) 10,000

12.000

14,000

*J

m 16,000

*J 1

I

I

I

20

40

60

80

% illite layers in illite/smectite

Fig. 3-20. Relation between sample burial depth and percentage of illite layers in mixed-layer illite-smectite in <0.5-pm size fraction in Gulf Coast sediments of Louisiana, U.S.A. Arrows denote samples selected for K-Ar studies. (After Perry, 1974, p. 828, fig. I . )

TABLE 3-V Variation with maximum burial depth of analyzed samples from Lower Cretaceous shales of North Eastern British Columbia of: (1) discrete layer-silicates of Na-saturated, <2-pm oriented clays; (2) crystallinity index and (3) sharpness ratio of K-saturated, <2-pm oriented specimen of illites; (4) content (in %) of 2-M illite polymorphs of K-saturated, 2-0.2-pm size fraction of unoriented clay specimen (after Foscolos and Kodama, 1974, p. 322) ~~

Discrete layer silicates

Crystallinity index

Sharpness ratio

Illite [(2 M)/(2 M

illite-chloritekaolinite

26

1.2

20

2217

illite-chloritekaolinite

24

1.2

30

7300

2433

illite-chloritekaolinite

26

1.2

30

278

8OOO

2667

illite-chlorite

19

1.4

30

20 1

8650

2883

illite-chlori tekaolinite

20

1.5

40

109 437

9350 10.100

31 17 3367

illite illite

15 13

1.8 2.2

50 45

Sample No.

Maximum burial depth of analyzed samples

1063

(ft) 5900

(m) 1967

999

6650

823

+ 1 Md)]X 100

166 of water; this last process occurs very slowly and depends on various geological conditions. Meanwhile, it became apparent that the smectite illite conversion is not that simple, with a number of other processes being involved. Low-charge smectite layers do not convert to illite by simply being exposed to a source of soluble potassium (Weaver, 1958). The fixation of potassium ,in the newly formed illite layers involves an increase in the net negative interlayer charge of the expandable layers. This increase can result either from the substitution of aluminium for silicon in the tetrahedral layer or by the substitution of divalent for trivalent cations in the octahedral layer. Powers (1959) suggested that on deep burial Mg2+ migrates into the octahedral layer, where it replaces Al'+, which in turn displaces tetrahedral Si. This causes an overall increase in the layer charge. Perry and Hower ( 1970) enumerated three possible mechanisms for converting expanding into nonexpanding clay: substitution of A13+ for S i 4 + , substitution of Mg" and/or Fez+ for Al" , and/or reduction of F e 3 + . They concluded, however, that there is insufficient evidence to determine which chemical reactions are involved. Eslinger et al. (1979) attributed an increase in FeZ+/Fe3+ ratio with increasing percentages of illite layers in bentonites in the Disturbed Belt, Montana, to a redox reaction involving the oxidation of organic matter. They estimate that iron reduction may have contributed as much as 10-30% of the increase in total structural charge. Foscolos and Kodama (1974) showed that aluminium content in the tetrahedral positions increases with burial depth (Table 3-VI). As diagenesis proceeds, increasing amounts of aluminium substitute for silicon in the tetrahedral layers. In the octahedral position, aluminium content also increases with increasing depth of burial. Iron and, to a lesser extent, magnesium contents decrease with increasing burial. As shown in Table 3-VI, potassium content in the overall structural formulae increases with depth of burial. On the other hand, calcium content decreases with burial depth. Pollard ( 1971) proposed a semi-displacive mechanism for structural changes occurring in the course of the smectite- illite conversion. In the first stage, interlayer Al' enters hexagonal holes. Subsequently, under conditions of low-grade metamorphism, temperatures become sufficient to permit switching of Si-0 to AIL0 bonds, accompanied by rotation of the tetrahedral geometric adjustment of the octahedral layers and migration of protons from O H groups to apycal oxygens. This mechanism, which involves stretching of Si-0 bonds followed by expulsion of Si4+ from tetrahedral positions, requires relatively high energies and is, therefore, not very likely. On the basis of results of experiments involving smectite-alkali-halide reactions, Heller-Kallai ( 1975) proposed a process of irreversible deprotonat i o n of the clay, in the presence of proton acceptors such as alkali halides. +

+

TABLE 3-VI Summary of the chemical formulae for the seven interstratified clays from Lower Cretaceous shales of northeastern British Columbia (after Foscolos and Kodama, 1974, p. 328) Sample No.

Overall layer structure

Non-swelling layer structure

Swelling layer structure

0

B

h

0

0

0

0 h

,

-d8

0 0 0

Vi

-

3

N

-d8

0

167

168 Potassium source. In their study of sediments from the Gulf Coast, Hower et al. (1976) noticed that the variations in the bulk chemical composition of the shale with depth show only minor changes, except for a marked decrease in CaO, concomitant with the decrease in calcite content. By contrast, the < 0. I-pm fraction, consisting of virtually pure illite-smectite, shows a large increase in K,O and A1,0, and a decrease in SiO, content (Figs. 3-21, 3-22). The atomic proportions closely approximate the reaction: smectite A13+ K + = illite Si4+. The potassium and aluminium appear to be derived from the decomposition of potassium feldspar (and possibly mica) and the excess silica probably forms quartz (Fig. 3-23). Hower et al. concluded that the shale during the mineralogical and chemical changes occurring in response to burial diagenesis, acted as a closed system for all components except H,O, CaO, Na,O and CO,. Thus, whereas K,O required for the smectite-illite transformation was taken up in the fine clay fractions, coarse-grained fractions lost K,O (Fig. 3-2 1). Similar decreases of potassium contents in the silt fractions of the Gulf Coast sediments, with increasing diagenetic changes, were reported earlier by Perry and Hower ( 1970) and Perry (1974), who attributed this to the decomposition of coarse micas in addition to K-feldspars. In other words, the potassium is being redistributed between the detrital and diagenetic phases. This supports the additional evidence against the hypothesis that diagenesis proceeds in the

+

+

+

Weight percent K O , 1

2

3

(Ignited,CaO free) 4

5

6

Fig. 3-21. Distribution of potassium among various size fractions of shale samples (ignited, Ca-free) from different depths from a well in Oligocene-Miocene sediments of the U.S.A. Gulf Coast. (After Hower et al., 1976, p. 735, fig. 11.)

169

Weight percent o x i d e

Fig. 3-22. Depth dependence of the A1,0, and SiO, contents of
sediments by extraction of K from sea water (Weaver, 1967a; Weaver and Beck, 1971). This also supports the early opinion of Powers (1 959) that the necessary I

1

I

I

I

I

Percent quartz

Fig. 3-23. Quartz content of various size fractions of shale samples from the same well as in Fig. 3-21. Quartz was not detected in the
170 potassium is not supplied by pore solutions from a distant source outside of the sediment, but is derived from the potassium feldspar within the sediment. Heling (1978) also does not believe that the potassium required for illite diagenesis in Rhinegraben sediments was supplied by fluid movement, considering the very low permeability of compacted sediments. If potassium was carried by pore solutions, illite diagenesis should be expected to be more advanced in those strata where grain size is coarse and permeability is high as compared with fine-grained horizons; however, a reverse relationship is observed. A rough estimate proves that the K,O content of the potassiumfeldspar presently found in the Rhinegraben shales is not sufficient to provide the amount of potassium necessary for the observed illite diagenesis. But considering that the initial K-feldspar abundance (i.e., before the carbonate had replaced 30-40% of the sediment) was greater by about one third, potassium supply and demand could be assumed to have been balanced. The major process involved in the case of the smectite-+illite transformation, according to Heling, is the diffusion of K-ions into the interlayer positions of the smectite minerals. I Kinetics of illite diagenesis. When the charge on the expandable layer-silicate reaches about 0.75 equivalents per half cell, water is expelled from the interlayer positions and potassium is fixed, converting the expandable smectite structure into that of illite (Hower and Mowatt, 1966). Perry and Hower (1970) suggested that the occurrence of that reaction is directly related to temperature. For example, data from a typical Gulf Coast well are shown in Fig. 3-24. Mixed-layer clay from this well initially contains 80% expandable layers (that is, the illite/smectite ratio is 1 :4). Reaction begins at a depth at which the temperature reaches 50°C and ceases at 100°C where the clay contains 20% expandable layers. With 35% expandable layers content an allevardite-type ordering develops. Thus, reaction stages are controlled by burial depth and geothermal gradients. According to data compiled by Dunoyer de Segonzac (1970) and presented in Fig. 3-25, the temperatures of transformation of smectite into illite-smectite range from 70 to 95°C and take place at depths of burial ranging from 1000 m to 2500 m. The depth of burial at which these conversions take place depends also on the geothermal gradients. For example, Heling (1978, 1979) observed that the amount of smectite layers in illite-smectite interlayers from sediments in the Rhinegraben was already reduced at a depth of 700-900 m (at temperatures of about IOO'C), probably as a result of the high geothermal gradients common there (Fig. 3-26). Similarly, formation of illite-smectite interlayers

'

In the opinion of the editors, compaction fluids could have played an important role in supplying K-ion, irrespective of apparent sediment permeability.

171 00

1

I

I

I

In 0

-

>,

-

vC0 0

a x w 20-

01

I

I

I

I

I

Ternperature,'C

Fig. 3-24. Relationship between expandability ( S of expandable layers) and temperature for illite-smectite from well E of Perry and Hower (1970, p. 175, fig. 15.)

0

1000

2000

JOOC

400C rn

Fig. 3-25. The disappearance of montmorillonites with burial. Dotted lines: montmorillonites are transformed into mixed-layer illite-montmorillonite. Circles indicate depth of burial at which transformation begins. (After Dunoyer de Segonzac, 1970, p. 296, fig. 5.)

172

I

1

0

'

W

U

3

tU w

n

3+

1;

It

'

Paleozoic samples Mesozoic samples

I

I ion'

70

I

40

60

eo

100~0

ILLITE CONTENT IN < 2pm FRACTION ,%

Fig. 3-26. Relative illite abundances as a funotion of temperature in sediments from the Rhinegraben, West Germany. (After Heling, 1979, p. 81, fig. 21.)

at burial depths ranging only from 300 m to 1300 m was reported by Matter (1974) for the Arabian Sea. Hydrothermal experiments undertaken by Eberl and Hower (1976) give some indications on the kinetics of illite formation. With potassium as the only interlayer cation, smectite initially synthesized from glass reacted to

173 form randomly oriented illite-smectite and finally ordered illite-smectite. Expansibility was inversely related to the duration of the test. Thus, a sequence of reactions similar to those that appear during burial diagenesis could be followed in the laboratory. The reactions can be summarized as follows: (2)

(1)

glass + 100%expandable smectite + illite-smectite (or pyrophyllite

+ kaolinite

+ quartz)

Whereas the first reaction goes to completion within a period of few days, the second one takes place over a period of months or years, depending on the temperature. The appearance of illite layers in the second reaction can be described with a first-order kinetic reaction. The integrated form of this reaction is:

where a is the initial concentration of smectite (100%)and x is the amount of smectite that has reacted to form illite after t days. The ( a - x), therefore, is equal to the percentage of smectite layers (or the expandibility) present in the total amount of the illite-smectite synthesized, and eq. 3-1 reduces to: In

100 = kt % ' expandable

(3-2)

The first-order rate constant k for a specific temperature is given in terms of (days) I . After determining k graphically and introducing into the Arrhenius equation, Eberl and Hower determined an activation energy for the formation of illite layers from smectite of 19.6 3.5 kcal/mole. The magnitude of this activation energy and the rate constants suggest that the transformation of smectite into illite during diagenesis involves the breaking of chemical bonds in the tetrahedral sheet so that aluminium can substitute for silicon, thereby building a negative charge on the 2: 1 layers. It is not all clear, however, whether the relationship between temperature and expansibility represents reaction kinetics or chemical equilibrium. According to the experimentally established rate constants, the expansibility of the Miocene Gulf Coast clays should be considerably lower than actually determined. The higher expansibilities prevailing would then indicate that illite-smectite is stable and that its expansibility is determined by prevailing P-T conditions rather than by the length of time the clay was reacting. Eberl and Hower (1977) noted that the presence of Nai , Mg2+ and Ca2+ slows down the experimentally determined reaction rates. A higher layer charge was required to produce sodium mica-like layers, which is due to the ~

*

174 higher hydration energy of sodium. The difference in hydration energy between potassium and sodium may account for the fixation of potassium rather than sodium in illite during burial diagenesis. The greater the hydration energy of the cations adsorbed in the interlayer region of expanding clays, the greater the charge that must develop on the 2: 1 layers to cause dehydration. In addition, more A13+ for Si4+ substitution is required with consequent slower reaction rate. That may be, at least partly, the reason why ( 1) the interlayer region of illite concentrates potassium rather than sodium during diagenesis, and (2) shales are enriched in alkali elements of low hydration energy, such as Cs, Rb, and K (Eberl, 1978a). Inasmuch as kaolinite content does not appear to increase with depth, but instead decreases (Dunoyer de Segonzac, 1970), the reactions suggested by Eberl and Hower ( 1976) may be incorrect. Other factors that may affect reaction rates are suggested by Heling (1978), who pointed out that small grain size and high potassium concentrations in the pore solutions shorten the time necessary to complete the transformation. In his study of argillaceous sediments o f ,the Rhinegraben, he concluded that the transformation of smectite into illite primarily depends on the cumulative energy supply, in terms of temperature X time. Morphology of illite diagenesis. Various observations have shown that the behavior of the 2: 1 crystallites in the fine clay fraction of sediments is different from that in the coarse clay fractions. It appears that the smallersized fractions contain more expanding layers, more 1-M, illite, and less 2-M illite than the larger-particle fractions. That suggests that small 2: 1 crystallites have lower tetrahedral aluminium, lower interlayer cation population density, and greater ease of expansion. T h s led Jonas (1975) to relate the smectite -,illite transformation to the simple crystal growth process of the 2 : 1 clay minerals. Figure 3-27 shows the progressive changes brought about as a series of silicate layers increases in areal extent during aggradation. Increased interlayer attraction allows these larger silicate layers to gather into thicker packets which increase in thickness as the growth proceeds. Crystallites of smallest areal extent and weakest interlayer interaction behave as single-layer particles, separated by water envelopes. Whenever the areal extent of the crystallite increases to the size that it can permanently bond itself to a neighbor, packets created increase in size (Fig. 3-27). Such a larger unit would be recognized as an illite-montmorillonite-illite (IMI) diffraction sequence. The next stage would be the creation of IMII units, and so forth.

INite diagenesis stages. In nature, the process outlined above is not a continuous one, but normally occurs in stages as related to a continuous

175

Eq M

Ei M'

a

b

d

C

Fig. 3-27. Stages in crystal growth of 2 : 1 clay minerals showing the currently recognized states. (After Jonas, 1975, p. 11, fig. 5.)

variation in the geological conditions. Foscolos and Kodama ( 1974) proposed the following scheme of stage subdivision in these various stages: (1) Discrete expandable clays are confined to the early diagenetic stage and they, therefore, commonly do not occur below 4,500 f t of burial depth. (2) Division between early and middle diagenesis cannot be drawn sharply, but is defined as the point at which: (a) discrete expandable layer-silicates and kaolinite are present; (b) 1-M 1-M, illite polymorphs comprise 75% of the illite or more; (c) illite contribution in the interlayered clays is 25% or lower; (d) crystallinity index is 2 20 mm and sharpness ratio is d 1.5. (3) Transition from middle diagenesis to late diagenesis is defined as the point when: (a) kaolinite just disappears from the clay-mineral assemblage; (b) the illite content in the interstratified clay is 80%; (c) 50% of the illite is the 2-M polymorph; (d) crystallinity index is d 15 mm and sharpness ratio is 2 2.0. (4) In general, kaolinite disappears below a depth of 9000 ft, whereas illite-smectite interlayers generally occur between depths of 3000 and 12,000 ft. The maximum depth at which they are found is around 15,000 ft.

+

Changes in clay mineral composition-kaolinite and chlorite diagenesis The statistical distribution of clay minerals during geological periods compiled by Weaver (1967b) indicates that contents of both kaolinite and swelling 2: 1 clay minerals are diminishing with progressive age of the sediments (Fig. 3-28). Hower et al. (1976) showed a distinct decrease in kaolinite content with sediment depth in the fine-clay fraction from shales of

176 swelling

clay-

20

kaolinite

0c h l o r i t e

[i l l i t e LO

60

A0

Fig. 3-28. Statistical distribution of clay minerals during various geological periods. (After Weaver, 1967, cited by Dunoyer de Segonzac, 1970, p. 294, fig. 4.)

the Gulf Coast, U.S.A. (Fig. 3-29). Parham (1966), who noticed the scarcity of kaolinite in ancient sediments of the U.S.A., tended to believe that either the nearshore regions of basins, which constitute the areas richest in kaolinite, could have been preferentially eroded, or that the climate of ancient areas was not favorable for the formation of kaolinite. From the study of a series of sedimentary sequences, Dunoyer de Segonzac ( 1970) suggested that kaolinite is progressively eliminated during burial diagenesis and possibly transformed into illite. He stated that the instability of kaolinite during burial is determined by geochemical factors, which are not strictly related to regular gradients of temperature or pressure. On the other hand, Curtis and Spears (1971) from considerations of solubility data for gibbsite and kaolinite, suggested that silicification of hydrated aluminium oxides, deposited in sedimentary basins, and their diagenetic transformation may represent an important mechanism of kaolinite formation in ancient sediments. Together with illite, chlorite content appears to increase with increasing

177 sediment age (Fig. 3-28). Very old sediments are frequently composed solely of illite and chlorite. In their study of Oligocene-Miocene shales from the Gulf Coast, U.S.A., Hower et al. (1976) observed an increase in chlorite content with burial depth and attributed this increase to deep-burial diagenesis. According to Dunoyer de Segonzac (1970), the diagenetic formation of chlorite is preceded by the formation of corrensite, mixed-layer chloritevermiculite or “chlorite labile to heating”, i.e., a chlorite phase that contracts partly on heating. The lability to heating decreases with burial depth and disappears completely during the deep-burial diagenetic stage. X-ray analysis indicates the presence of a chlorite-like component in Tertiary shales from the Canadian Northwest Territories (Powell et al., 1978). It is suggested that the mineral diagenetically formed here during burial by adsorbing the amorphous Fe,O, and/or A1,0, in the expandable layer silicates, according to the following scheme: illite or glauconite

~

smectite

+AP+, K +

- S i 4 + , H,O

vermiculite

+ Fe(OH), chlorite

The association of corrensite, expandable chlorite, and illite-smectite in the Lower and Middle Ordovician strata of Cow Head Klippe, western Newfoundland, suggests a qualitative sequence of diagenetic events (Suchecki et al., 1977). During burial diagenesis, Mg2+-rich smectite, which may have formed during early diagenesis, and/or volcanic detritus are altered to dioctahedral illite-smectite of low expansibility and a trioctahedral mixedlayer clay (corrensite and/or expandable chlorite). Corrensite can continue to exist with a higher proportion of expandable layers than the illite-smectite because of its trioctahedral character. According to these authors, the larger hydration energy of Mg2+ contributes to the higher thermal stability of trioctahedral expandable phases, such as corrensite, as compared to dioctahedral illite-smectite with a comparable proportion of expandable layers. Relatively uniform illite-chlorite parageneses are almost always formed from the different primary clay-mineral associations as they approach the boundary between diagenesis and metamorphism (Eckhardt, 1958; Frey, 1970; Millot, 1970). Increase in the chlorite content takes place more slowly than that of illite. Chlorite forms in larger amounts only under lowtemperature metamorphism; at the same time illite is transformed into

178 I

i

I

I

Q h

t L n

a,

c

?f3

-

0 5a 2

- = O . l -

0.5!.irn

4

.

I

I

I

10

20

30

Percent koolinite

Fig. 3-29. Kaolinite content of the
sericite (Eckhardt, 1958). Montmorillonite, kaolinite and, occasionally, illite act as source material for the formation of chlorite during the deep-burial stage. The sprouting of chlorite crystals observed in argillaceous sediments by Eckhardt supports the latter mode of genesis. In the case of montmorillonite, however, a transformation into a tetramorphic layer-silicate can take place fairly easily through fixation of a brucite layer in the interlayers (cf., the formation of chlorite-sudoite in the pre-burial stage), but the mineral formed in this manner would be dioctahedral (sudoite). In fact, such dioctahedral tetramorphic layer-silicates (sudoites) have been described as diagenetic neoformations in argillaceous sediments. Weaver ( 1959) found that sudoite (“dioctahedral chlorite”) is a common mineral in the Middle Ordovician K-bentonite beds from Pennsylvania, Tennessee, Virginia, Alabama and Kentucky. Volcanic ash presumably altered to a dioctahedral montmorillonite-like mineral; the presence of Mg in interstitial water enabled precipitation of brucite between the layers. Von Engelhardt et al. (1962) found that sudoites are important clay minerals in weakly consolidated clays and marls of the Middle Keuper of southwest Germany. The sudoite content increases with decreasing amount of kaolinite present. This led Kromer (1963) to believe that sudoite forms from kaolinite, because primary differences in the kaolinite distribution are not probable, even though they cannot be excluded.

179 The transformation of biotite The diagenetic behavior of clastic biotite (or ledikite) in platform and geosynclinal successions of the Soviet Union (marginal part of the Siberian platform; Verkhoyano-Kolimskaya syncline) has been studied by Kossovskaya et al. (1965). This behavior is completely different from the transformation of biotite into vermiculite during the process of weathering. Under diagenetic conditions, modifications of biotite to montmorillonite and kaolinite, can be traced going through a series of intermediate stages of regular and irregular interstratified phases, depending on environmental conditions and the rate of immersion of the sediments. In platform succession of clays (total thickness of 500 m), with increasing depth the following phases derived from biotite were detected: (1) trioctahedral mica-chlorite; (2) chlorite-“mobile” chlorite; and ( 3 ) “mobile” chlorite-montmorillonite. The succession of the changes may be pictured as follows: (1) potassium replaces Mg, which forms brucite layers (appearance of chlorite); (2) gradual hydration and decomposition of the brucite layer (appearance of “mobile” chlorite); ( 3 ) presence in the idterlamellar spaces of only water molecules and cations of the exchange complex (formation of montmorillonite). Besides the listed minerals, kaolinite and illite are also formed, The formation of kaolinite represents the final stage of the decomposition of biotite. Kaolinite clays which are usually encountered at the bottom of coal beds, formed under conditions of very slow sedimentation and lengthy diagenesis in acid medium. In geosynclinal succession of clays (total thickness of 3500 m) the following associations can be pointed out: (1) association with chlorite and mixed - layer montmorillonite - hydromica formations; (2) chlorite dioctahedral hydromicaceous association with relics of the transitional phase, and ( 3 ) association of chlorite and dioctahedral illite. Transformation sepiolite-talc. Fuchtbauer and Goldschmidt ( 1963) pointed out that sepiolite and palygorskite frequently occur in younger evaporites, whereas talc is more frequent in older ones. Under certain conditions there could have been a transformation of sepiolite to talc. To what degree such a structural change is possible, however, should be investigated. Formation of authigenic feldspars According to Taylor ( 1964), authigenic orthoclase may b e extensively developed in shales containing pyroclastic material. In the Tertiary John Day Formation of Oregon, it occurs over an area of about 600 sq. miles. The results of potassium-argon age determinations from both authigenic and pyrogenic constituents suggest that orthoclase was formed at burial depths ranging from 400 to 2200 ft and at a temperature of 20-40°C.

180 Authigenic K-feldspars have also been described from several DSDP core sediments. In basal sediments of leg 35, Bellinghausen Abyssal Plain, Kfeldspars were observed to have formed in association with porcellanite layers (Kastner, 1976; Kastner and Gieskes, 1976). The authigenic formation of K-feldspar in supersaline lake sediments has been mentioned previously. Cementation of argillaceous sediments during diagenesis Cementation is not prerequisite for the consolidation of a clay-rich sediment. In fact, numerous clay sediments without any pore cement are found at considerable depths, the clay minerals and other clay-sized material being kept together by cohesion. As most argillaceous sediments initially contain allogenic or biogenic carbonates, feldspars, SiO, -minerals, etc., it is very probable that partial or a complete cementation can occur at an early or a late stage of diagenesis. Calcite is by far the most important pore cement, which enables cementation to proceed even at the beginning of the shallow-burial stage, as indicated by the carbonate concretions. The cementation of argillaceous sediments, however, generally takes place during a later period in the deep-burial stage. The second most important pore cement is quartz. In addition to the quartz and opal present initially in the sediment, additional silica is released during most of the diagenetic mineral alterations and can serve as a pore cement in situ. Dapples (1967) pointed out the possibility of an early replacement of clay minerals by silica in argillaceous sediments rich in SiO, and the consequent formation of siliceous shales. Sensitivity in Canadian clay sediments is explained by cementation of the microstructure involving the interaction of primary mineral particles with amorphous matter (Bentley and Smalley, 1978). The particles appear to be coated with amorphous matter consisting of an iron-silicon-hydroxyl complex. During sedimentation links develop as the particles come into contact. Reaction series The concept of reaction series developed by Eberl (1978b) on the basis of hydrothermal experiments, represents a way of relating the origins of several geologically diverse clay minerals to a common mechanism. Although developed on the basis of hydrothermal experiments, paragenesis in a reaction series, which is a function of pressure (P),temperature ( T ) and stability fields, can also be used to describe or even predict deep-burial diagenetic processes. The directions of the reaction series for dioctahedral smectites depend mainly on interlayer cation and on the concentration of its salt in the solution. All of the alkali and alkaline-earth interlayer cations give rise to the reaction direction towards mica, namely: smectite mica-smectite mica. +

+

181

This series includes also a rectorite phase: Na-smectite Na-rectorite paragonite. The Mg-smectite follows the chlorite direction if the solution contains M g 2 + . A similar approach was developed by Velde (1972, 1977). Experimental data on hydrothermally treated natural clay minerals were used to construct phase diagrams relating composition, pressure and temperature, which define the phase relations of mixed-layered solid solutions in dioctahedral illite-montmorillonite and trioctahedral expanding chlorite and corrensitelike minerals (Fig. 3-30). Three major relations were established: (1) The R3+-i.e., A13+ and/or Fe3+-content of the assemblage will control whether or not an expanding chlorite or corrensite phase will appear. These minerals are rich in R3+ as well as R2+ (Fe2+ and Mg*+) contents. (2) Temperature and pressure control the type and composition of the mixed-layered illite-montmorillonite mineral which is stable with aluminous phases. A sequence of different types of mixed-layered ordering can be established, which might be correlated with diagenetic or epimetamorphic grade. (3) The presence of “metamorphic” trioctahedral phyllosilicate phases, i.e., those due to the effects of earliest metamorphism, is correlated with the P-T-X variables. These phases include 7 A chlorite (iron-rich), 14 A chlorite, +

K-MICA

-+

R2’R3’

Fig. 3-30. Possible general phase diagram for illite and related phases as a function of composition and increasing pressure and temperature. Ill = illite, either predominantly as 1 Md or 2-M polymorph; I = illite mica; I , = I-MI1 type ordered mixed-layered mineral ( <20% expandable); M L = mixed-layered illite-montmorillonite; ML,<,=pressure of ordering reflection; Kuol= kaolinite; Chl =chlorite; EXP3 =expandable trioctahedral phase which includes expanding chlorite and corrensite. (After Velde, 1977, p. 269, fig. 8.)

182 expanding chlorite, and corrensite-like minerals. The general sequence of assemblages and the phase relations proposed should correspond to those which are encountered in nature with changing temperature and pressure. Parallelism between organic and inorganic matter diagenesis The burial diagenesis of silicate minerals is frequently accompanied by a parallel diagenesis of organic matter contained in the sediment. Organic matter undergoes cracking reactions to yield petroleum hydrocarbons. There is an increase in the proportions of extractable organic matter and an increase in the proportion of hydrocarbons in the extract (Tissot et al., 1971). With increasing degree of diagenesis, lower-molecular-weight hydrocarbons become prominent, whereas with excessive thermal alteration, petroleum hydrocarbons are cracked to methane and the composition of the residual organic matter approaches that of graphite. These transformations are reflected in the physical characteristics of the organic matter, such 'as spore coloration and reflectance, which can be correlated with the petroleum generation process. The rank of associated coals or coal fragments in the rock increases with increasing degree of diagenesis and can be determined by reflectivity measurements (Foscolos et al., 1976). Changes in vitrinite reflectance have been correlated with changes in clay mineralogy (Heling and Teichmiiller, 1974).Apparently, the transformation of smectite to mixed-layer clays occurs during the stage when the reflectance range reaches 0.4-0.5% R, (mean). The organic and mineral diagenesis from the Bassin de Douala in Cameroun was correlated by Durand et al. (1975). Zone A down to a depth of 1200 m, where organic matter is little developed, contains unchanged detrital clay minerals. In zone B (1200-2200m), which is the zone of hydrocarbon formation, smectite and kaolinite disappear and formation of interlayers starts. In zones C (2200-2800 m) and D (2800-4000 m), cracking develops, the contents of extractables and hydrocarbons diminish strongly, while the formation of illite and chlorite continues. They attribute these parallel processes to the effects of temperature on both organic and mineral diagenesis. An experimental diagenetic study of a modern lipid-rich sediment indicated that the only factors which affected variation in components measured at the end of each experiment, were: organic matrix (presence or absence of soluble lipids), temperature, and the organic matrix- temperature interaction (Harrison, 1978). Fig. 3-31 is a model representing a sequence of diagenetic zones with depth and temperature for black shale (Tourtelot, 1979). With increasing temperature and depth of burial, the rocks reach the zone of

183 DEPTH (KM) ,

,

01

I

0.01 :1

TEMP POROSITY (OC) ( Y o ) DIAGENETIC ZONES AND PRODUCTS

_-. -_ -

--

SULFATE REDUCTION

=0---80-1

[

PYRITE CARBONATES [ l 2 C ENRICHED) PHOSPHATES

FERMENTATION METHANECARBONATES

ENRICHED)

DECARBOXYLATION SIDERITE

HYDROCARBON FORMATION LIQUID A N D GASEOUS HYDROCARBONS PYROCATALYTIC METHANE MONTMORlLLONlTE -1LLITE

Fig. 3-3 I . Zones of diagenesis in black shale; a zone of oxidation may or may not be present at top of sediment column. (Modified from Curtis, 1977, by Tourtelot, 1979, p. 318, fig. 5.)

decarboxylation where the organic matter begins to decompose by chemical instead of biological processes. Jackson (1977) observed a correlation between crystal structure of illite and degree of humification of the associated organic matter in the prePhanerozoic and Phanerozoic sediments. He, therefore, suggested that molecular structure of extractable organic matter could be used as an index of incipient metamorphism. Jackson considered the possibility that the original structure and composition of the organic matter influenced postdepositional changes in the crystallographic properties of sedimentary clay minerals. Foscolos et al. (1976) proposed a classification of diagenetic stages on the basis of alterations of both inorganic and organic material. Diagenesis is divided into three stages: ( 1) Eodiagenesis (early diagenesis) that corresponds to the zone in which pore water is lost from the shales, little hydrocarbon generation occurs, and coals are of lignitic or sub-bituminous types. (2) Mesodiagenesis (middle diagenesis) corresponds to the main phase of oil genesis, when coals are of the high-volatile to low-volatile bituminous type. The first stage of clay dehydration occurs during the early mesodiagenesis and the second one in the late mesodiagenesis. (3) During telodiagenesis (late diagenesis) extensive cracking of the organic matter occurs and dry gas is the main hydrocarbon product. Coals fall in the semi-anthracite type range during this stage. In a later paper, Powell et al. (1978) modified the previous classification of Foscolos et al. ( 1976). Whereas the second clay-dehydration step was assigned formerly to the zone of late mesodiagenesis, it is now

Ref lec tonce percent R O mox

Wet Gas

C~C$~C,~IOO o-o

or-

I

Extroct 6 Hydrocarbons mg per g organic c

Woody-

M i x e d Loyer Clays dool spacing lF-6:

0-0

-Amorphous

Diogenetic Processes 6 Stoges

I--; Orgonic

'ine

Motter

Clays

methane for mot i o n

I

-.-

odvent of wetgas CiC,

4 I

I

generation from kerogenof: 1 resins (Lospholtenes

1st cloy dehyd ration

I

i s morph i c

2 hydrocarbons- high

only from omorphouc 3n-alkanes b pristane from woody-herboceous

I

l o s s of adsorbed

------__--

\-

subst itution O f Al"for Si"

H,O

K,O

maximum hydrocorbon(oil) generation from

increase in

kerogen decreose in kerogen H/C

decrease inCEC

t ----_---_u c ..-C L

odvent of colalytic crocking of l i q u i d hydrocarbons to gos ..

-

--

P

.-

ll

-

n 0

? U

.-C

I

Fig. 3-32. Relationship between clay and organic diagenesis in the western Sverdrup Basin, Northwest Territories of Canada. (After Powell et al., 1978, p. 1194, fig. 13.)

185 proposed to occur only in the very latest stage of diagenesis (telodiagenesis). That implies that the second clay dehydration occurs below the oil-generating zone. As a result, the role of clay dehydration in oil migration is opened to question. The relationship between clay- and organic-matter diagenesis as proposed by Powell et al. is presented in Fig. 3-32. A vitrinite reflectance level of 0.5% R, maximum is a good indicator of the onset of hydrocarbon generation, both from the standpoint of wet gases (C,-C,) and liquid hydrocarbons; it corresponds to the onset of the first clay-dehydration process during early mesodiagenesis.

CHANGES IN STRUCTURE OF CLAYS DURING DIAGENESIS

The loosely packed structure of freshly deposited clays becomes more dense under the weight of new layers of sediment, with a simultaneous reduction of the water content, This process is called “compaction”. Compaction is defined as a “decrease in volume of sediments, as a result of compressive stress, usually resulting from continuous deposition above them” (American Geological Institute, 1957). The compaction process involves increases in sediment density and decreases in its porosity and remaining water content. Significant changes in the microstructure of the sediment components occur as well. Moreover, the compaction process, also significantly changes the chemistry of the interstitial solutions and, thus, indirectly affects the diagenetic processes involving mineral-phase transformations and neoformations. Once a certain overburden load has been reached and the sediment compacted, the process is irreversible, that is, even after later uplift and erosion of the upper layers, with consequent release of pressure, the porosity attained at the maximum burial depth does not change. A detailed review of compaction during clay diagenesis has recently been given by Von Engelhardt (1977). Changes in density

During the gravitational compaction of marine sediments, there is generally a rapid increase in bulk density within the first few hundred feet of burial. The curves presented in Fig. 3-33 show that with fine-grained clastics, bulk density increases rapidly on compaction. At depth, the bulk densities of sediments tend to approach the weighted averages of the grain densities because of decreasing pore volume. The rate of decrease appears to be related, according to %eke and Chilingarian ( 1974), to overburden pressure (or burial depth) and tectonic stresses, clay mineralogy, temperature, time, loading rate, and, in part, grain-size distribution, secondary cementing

186

I

DENSITY, g / c d

Fig. 3-33. Variation of shale bulk densities with depth in sedimentary basins. 1 =methanesaturated clastic sedimentary rock (probable minimum density); 2 = mudstone, Po Valley Basin, Italy; 3 =average Gulf Coast shale densities: values derived from geophysical data; 4 =average Gulf Coast shale densities derived from density logs and formation samples; S = Motatan-l, Maracaibo Basin, Venezuela; 6 =Gorgeteg No. 1, Hungary: calculated wet-density values; 7 =Pennsylvanian and Permian dry-shale density values, Oklahoma and Texas; Athy’s adjusted curve; 8 =Las Ollas-I, eastern Venezuela. (After Rieke and Chilingarian. 1974. p. 34, fig. 12.)

material, trapped salts in the pores and mineralogy of the non-clay fraction. From the measurements of bulk densities carried out on water-saturated Recent marine sediments by Preiss (1968), it can be seen that the rate of increase in density is highest in the uppermost 50 cm of sediment (Fig. 3-34).

187

DENSITY, g/cm3 Fig. 3-34. Relationship between bulk density and depth in deep-sea sediment. Solid circles = in-situ readings; open circles= repeated in-situ readings; solid squares = measurements made on cores by laboratory gamma-ray apparatus. N o allowance was made for instrument tower settlement or core shortening. (After Preiss, 1968, p. 640, fig. 3.)

The different shapes of the density-versus-depth curves obtained for argillaceous sediments indicate that no universal bulk-density curve can be constructed to characterize a specific type of sediment. Changes in porosity Porosity is an inverse function of pressure in homogeneous muds and shales. As shown by the curves in Fig. 3-35 compiled by Rieke and Chilingarian ( 1 974), the porosity of clays decreases rapidly with depth. Change in porosity is mainly a function of maximum overburden stress and of time, but is also affected by lithology, depositional environment, overpressured fluid zones, diagenesis, and tectonic stress. Apparently, even a very thin layer of younger sediments causes compaction. For example, according to Ziillig (1956) in the Zuger See even after the formation of a sediment layer 3.6 cm in thickness, the water content decreases from 83.6 to 70.6%. In Lake Constance, the decrease in water content is of the same order of magnitude (between 0 and 4.6 cm) below depositional interface (see Table 3-IV). The water content and the calculated porosity of clay-rich sediments of a 8.20m long core from Lake Zurich (Switzerland) were studied by Ziillig (1956) (see Fig. 3-36). The lowest layers of the core have an estimated age of

188

I

40

POROSITY, %

Fig. 3-35. Relationship between porosity and depth of burial for shales and argillaceous sediments. (see Rieke and Chilingarian, 1974, p. 42, fig. 17.)

about 5000- 10,000 years. The individual measurements show marked deviations probably due to the variation in grain size and composition of the core in the different layers; however, on the whole the decrease in porosity with increasing depth can be clearly seen. Between 0 and 8 m the porosity of the sediment decreases by about 30%, with the most marked decrease occurring in the uppermost few centimeters. With decreasing porosity, the compressive stress increases considerably, which clearly demonstrates the process of consolidation. The porosities of argillaceous sediments determined from sediment cores from three completely different environments (Black Sea, Santa Barbara Basin and Lake Zurich) are compared in Table 3-11. In spite of these differences in environment, mineral composition, and the completely different ages of the comparable layers (the lowest layers of Lake Zurich core

189 0 m

Y

Q 1

2

E. W

2 3 W c

E a

W

$

6

c W

1 B

5

u)

3

s m 1

I-

6

x 7

8

/r;

. Porosity PI.) (calc Sp 6p 7.0 Sp ' a 50 60 70

10 20 30

WATER CONTENT ( ahwet wt. )

100

300

500

700

900

COMPRESSIVE STRENGTH (dyne/cm*. 981 1

Fig. 3-36. Water content and compressive strength of sediments from Lake Zurich. (After Zullig, 1956, p. 75, fig. 15.)

are about 10-100 times younger than the lowest layers of the other two cores), there is a strong resemblance in the trend of porosity decrease with depth. This similarity apparently is limited to the shallow-burial stage of sediment compaction. At greater depth, in the deep-burial stage of sediment compaction, the porosity-depth relationship curves of sediments from different geologic provinces differ markedly (Fig. 3-37). I t is, therefore, probable that in the deep-burial stage of sediment compaction, other factors besides maximum overburden stress become prominent. For example, the work of Heling (1979) on the Upper Rhinegraben sediments shows that

190

-

0 30

>-

c

0 10

ti DEPTH, rn

Fig. 3-37. Porosity-depth relationships in argillaceous sediments from various geological provinces. (After Dickinson, 1951; Hedberg, 1936; Hosoi, 1963; and Athy, 1930; cited by Burst, 1976. p. 303, fig. 4.)

porosity-depth relationships vary with differences in the geothermal gradient (Fig. 3-38). There are detailed studies on the compaction (or density) of argillaceous sediments which are not influenced by tectonics, at depths up to 3000m for the Tertiary deposits of Venezuela, the Tertiary sediments of the Po Basin, and the Liassic deposits of Germany (Hedberg, 1936; Storer, 1959; and Von Engelhardt, 1960; cited by Muller, 1967). As shown in Fig. 3-39, a relation exists between the void ratio and the logarithm of the depth, which can be expressed by the following equations: E z E , -h.logt e E=100 - e

(3-3) (3-4)

where E = void ratio at a depth t (in m), E , = void ratio at a depth t = 1 m, b = index of compressibility of a certain clay, and e = porosity (in W ) . If the above equations are valid until the volume of pore space reaches 0, the depth t,, can be calculated at which porosity no longer exists ( e = 0): logt,, = E , / b . The constants given in Table 3-VII apply to the compression curves presented in Fig. 3-39. Von Engelhardt (1960) pointed out, however, that this simple relationship does not apply to all depths. Deviations are to be expected both for small and for great depths. Based on the previous discussion, it is probable that at greater depth the pore space increases more slowly than the empirical equation indicates. The calculated depth t o , therefore, has no real meaning. The decrease of initial volume of argillaceous sediments with increasing

191

MEDIAN PORE RADIUS, m . in A

Fig. 3-38. Porosity data (void ratio E), median pore radii ( r , , , ) and , abundances of rnontrnorillonite ( M ) and mixed-layer ( M L ) mineral in the < 2 pm fraction, as a function or dcpth. from four boreholes displaying different geothermal gradients from the Rhincgrabcn. Germany. (After Heling. 1979. p. 79. fig. 1 .)

depth of burial is considerable (Table 3-VIII). The rate of decrease is very high down to a depth of 500m. but with increasing depth of burial i t becomes less. According to Teodorovich and Chernov (1968), three stages can hc

192 Void

10

20

30

ratio

67

40

Fig. 3-39. Porosity and void ratio of argillaceous sediments related to depth of burial. (After H. Fuchtbauer, cited by Von Engelhardt, 1960, p. 39, fig. 21.) TABLE 3-ViI

'

Constants calculated from curves in Fig. 3-39 (after Von Engelhardt, 1960)

Tertiary, Venezuela Tertiary, Po Basin Liassic, northwestern Germany

1.844 1.700 1.160

0.527 0.48 1 0.3 17

65 63 54

3,160 3,500 4,570

'

E , =void ratio at a depth t = 1 m; b=compressibility of clay; e=porosity; r,, =calculated depth at which theoretically e=O. TABLE 3-VIII Decrease of volume of argillaceous sediments with increasing depth (after Von Engelhardt, 1960) Depth ( m )

500 I000 2000 3000

Decrease in volume (in % of original volume) Tertiary. Venezuela

Tertiary, Po Basin

Liassic, northwestern Germany

50.0 55.5 61.1 64.4

48.0 53.4 58.7 61.7

39.6 44.0 48.5 51.0

193 discerned in the compaction of clay horizons in the oil-producing deposits of Apsheron in the Azerbayjan S.S.R.: (1)During the first stage, occurring at burial depths down to 8-10m. there is a rapid compaction, resulting in a porosity decrease of clays from 66 to 40%. (2) During the second stage, in the depth interval from 8- 10 m to 12001400 m, there is a drastic decrease in the compaction rate; porosities decrease to 20-21%. (3)The third stage, in the burial interval of 1400-6000m, is characterized by very slowly compaction; the absolute porosity of shales decreases to 7-8%. More calcareous rocks have a lower initial porosity and the decrease in porosity of these rocks with depth is lower than in the case of less calcareous rocks. This effect is possibly due to the fact that the carbonate particles reduce the amount of plate-like and hydratable material.

'

Changes in residual moisture content

The loss of water from consolidating argillaceous sediments with depth is relatively rapid down to a depth of 250-300 m. On further burial, sediments become consolidated rocks and the rate of water escape decreases. Chilingar and Knight (1960) and Chilingarian and Rieke ( 1968) studied the relationship between moisture content and the applied pressure for various clays, soils, and marine' muds (Fig. 3-40). For montmorillonite clay hydrated in distilled water, a break in the curve occurs at 1000 psi. Possibly the pressure at which the rate of moisture loss decreases indicates when the expulsion of oriented water begins. Prior to this, possibly only the free (liquid) water is removed. Hedberg ( 1936) proposed a three-stage model for the dehydration of argillaceous sediments: Stage 1. Zero to 800 lb/in2 overburden pressure: (a) porosity of 90-75%; free water loss and first-order solid particle rearrangement: (b) porosity of 75-35%; loss of some absorbed water. Stage 2. 800-6000 lb/in2 overburden pressure: porosity of 35- 10%; mechanical deformation of sediments and continued loss of absorbed water; some incipient recrystallization. Stage 3. > 6000 Ib/in2 overburden pressure: porosity < 10%; high-pressure recrystallization.

'

See Rieke and Chilingarian (1974) for detailed discussion of this subject. Some carbonate muds have compactability similar to those of clays; however, the slope of the remaining moisture content versus pressure curves decreases with increasing carbonate content. (Editorial comment.)

194

110

.-m VI

I20

n a

2 TI

IW

s

5

80

W

c 2

8

KO

W

w 3 I-

2=

10

20

0

PRESSURE, p.s.i.

Fig. 3-40. Relationship between moisture content M (% of dry weight) and pressure p (in p.s.i.) for various clays, gum ghatti, and gum tragacanth. I =gum ghatti (natural organic colloid); 2 =gum tragacanth (natural organic colloid); 3 =silicic acid, M = 186-33 log p ; 4 = montmorillonite no. 25, Upton, Wyoming, John C. Lane Tract (bentonite), M = 104- 18.06 l o g p (straight-line portion of the curve); 5 =montmorillonite no. 25, M =58-10.2 l o g p (hydrated in sea water); 6 =illite no. 35, Fithian, Illinois, M =50-8.7 l o g p ; 7=kaolinite no. 4, Macon, Georgia, Oneal Pit, M=33.9-5.96 logp; and 8=dickite no. 15, San Juanito, Chihuahua, Mexico, M =26.7-5.04 logp. (After Chilingar and Knight, 1960, and %eke and Chilingarian, 1974, p. 59, fig. 30.)

Van Olphen (1963b) differentiated between two stages of compaction: one in which the particles are relatively far apart, and a second one in which they are separated by only a few monomolecular layers of water. In the first stage, compaction is primarily controlled by the double-layer repulsion (“osmotic swelling”). The pressures range from a fraction of 1 atm to tens of atmospheres. In the second stage, compaction is largely controlled by the forces of adsorption of the water layers on the clay surfaces. The removal of adsorbed water from between the surfaces of clay particles, as well as from quartz particles, will generally require extremely high compaction pressures, which are higher than those usually encountered in nature. The release of adsorbed water, however, will be facilitated by an increase of temperature. Pressure- temperature curves for dehydration reactions of montmorillonite minerals, saturated with different cations, show that water can be removed from clay surfaces at fairly low temperatures (Crowley and Roy, 1959). Meade (1964) concluded that “temperature might be as important a factor as pressure in removing the last increments of water from clay”. The results of

195 these and other investigations raise the question as to the nature of the water being removed and that of the driving forces. Measurements by Von Engelhardt and Tunn (1954) show that the fluid volumes adsorbed by clays exceed the pore volumes of dry powders. In addition, the volume of fluid adsorption by montmorillonite considerably exceeds the volume of its intracrystalline swelling. I t is concluded, therefore, that the accumulation of densified water layers on the outer surface of mineral particles represents a significant portion of the total immobilized water component (Burst, 1976). The precise nature of this adsorbed water envelope is not clear. Suggestions vary from low-density ice-like structures with an hexagonal net configuration, to those of high-density in-filled hexagonal structures (Bradley and Serratosa, 1960). I t is very probable that all these forms exist within a threedimensional special arrangement that includes the entire interfacial region between the 0 - O H plane of the clay crystal and bulk liquid water. It is very probable too that besides the liquid (pore) water and structural water adsorbed on the solid phase surfaces, other water forms, such as clay-interlayer water, structurally bound hydroxyl groups, and compositionally bounded water molecules (such as those contained within the gypsum lattice), also are affected by the sediment dehydration process. The accurate nature of these processes is not known because of “the imprecise knowledge of clay interlayer water density” (Magara, 1975). The principal driving forces for the expulsion of liquid pore water by compaction and fluid flow are gravity, pressure gradients, and temperature. Additional processes are capillarity, osmosis, and ionic filtration. Magara (1974) suggested that the increase in shale pore-water volume owing to expansion coefficients reacting to geothermal effects, could be considered an additional driving force sufficient to induce fluid movement. Another temperature-dependent process responding to geothermal gradients is clay diagenesis, more specifically the conversion of smectite into illite, with the attendant water expulsion. The amount of water transferred from lattice to pores during clay diagenesis has been estimated to range between 10 and 15% of the compacted sediment volume (Burst, 1976). Powers (1967), noting the difference between the magnitude of initial compaction and compaction attributable to the smectite conversion to illite, designated the two effects as stage I and stage 11. Burst (1969) reduced stage I 1 (smectite to illite transition) to a relatively narrow, temperature-dependent zone and added a deeper, third stage in order to accomodate deep-burial water loss subsequent to the major diagenetic conversion (Fig. 3-41). Perry and Hower (1972) refined Burst’s interpretation, by splitting his stage I1 into two parts representing: (a) an initial high rate of water expulsion caused by the relatively rapid collapse of approximately 65% of the smectite layers at the

196 Density = 1.32 Recent buriol P o r e woter

n

Interloyer woter

Swelling clay solids

\ \

!

Non-swelling clay solids

! \ \ \

Non-cloy soItd5

\ \

\ ! I

I

\

dz1.96 ofter

\

\

d = 2.28

1st

Dehydrotion \

otter

. \

\,

2 nd Denydrotion

d = 2 57

Fig. 3-41, Marine shale bulk-composition during dehydration. The compaction model is based on a three-stage dehydration sequence and the transformation of smectite (swelling clay solids) into mixed-layer varieties (non-swelling clay solids). (After Burst, 1969, p. 81. fig. 6.)

top of the diagenetic zone, and (b) a second high rate of water expulsion resulting from the quick collapse of that fraction of the smectite lattices representing the interval between 65 and 80% of total collapsed lattices. The net effect of clay-mineral diagenesis on sediment dewatering is to restrict water movement in the mid-range depths (3000- 12,000 ft in the Tertiary sediments) and to encourage it at deeper burial depths where interlayer water has been liquified and acts as an effective porosity breaching agent (Burst, 1976). The influence of different factors during compaction in the shallow-burial stage of diagenesis, which are also largely responsible for the formation of the primary porosities, is shown in Fig. 3-42. With increasing overburden

'

197 Highly clays

A o

10

100

E f f e c t i v e overburden load ( k g / c m 2 )

1

10

Pressure

100

(k g / cm2)

(1

B

0

D

0

=

Kaolhitc

1 10 100 Pressure (kg/cm2)

1 10

100

P r essur e ( k g / c m 2 )

E o 10

100

Pressure (kg/crn2)

Fig. 3-42. Relationship of void ratio to other factors, observed in natural sediments and in laboratory experiments. Void ratio is ordinate in all graphs: note different void ratio scales. (After Meade, 1963, cited by Miiller, 1967, p. 164, fig. 8.) A. Generalized relation to effective overburden load and particle size in sediments. (Modified after Skempton, 1953, p. 55.) B. Experimentally determined relation to pressure and claymineral species. (Modified after Chilingar and Knight, 1960, p. 104, to show their results to 100 kg/cm2.) C . Experimentally determined relation to pressure and adsorbed cations in (0.2 p m fraction of montmorillonite. (Modified after Bolt, 1956, p. 91.) D. Experimentally determined relation to pressure and electrolyte concentration in unfractionated Fithian illite (about 60% by weight coarser than 2pm). (Modified after Mitchell, 1960, fig. M3.) E. Experimentally determined relation to pressure and electrolyte concentration in (0.2 p m fraction of Fithian illite. (Modified after Bolt. 1956. p. 92.)

pressure, the influences of the various factors become similar. Meade (1964) concluded that at overburden pressures greater than about 50 kg/cm2 the important influences on the water content of clayey sediments seem to be only particle size, type of clay minerals, and temperature. Compression experiments (30-3200 kg/cm2) were carried out by Von

198 Engelhardt and Gaida (1963) on pure smectite and kaolinite clay muds treated with NaCl and CaCl, solutions and distilled water. They showed that the equilibrium porosity with a distinct void ratio which is reached at a certain pressure, does not depend on electrolyte concentration. Meade (1964) concluded from this and other observations that “apparently the physicochemical influences of the different cation types and electrolyte concentrations do not affect the amount of water held by a clay unless the amount exceeds a certain minimum necessary to develop diffuse double layers around the particles. The minimum amount seems to be about 50% by weight in very fine-grained ((0.1 pm) Na-smectite; i t should be somewhat less in coarser-grained and more silty clays. When overburden loads have reduced the amount of water in a clayey sediment below this minimum, the forces that must be overcome in order to compact the sediment further are more conveniently thought of as forces of hydration-the attractions between the clay surfaces and water or between cations and water-rather than as forces of repulsion or attraction between particles.” The Na-adsorbed expandable clay minerals may restrict water flow through the mechanism of clay blocking (Von Engelhardt and Gaida, 1963). The newly liquefied water, which is derived from the clay-mineral diagenesis, has a low electrolyte content. Its entry into the surrounding saline, watersaturated formation matrix could cause blocking of the clay and restrict flow in rock pores (Burst, 1976). Chunges in microstructure and buildup of clay particles on compaction

Clay particles respond to the confining pressure of compaction in terms of their geometrical arrangement. Meade (1968) and Pusch ( 1 966, 1970) have studied the particle arrangement of clays in Recent argillaceous sediments and soils and have proposed various idealized models from the observed microstructures. These and other studies have been reviewed by Moon (1972) and Rieke and Chilingarian (1974). Lambe (1958a, b) proposed three main models of clay structure: the “salt-flocculated” structure, formed in an electrolyte-rich solution; the “non-salt-flocculated” structure, produced in electrolyte-free conditions: and the “dispersed” structure, produced by the action of dispersing agents. Upon flocculation, clays were supposed to form random, open structures, dominated by edge-to-face particle contacts. These structures, described also by Tan (1959) and later by Rosenquist

’

~

I

I t is important to note here that one of the editors (G.V.C.), who conducted extensive

research on drilling fluids, -observed that “flocculation” of clays in electrolyte solutions gives rise t o surface-to-surface orientation of clay particles due to the neutralization of negative clay charges by positive cations.

199 ( 1962) were thought to result from the attraction of positively charged plate

edges to negatively charged plate surfaces. In 1960, Aylmore and Quirk introduced the concept that lightly compacted clays were composed of “domains” or “turbostratic” groups (later also described as tactoids) of oriented packets of clay flakes in random array. This arrangement provides perfect orientation of the clay particle within the packets or domains, but only random orientation between the domains themselves (Fig. 3-43). Similar structures have also been proposed by Van Olphen (1963, 1964) and termed “oriented aggregates” or “books”. The orientation of the domains appears to be greatly influenced by electrolyte concentrations. Low concentrations are conducive to structures with a high degree of preferred orientation of domains, whereas high concentrations may decrease the degree of orientation. Experimental studies indicate that most of the reorientation of clay particles takes place during the very early stages of compaction. Very low pressures are sufficient to produce orientation of kaolinite or illite in clays. Pressures around or greater than 100 kg/cm2 produce Preferred orientation in the case of any platy clay mineral, including smectite (Meade, 1968). I t is, therefore, probable that the domain-structure is the predominant fabric of fresh or lightly compacted clays. Single-plate structures are possibly prevalent only in very dilute clay suspensions, having very low electrolyte concentration.

FLOCCULATED

DISPERSED

a

b

C

d

Fig. 3-43. A suggested scheme of particle arrangement in clay sediments. a. Open, random arrangement of domains of 2-3 particles per packet. b. Parallel or sub-parallel arrangement of domains of 2-3 particles per packet. c. Increased parallelism and more particles incorporated into each domain than in a, i.e., mudstone. d. Complete parallelism and more particles per packet than in b, i.e. shale. (After Moon. 1972. p. 318. fig. 7.)

200 The concept of the card-house structure has been utilized to explain the anomalous behavior of “sensitive” clays (“quick clays”). Rosenquist ( 1966) attributed the deformational sensitivity of these sediments to postdepositional changes in their microstructure. Sedimentation of the particles into a marine environment leads to their flocculation into an open “cardhouse” structure, which is stable. Subsequent decreases in electrolyte concentration of the pore water, during the postdepositional period, increase the effective zeta-potentials and thereby the mutual repulsion between the mineral particles. This model seems to produce a system of sufficient porosity in which the moisture content can be in excess of the liquid limit. Pusch (1966, 1970) and others have suggested that the unstable nature of quickclays may be due to the presence of organic substances which act as dispersing agents and prevent formation of oriented structures. Bentley and Smalley (1978), on the other hand, proposed a mechanism that involves cementation of the microstructure and is based on the interaction of primary minerals particles with amorphous matter. With compaction, there is an increase in particle orientation and domain dimensions. Pusch ( 1970) followed the microstructural changes in a soft marine quick-clay undergoing unconfined compression in the laboratory. The natural microstructural pattern of the clay was characterized by a network of small aggregates connected by bridges formed by particles. Links between the particles broke down successively on increasing shear deformation and formed domain-like groups of particles with preferred orientation. According to Heling (1970), only in the shallow-burial stage the geometric model of increasing parallel orientation with increasing compaction pressure agrees with the data obtained from the Tertiary shales of the Rhinegraben. In the deep-burial stages, the clay-mineral fabric is no longer controlled by mechanical rearrangement but instead by mineral transformations. Fissile shales represent, structure-wise, the opposite of quick clays. Fissility of shales appears to be correlated with clay flake orientation (O’Brian, 1970). Shales with high degrees of preferred orientation also have well-developed fissility. Random clay platelet structures prevail in non-fissile claystone. Low electrolyte concentrations in the sedimentation medium or high organic-matter contents, both preventing flocculation, are, therefore, supposed to promote fissility. It is, however, also quite possible that the processes responsible for the formation and preservation of the laminations in fissile rocks will also produce a well-oriented clay fabric. Thus, fissility and clay orientation are only indirectly related (Spears, 1976). The studies of microstructures of clay sediments have recently been greatly advanced by scanning electron microscopy (Keller, 1978).

’

’

In the opinion of one of the editors (G.V.C.), based on research, the reverse is true.

20 1 CORRELATION OF MECHANICAL A N D CHEMICAL - MINERALOGICAL CHANGES WITH DEPTH OF BURIAL, PRESSURE, TEMPERATURE, AND DURATION OF BURIAL

In Fig. 3-44 an attempt is made to correlate the main mechanical and chemical-mineralogical changes occurring during diagenesis with the depth of burial, increase of temperature and pressure, and the duration of burial. The figures given in the graph should only be considered as average values; for any particular argillaceous sediment, the true values may vary within greater limits.

THE TRANSITIONAL ZONE BETWEEN DIAGENESIS AND METAMORPHISM

'

At greater depth, argillaceous sediments are subjected to elevated temperatures. Some new minerals formed under these conditions are unstable at shallow depths. The critical transition temperature between diagenesis and metamorphism is 300°C and is almost independent of pressure (Winkler, 1965). Temperatures during low-grade metamorphism of pelitic rocks in the Glarus Alps are given by Frey (1970) as ranging between 250 and 300°C. In burial metamorphism at normal geothermal gradients (30"C/km), this temperature is reached at depths of ca. 10,000m. In regional metamorphism connected with large-scale orogenesis, additional thermal energy is produced so that critical temperature occurs at much shallower depths. For example, R. Trumpy (cited by Frey) suggested that the mass of the Glarus Alps overthrust, associated with Alpine orogenesis, may have been 5000-6000 m thick. This corresponds to the P-T conditions of the greenschist facies, the low-temperature subfacies of which are characterized by the quartz-albite-muscovitechlorite assemblage. If the original argillaceous material was poor in K + and rich in A13+, pyrophyllite becomes the characteristic mineral of the greenschist facies above 400°C (Winkler, 1965): 1 kaolinite 2 quartz = 1 pyrophyllite 1 H,O.

+

+

The same reaction is given for the formation of pyrophyllite in the anchimetamorphosed Upper Triassic and Lower Liassic clays and mark of the Alpine border region in Switzerland (Frey, 1970). Phengite and Al-rich chlorite formed from mixed-layer illite-montmorillonite. For the formation of paragonite, the following sequence is proposed: irregular mixed-layer

'

See Chapter 5 .

202

I

I I

I

0

0

I

I

I

I

I

I

x

D

g

t

8

I

8

w

I

Fig. 3-44. Relationship of mechanical and mineralogical-chemical changes occurring in argillaceous sediments during diagenesis with depth of burial, pressure, temperature, and duration of burial. (After Miiller, 1967, p. 171, fig. 1 I.)

203 illite-montmorillonite + regular mixed layer mica-montmorillonite + mixed -layer paragonite-phengite paragonite. The so-called “zeolite facies” has been described from New Zealand, Australia, the Soviet Union, and North America. It is the typical mineral facies of the burial metamorphism of sediments containing zeolites (for literature, see Winkler, 1965, and the present volume, Chapter 5). The critical zeolite of this facies is laumontite (CaAl,Si,O,, 4 H,O), which never occurs in unmetamorphosed sediments. Winkler, therefore, more accurately calls the zeolite facies the laumontiteprehnite-quartz facies. The sedimentary assemblages of both analcime plus quartz and heulandite (or clinoptilolite) disappear under the same pressure and temperature conditions and are replaced by albite and laumontite: --f

analcime

+ quartz = albite + H,O

heulandite = laumontite

+ 3 quartz + 2 H,O.

At only slightly higher temperatures these reactions are followed by a reaction between laumontite and calcite, with formation of prehnite, which is also characteristic of the zeolite facies: laumontite

+ calcite = prehnite + quartz + 3 H,O + CO,.

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212 Velde, B. and Odin, G.S., 1975. Further information related to the origin of glauconite. Clays Clay Miner., 23: 376-381. Von Engelhardt, W., 1960. Der Porenraum der Sedimente. Springer, Berlin, 207 pp. Von Engelhardt, W., 1977. The Origin of Sediments and Sedimentary Rocks. Part I11 of: W. von Engelhardt, H. Fuchtbauer and G. Muller (Editors), Sedimentary Petrology. A Halstead Press Book, Wiley, New York, N.Y., 359 pp. Von Engelhardt, W. and Gaida, K.H., 1963. Concentration changes of pore solutions during the compaction of clay sediments. J . Sediment. Petrol., 33: 919-930. Von Engelhardt, W. and Tunn, W.L.M., 1954. The flow of fluids through sandstones. Ill. State Geol. Surv. Circ., 194: 17 pp. Von Engelhardt, W., Miiller, G. and Kromer, H., 1962. Dioktaedrischer Chlorit (“Sudoit”) in Sedimenten des Mittleren Keupers von Plochingen (Wurtt.). Natumissenschaften, 49: 205-206. Von Rad, U., Riech, V. and Rosch, H., 1977. Silica diagenesis in continental-margin sediments off northwest Africa. In: Y. Lancelot, E. Seibold and co-workers, Initial Reports ofthe Deep Sea Drilling Project, 41. U S . Gov. Print. Off., Washington, D.C.. pp. 879-905. Weaver, C.E., 1958. The effects and geologic significance of potassium “fixation” by expandable clay minerals derived from muscovite,‘ biotite, chlorite and volcanic material. Am. Mineral., 43: 839-86 1. Weaver, C.E., 1959. The clay petrology of sediments. Clays Clay Miner., Proc. 6th hail. Conf: Clays Clay Miner., pp. 154-187. Weaver, C.E., 1967a. Potassium, illite and the ocean. Geochim. Cosmochim. Acta, 31: 21812196. Weaver, C.E., 1967b. The significance of clay minerals in sediments. In: B. Nagy and U. Colombo (Editors), Fundamental Aspects of Petroleum Geochemistry. Elsevier, Amsterdam, pp. 37-75. Weaver, C.E. and Beck, K.C., 1971. Clay water diagenesis during burial: how mud becomes gneiss. Geol. SOC.Am. Spec. Pap., 134: 1-78. Weaver, C.E. and Beck, K.C., 1977. Miocene of the S.E. United States: a model for chemical sedimentation in a peri-marine environment. Sediment. Geol., 17: 1-234. Whitehouse, U.G. and McCarter, R.S., 1958. Diagenetic modification of clay-mineral types in artificial sea water. Clays Clay Miner.-Proc. Natl. ConJ: Clays Clay Miner., 5: 81-1 19. Wilson, M.J., Bain, D.C., Bettardy, W.J. and Berrow, M.L., 1972. Clay-mineral studies on some carboniferous sediments in Scotland. Sediment. Geol., 8: 137- 150. Windom, H.L., 1975. Eolian contribution to marine sediments. J. Sediment. Petrol.. 45: 520-529. Windom, H.L. and Chamberlain, C.F., 1978. Dust storm transport of sediments to the North Atlantic Ocean. J . Sediment. Petrol., 48: 385-388. Winkler, H.G.F., 1965. Die Genese der metamorphen Gesteine. Springer, Berlin, 218 pp. Wise, S.W. and Weaver, F.M., 1974. Chertification of oceanic sediments. Int. Assoc. Sedimentol. Spec. Publ., 1: 301-326. Yariv, S . and Cross, H., 1979. Geochemistry of Colloid Systems. Springer, Berlin, 450 pp. Zimmermann, H.B., 1977. Clay-mineral stratigraphy and distribution in the South Atlantic Ocean. In: P.R. Supko, K. Perch-Nielsen and co-workers, Initial Reports of the Deep Sea Drilling Project, 39. U S . Gov. Print. Off., Washington, D.C., pp. 395-405. Zullig, H., 1956. Sedimente als Ausdruck des Zustandes eines Gewassers. Schweiz. Z. Hvdrol., 18: 5-143.

213 Chapter 4 DIAGENESIS O F DEEP-SEA CARBONATES HARRY E. COOK and ROBERT M. EGBERT

INTRODUCTION

The enormous effort devoted to the study of carbonate sediments, especially during the past 25 years, has yielded a number of fundamental advances on the origin, facies distribution, and diagenesis of marine carbonates (e.g., see Ginsburg, 1957; Pray and Murray, 1965; Bathurst, 1971, 1975; Bricker, 1971; Milliman, 1974; Wilson, 1975; Chilingar et al., 1979). As pointed out by Matter et al. (1975) and Cook and Enos (1977), the geology of carbonate deposits has focused on shallow-water carbonates and, therefore, most of the ideas about their diagenesis have been based on these studies. Until the mid-1960s it was still widely believed that lithification of a carbonate sediment occurred mainly in the subaerial, fresh-water, and intertidal environments (Ginsburg, 1957; Bathurst, 1971, p. 323). Many workers continued to minimize the potential importance of submarine lithification (for example, Purdy, 1968; Dunham, 1969). This point of view finally had to be revised when, during the past 15 years, more and more occurrences of both shallow (Alexandersson, 1960; Ginsburg et al., 1967; and others) and deep-sea cementation (Friedman, 1964; Gervirtz and Friedman, 1966; Milliman, 1966; Fischer and Garrison, 1967; Bartlett and Greggs, 1969; Bricker, 1971; Milliman, 1974; Mullins et al., 1980; and others) of carbonate sediments were reported. Indeed, these studies presented persuasive evidence to demonstrate that deep-sea cementation can produce hardgrounds at the sediment-water interface at water depths of at least 3500 m. Therefore, although the role of fresh water in the lithification process credibly explained the data obtained on shallow-water carbonates (Ginsburg, 1957), it did not explain lithification of deep-water oozes on the sea floor. Before the start of the Deep Sea Drilling Project (DSDP) in 1968, deep-sea carbonates could be studied directly only from dredge hauls and piston cores, which did not penetrate much deeper than 10m below the sediment-water interface. With the advent of the Deep Sea Drilling Project, long cores (1000 m), sometimes continuous down to oceanic basement, have allowed investigators to study the two crucial diagenetic factors of depth of burial and age. Thus, for the first time progressive diagenesis could be studied. One of the most fundamental results of the DSDP is that in

+

214 deep-sea carbonate sections there is a general progression from ooze to chalk to limestone with increasing burial depth and age of the sediment. Interestingly, Hamilton (1959) suggested that the second seismic layer in deep-sea basins could be lithified carbonate ooze, but his suggestion generally was not accepted. The most important sedimentological results of the DSDP up to 1973 are presented by Davies and Supko (1973). A number of excellent review articles, research papers, and symposia volumes are now available that include various aspects of deep-sea carbonates. A partial list includes Funnel and Riedel (1971), Lisitzin (1972), Schlanger et al. (1973), Hsii and Jenkyns (1974), Milliman (1974), Schlanger and Douglas (1974), Bathurst (1975), Honjo (1975), Matter et al. (1975), Sliter et al. (1975), Van der Lingen and Packham ( 1975), Berger ( 1976), Anderson and Malahoff ( 1977), Milliman and Muller (1977), Scholle (1974, 1977), Wise ( 1977) , Schlager and James ( 1978), and Mullins et al. (1980). A point that must be emphasized at the outset is the fundamental difference in chemical properties between unlithified shallow-water carbonate sediments and pelagic deep-sea carbonate oozes. Constituents of modern shallow-water carbonates are mainly a mixture of unstable aragonite and high-magnesium calcite, which are very susceptible to wholesale solution and cementation by fresh water. In contrast, the constituents in deep-sea carbonates (i.e., foraminifera and coccolithophorids) consist largely of stable low-magnesium calcite. Thus they undergo diagenetic changes that are different from changes occurring in shallow-marine carbonate sediment. Pelagic carbonate is not particularly susceptible to fresh-water alteration because of its initial composition and deep-marine environment. In this setting overburden pressures with increased burial depth promote gravitational compaction and dissolution-reprecipitation processes. These are the dominant processes that reduce porosity in deep-sea carbonates and promote lithification. The main focus of this paper is to emphasize significant advances relevant to diagenesis rather than the total literature. Thus, those studies that present new information and concepts are emphasized here. This review consists of three main parts: (1) those aspects on the origin and characteristics of modern deep-sea carbonates that bear on diagenesis; (2) changes in the physical properties and biotic constituents with increased burial depth and age, and the current thinking on the processes responsible for these changes; and (3) a model and summary of progressive diagenesis, the possible relationships between diagenesis and acoustic stratigraphy, potential of petroleum occurrence in deep-sea carbonates, and unsolved problems and new research directions. It is important to stress here that in preparing this review it became

215 painfully apparent that a titanic amount of literature has been published since 1965 that relates directly to this subject. Thus, certain omissions were unavoidable.

TERMINOLOGY

In this chapter, pelagic refers to “of the open sea”, as in the classical use of the word. Sediments deposited near shores or in narrow bodies of water are not pelagic (Riedel, 1963). Planktonic refers to marine organisms that are incapable of strong directed movement, their distribution being’ largely dependent upon water currents (Riedel, 1963). Lithijication is the total result of all diagenetic processes that convert an aggregate of loose grains into a coherent rock. These processes include, but are not limited to, gravitational compaction, dissolution, pressure-solution, grain interpenetration, and cementation (modified from Schlanger and Douglas, 1974, p. 118). Cementation applies to the single process in which calcium carbonate precipitated from an intergranular or intragranular solution, reduces or fills void space between and within grains and binds them into a rigid framework (modified from Schlanger and Douglas, 1974, p. 118). Ooze refers to deep-sea carbonates that “have little strength and are readily deformed under the finger or the broad blade of a spatula” (Winterer et al., 1973, p. 9-10). A stiff ooze can be cut with a wire cheese cutter. Chalk is a partly indurated ooze that is “readily deformed under the fingernail or the edge of a spatula blade” (Winterer et al., 1973, p. 10). Chalks are essentially friable limestones and require cutting with a bandsaw. Limestone connotes chalks that have been further indurated and can only be cut with a diamond saw. Dissolution and pressure-solution are used for two different processes both of which change a solid carbonate grain into its ionic components. Dissolution is herein used for the process whereby carbonate particles dissolve while in the water column, at the sediment-water interface, and at shallow burial depths before a rigid grain-supported framework is achieved by compaction. Dissolution thus mainly occurs in response to increased hydrostatic pressures and the resulting effect this has on the solubility of carbonates. Consequently, the grains are not yet undergoing any elastic strain and conceptually the solubility product constant, [Ca2+]. [CO:- 1, is the same on all parts of the grains’ surface. With increased compaction, the grains will be subjected to a gradually increasing elastic strain at their points of contact where the solubility of the grains will be greatest. Pressure-solution then is the

216 process whereby grains dissolve at their contact points (Bathurst, 1971; Milliman, 1974).

ORIGIN A N D CHARACTERISTICS OF DEEP-SEA CARBONATES

Distribution About 50% of the deep-sea floor is covered by calcareous ooze (Berger, 1976). Because the pelagic realm comprises half of the earth’s surface (Fig. 4-1), calcareous ooze is the most abundant crustal covering material ( - 25%) (Lisitzin, 1971, 1972; Parker, 1971 ; Ramsay, 1974; Berger, 1976, and others). Deep-sea carbonates are composed of virtually 100%biotic components with the principal groups being coccolithophorids and planktonic foraminifera (Riedel, 1963; Berger, 1976). Since the Jurassic, when planktonic foraminifera and coccolithophorids evolved (Black, 197l), deep-water carbonates have comprised about 67% of the world’s total marine carbonates (Hay et al., 1976). These carbonates accumulate at rates of about 1-30 m/m.y. (Van Andel et al., 1975), but some Cretaceous chalks may have sedimentation rates as high as 54 m/m.y. (Hancock, 1976). The distribution of modem deep-sea carbonates depends mainly upon carbonate productivity in the surface water mass, terrigenous sedimentation (in the sense that it can dilute carbonates), and the degree of carbonate saturation (Lisitzin, 1971; Milliman, 1974). High carbonate productivity correlates well with areas of high fertility in the euphotic zone (i.e., the zones of high productivity coincide with global divergences), because all the carbonate is precipitated by planktonic organisms (Berger, 1975, 1976). Thus, surface productivity (Fig. 4-2) to a major extent determines the basic patterns of deep-sea carbonate-ooze distribution on the sea floor (Fig. 4-l), which explains why the surface and general sea-floor distributions of coccoliths and planktonic foraminifera are the same (Barash, 1971 ; Berger, 1971, 1976; Lisitzin, 1971; McIntyre and McIntyre, 1971; Parker, 1971). The increased solubility of calcium carbonate with water depth is related to increased hydrostatic pressure, increasing CO, content within the ambient water, and decreasing temperature (Milliman, 1974). Thus, a depth level exists in the water column that separates well preserved from solution.-etched, poorly preserved foraminifera and coccoliths. This depth level has been termed the lysocline by Berger (1968) and generally lies at water depths between 3000-4000m (Berger, 1976). At greater water depths, the rate of supply of biogenic carbonates equals their rate of dissolution. Bramlette ( 1961) called this the calcium carbonate compensation depth, which is commonly referred to as the calcite compensation depth (CCD). The position of

Fig. 4-1. CaCO, distribution in bottom sediments of the oceans and seas (in % of dry sediment). I = < 1%; 2 = 1-30%; 3 = 30-70%; 4 = > 70%.(After Lisitzin, 197 I , fig. 1 1.5. p. 204; courtesy of Cambridge Univ. Press.)

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I

IZW

Fig. 4-2. CaCO, distribution in the surface suspended matter (0-5 m water layer) measured in percent of the dry suspension. 1 =el%;2=1-3%; 3=3-5%; 4=5-10%; 5=>10%. (After Lisitzin. 1971. fig. 11.1, p. 199;courtesy of Cambridge Univ. Press.)

the CCD varies, ranging between 3500 and 5500m water depth in the Atlantic and between 3000 and 5000m water depth in the Pacific (Berger, 1976). About half of the deep ocean bottom consists of calcareous ooze with the other deeper half, below the CCD, by material with only a few percent of calcium carbonate (Berger and Winterer, 1974). T.he boundary between the two lithofacies roughly follows depth contours of the CCD (Berger and Winterer, 1974). Coccolithophorids

Coccolithophorids which date back to the early Jurassic, are yellow-green planktonic algae (Milliman, 1974) with individuals consisting of round coccospheres that average about 5- 100 p m in diameter. Each coccosphere is made up of from 10 to 30 calcareous plates (Berger, 1976) called coccoliths

I

219

Fig. 4-3. Coccospheres and coccoliths. a , b =coccospheres, SEM scale=3 pm. c =single coccoliths, SEM scale= 2 pm. d , e =single coccoliths, SEM scale= 1 pm. (After Milliman. 1974, plate V, p. 57; courtesy of Springer-Verlag.)

(Figs. 4-3, 4-4), which in turn are composed of very small low-magnesium calcite rhombohedrons, 0.25-1.0 pm in size (Fig. 4-5) (Gaarder, 1971; Bathurst, 1971; Milliman, 1974). Coccospheres most frequently occur in the upper 200m of the water column (Uschakova, 1971), where up to 1.5 - lo7 coccospheres/liter of sea water can be present (Okada and Honjo, 1972).

220

C

d

e

f

9

h

Fig. 4-4. Coccolithophorids and related forms. a - c = discoasters; d =pentalith of Braurudosphaeru; e,f=rhabdoliths; g, h =coccoliths. (After Riedel, 1963, fig. 1, p. 869; courtesy of In terscience.)

Foraminifera

Tests of these protozoans range from about 50 to 1000 pm in size with planktonic forams being much more abundant (> 99%)in deep-sea carbonate oozes than benthonic forms (Figs. 4-6, 4-7) (Riedel, 1963; Milliman, 1974; Berger, 1976). Planktonic forams inhabit the upper part of the water column down to a depth of about 100-300m below the surface (Berger, 1976). As Berger (1971) pointed out, there must be a huge turnover of living forams because they are preyed upon to a marked degree.

Fig. 4-5. Coccolith constructed of overlapping calcite crystals. (After Bathurst, 1971, fig. 107, p. 68.)

22 1 Even though benthonic forams comprise less than 1% of the carbonate in deep-sea sediments they are significantly more resistant to dissolution than their planktonic counterparts (Parker and Berger, 197 1). Benthonic forams

Fig. 4-6. A living planktonic forarn (Globoquadrina hexagonu) from 3000 m water depth: scale= 100 prn. (After Adelseck and Berger, 1975. plate 1. p. 7 5 : courtesy of Cushman Found.)

may be calcitic (low-magnesium calcite), aragonitic, tectinous. or agglutinated with arenaceous grains (Milliman, 1974; Berger, 1976), whereas the tests of planktonic forams are made up of low-magnesium calcite (Berger, 1976, and others). Benthonic forams span the Phanerozoic; however, planktonic forams are not found until the Jurassic (Berger, 1976) and it was not until the Cretaceous that they became abundant (Riedel. 1963). Pteropods

Pteropods are planktonic gastropods up to several millimeters in length that are common throughout the world in the upper 200m of the water

I mm

Fig. 4-7. Some Recent planktonic forams. (After Riedel, 1963, fig. 4, p. 872; courtesy of Interscience.)

222 column (Fig. 4-8)(Riedel, 1963; Milliman, 1974). About equal numbers of pteropods and planktonic forams fall to the bottom in the pelagic realm (Berner, 1977). Pteropods are minor constituents in deep-sea carbonates,

Fig. 4-8. Pteropods from Recent sediment in the Tuamotu Archipelago. (After Riedel, 1963, fig. 5 , p. 873; courtesy of Interscience.)

however, due to the instability of their aragonitic shells (fiedel, 1963; and others).

DIAGENETIC ASPECTS

Introduction Diagenesis, strictly speaking, only refers to the changes that occur to sediments after deposition, but before metamorphism. It is important in diagenetic studies, however, to understand the modifications that can occur to biogenic particles before they reach the sediment-water interface. This is useful in order to gain a clearer perspective of what sediment features are inherited versus those changes that are of a postdepositional origin. Thus, the character and fertility of the upper waters determine what species of micro-organisms are potentially available to reach the sea floor. However, it is the biological, mechanical, and chemical processes in the water column, on the ocean floor and during burial, that control the degree and nature to which that supply is preserved (e.g., Parker and Berger, 1971; Ramsay et al., 1973). Commonly, the subject of carbonate diagenesis is discussed in terms of one diagenetic process after the other. This approach is reasonable, especially with ancient rocks, because in ancient sequences an end product of

TABLE 4-1

Diagenetic realms (After Schlanger and Douglas, 1974, table 2, p. 132; courtesy of Int. Assoc. Sedimentol.) Depth

Realm

Residence Time

0-200 m (surface water)

I Initial Droduction

Weeks

II Settling

Days to weeks for forams; months to years for coccoliths depending c n pelletization (Smayda. 1971)

Pelletized coccoliths, ratio of broken to whole nannoplankton increases downward, ratio of living to empty foram tests decreasing during settling (Lisitzin. 1971 ; Berger. 1971)

Inversely proportional to sedimentation rate and dissolution rate

'Honeycombed' structure (Tschebotariof, 1952). Large foram tests supported by chains of coccolith discs. This surface is actually part of Realm IV

50,000 years (at 20 m/10' years sedimenta-

Remoulded 'honeycomb', slight compaction, burrowing, destruction by ingestion and solution

-

~~

200 m to sea floor

3000-5000 m (KC Diagenetic

~

Porosity (I)) Velocity (V,) km/s

Highly dispersed calcite-sea water system; 10-10' forams/m3. 104-100nannoplankton/m' (Lisitzin, 1971: Berner. 1971)

I I I I

Ill Deposition

IV Bioturbation

tion rate)

I I

I I

I

1-200 rn

V

(sub-bottom)

Shallow-burial

2WloOO m t (sub-bottom)

VI

Deep-burial

VII Metamorphic

I'CN

4

1.45-1.50

$

Up t o N 1 2 0 10'years ~ (by then either subducted or uplifted)

Chalk with strong development of interstitial cement and overgrowtbs; transition down to limestone with dissolution of forams, pervasion by cement and overgrowths-grain interpenetration. welding and 'ameboid mosaics' (Fischer er a/.. 1967)

4

I

~ 7 5 - 8 0 ~

I

Recrystallization trending to 'pavement mosaic' (Fischer rr 01.. 1967) o f completely interlocking crvstals

?,

I

~

t X 75-60% V , c r 1.6-1.8

rU_

V,=

+-

,

,

V c C 1.45-1.6

Ooze affected by gravitational composition, establishment of firm grain contacts; dissolution of fossils and initiation of overgrowths

slope , a t 3000 m +Slope at 5000m I

(Nafe & Drake, 1963)

10 x 10' years (at 20 m/lO' years sedimcntation rate)

10'-107 years

I I

$I N 80% i -

~~

1-10 km sub-surface

Diagenetic potential 0 4 * WJ

~~

Potential) 0-1 m (sub-bottom)

Petrography

/ I

I I

6 0 x down to 3540% I .8 increasing to 3.3 km/s

I

I I I

I

N

40% down to

< s% V < J x 3 + u p to h krn/s

,' I

I N h)

w

224 diagenesis is being studied. Unfortunately, the “process”-by-“process” approach does not convey a clear understanding of what happens to a sediment during its journey to become a rock. Prior to the advent of the DSDP in 1968, this same approach was required for ancient deep-sea carbonates. Investigators looked at ancient deep-sea chalks and limestones and speculated on how they evolved. Did they compact? What was the source of CaCO, for their cement? What processes were involved in going from an ooze with 80-95% porosity to a highly altered limestone with less than 5% porosity? What were their original biogenic components? These and other questions can now be answered with much more confidence than 10 years ago by studying the progressive changes that occur in deep-sea carbonates with increased burial depth and age. Coupled with this, the scanning electron microscope, which became available in 1965, allows diagenetic processes like dissolution, pressure-solution and cementation to be studied in minute detail on a routine basis. The approach taken in this chapter i s to discuss diagenesis in the framework of diugenetic realms, i.e., changes that occur in the water column, at the sediment-water interface, and during shallow-burial, and deep-burial (Table 4-1) (Schlanger and Douglas, 1974). Within each realm, diagenesis is discussed in the context of the properties of the sediment, changes that occur in these properties, and the interpreted processes responsible for these changes. The purpose of this approach is to instill an awareness that the diagenesis of deep-sea carbonates is a progressive, dynamic event that begins the moment a biogenic constituent is “born” and continues through a variety of diagenetic realms until these constituents lose all identity. At this point they have changed from soupy mushes to hammer-ringing limestones. It is understandable that the most significant advances in the diagenesis in deep-sea carbonates have come through the study of DSDP cores, and the authors draw heavily on this data base. In deep-sea carbonates, four major processes modify physical properties, state of fossil preservation, texture, and geochemistry. These include: ( 1) gruvirutionul compuction (mainly within the first 200 m of burial), (2) dissolurion (at the sediment-water interface, at shallow burial depths, and some within the water column), (3) pressure-solution (mainly during the deep-burial phase), and (4) cementution (beginning within the first few centimeters of burial and continuing with increased burial depth and age). Changes in rhe water column A number of excellent papers have assessed the problem of whether biogenic carbonate particles undergo dissolution while settling through the undersaturated water column (Chave, 1965; Peterson, 1966; Berger, 1971;

225 McIntyre and McIntyre, 1971; Morse and Berner, 1972; Adelseck and Berger, 1975; Berger, 1975; Be et al., 1975; Hecht et a]., 1975: Roth and Berger, 1975; Takahashi, 1975; Honjo, 1975, 1977). Honjo (1977) has reviewed this problem and his paper is an excellent source of data on this subject . I t is considered that most biogenic carbonate constituents, including coccoliths, do not undergo significant amounts of dissolution in the water column while settling. Adult planktonic forams have high sinking rates and protective organic tissues, both of which preclude appreciable dissolution. Honjo (1977) estimated that planktonic forams greater than 125 pm in size may only undergo 0.7-1.8% dissolution while sinking 5000 m (Table 4-11). The fate of smaller forams is uncertain. Aragonitic pteropod shells should also sink rapidly and their rarity in deep-sea carbonates is probably due to dissolution at the sediment-water interface. An individual coccosphere of about 10 p m size settles at a rate of 0.3-13 m/day, a rate far too slow to survive dissolution through thousands of meters of water (Table 4-11). Aggregates of coccospheres and coccoliths in fecal pellets up to 1 mm in size, however. could settle thousands of meters per day (Berger, 1976; Honjo, 1975, 1977). Because large numbers of well-preserved coccoliths do occur in deep-sea areas, it is generally accepted that coccoliths descend rapidly through the water column within the fecal pellets of pelagic grazers and filter feeders (Honjo, 1977). Coccolithophorids eaten by grazers and filter feeders do not dissolve during ingestion and excretion. Furthermore, some pellets have a protective organic film of pellicle (Honjo, 1977) that provides a chemical barrier to dissolution. Dissolution of coccoliths can eventually occur in unsaturated deep water, TABLE 4-11 Estimate of the amount of dissolution that individual planktonic forarns and coccoliths undergo while settling through a 5000-rn water column. (After Honjo, 1977. table 3. p. 287; courtesy of Plenum Press.) Residence time days to s i n k 5000 m

Planktonic foraminifera 125-250 urn > 250 Um

Dissolution in this time interval (X)

1.8 0.7

29 6

Penetration distance

any vater column any water column any vater column

Coccoliths

E. huxleyi C. neoh&

c.

*

eeptopoha

36 103 32 103 6 x lo3

Distance survived when the dissolution rate of 0.03 rng between 500 m to 3,000 m) was applied.

3.9 3.2

104 104

850

ern-'

82 m* 90 m* 1,250 m*

y-’ (Peterson‘s rate

226 however, when the fecal pellets disintegrate and the individual coccoliths are released. Also, the most fragile coccoliths are probably eliminated in the upper parts of the water column through fragmentation by plankton feeders or by dissolution (Berger, 1973; Wise, 1977). Sediment-water interface (0-

- I m)

Because the two most abundant particles in deep-sea carbonates are planktonic forams and coccolithophorids, diagenesis of these sediments involves the changes these two components undergo. Likewise, lithification of these sediments from ooze to chalk to limestone is the collective response of these biotic grains to burial. Physical properties. Figures 4-9 through 4- 1 1 show the overall relationship between physical properties and burial depth for some DSDP stratigraphic sections. The physical properties that are discussed in t h s paper include porosity, permeability, bulk density, and compressional velocity. Porosity is probably the single most important property in any diagenetic study. Porosities at or near the sediment-water interface in a deep-sea calcareous ooze range from 70 to 95% (Tracey and co-workers, 1971; Cook and Cook, 1972; Schlanger et al., 1973; Matter et al., 1975; Van der Lingen and Packham, 1975). Carbonate ooze that contains mixtures of coccoliths and forams have both interparticle porosity and intrabiotic porosity (Cook and Cook, 1972). A number of workers note that initial porosity in foram oozes is higher than would be expected from their grain-size distribution (Johnson et al., 1977). This is because the volume occupied by an individual foram test is about 80% intrabiotic porosity (Schlanger and Douglas, 1974). Deep-sea sediments in some localized areas on modern and ancient sea floors have had porosities reduced from initial values of 70-95% to 10-20% by calcite cement, forming hardground. The calcite cement is derived locally from the dissolution of biogenic constituents and reprecipitated. On modern sea floors these hardgrounds are underlain by unlithified ooze (Wise and Kelts, 1972). Ancient hardgrounds have been studied extensively (Bromley, 1968; Bathurst, 1971; Kennedy and Garrison, 1975; Wilber, 1976; Mullins et Fig. 4-9. Porosity-depth relations in pelagic carbonates: Plots to the left of the depth scale are GRAPE density and porosity data for DSDP Site 167 (Winterer et al., 1973). The points plotted to the right of the scale are based on averaged porosity data from DSDP Sites 62, 63, 64. 71, 77. 78, 158, 161, 214, 216, 217. These sections are all characterized by high CaCO, (80% or higher) content. Inset is from Hamilton (1959) based on experimental data on the compaction of a Globigerina ooze containing 54%CaCO,. Note the rapid decrease of porosity with depth in the upper 100-200m. (After Schlanger and Douglas, 1974, fig. I , p. 119; courtesy of Int. Assoc. Sedimentol.)

221

S I T E 167 LITHOLOGY

'OROSITY $(XI

G R A P E $(XI DEPTH IN METRES

-

FORAM N~NNOFOSSII

50

-

100

-

150

-

200

-

250

-

300

-

OOZE

1

-

..-:-:-.. * . ........ ..... !I'.."" .. ........ ....... .. ... . . . . ....... .--. \ . ... . : .::*; ; ...,., ?.. . ... .. ,.., . . . ..... . .,:.. ..'..'? . :. *!',':.. . .. . . . . . ;.. . . .. ....:... *:._.. . .. ..:

hhhhr

R?..

FORAMNANNOFOSSIL CHALK

*

I

350 400 450

-

CHERT NANNOFOSSIL CHALK A N 0 LIMESTONE

--____

*

-...*. .

-

i

..

.

- .. .. . .,. - . . .. '.

..

500

-

550

-

'

.

.._ . ,

.

,.i :.

NANNOFOSSIL CHALK AND CHERT

MARLY LIMESTONE

______

CHERT AND LIMESTONE

I

POROSIT Y (%I

50

.... '. I

, ..

60

70

228

e

S E M OAlLLlNG SAMPLES OlSTUREANCE Metres

LATE PLIOCENE

1

MIOOLE MIOCENE

LATE OLIGOCENE 400

96

SILICEOUS SEISMIC FOSSILS L\ REFLECTORS

INDURATION

- 100

M -

--?Do

- >pc

-4490

-

1 EARLY OLIGOCENE

500

- 500

I 600

cream - 600

700

CRt3ACEOUS

900

EARLY CRETACEOUS

Fig. 4-10. Stratigraphic column and physical properties. DSDP Leg 30, Site 288, Ontong-Java Plateau. South Pacific. (After Van der Lingen and Packham, 1975, fig. I. p. 448, 449.)

229 RELATIONSHIPS BETWEEN DlACENESlS AND PHYSICAL PROPERTIES OF BlOGENlC SEDIMENTS

POROSITY (%)

BULK OENSlTY (p/cm’l

SONIC VELOCITY (km/secJ (horuontal component)

230 Sf M WRtS

DRILLING OISNAEANCE

)

PLRCENTAGE ULLCIUM

23 1 RtLATlONSHlPS BCTHELN DIACENESIS AND PHYSICAI PROPkRITES Ok BlOGl NIC SFDIVI NTS

232 al., 1980). Most oozes, however, remain unlithified at the sea-floor interface except for the occasional hardgrounds (Scholle, 1977). Wet bulk densities in surficial ooze range from about 1.1 (94% porosity) to 1.60 (70% porosity) (Cook and Cook, 1972; Schlanger and Douglas, 1974; Van der Lingen and Packham, 1975). Sound velocity is an important but complicated physical property related to porosity, bulk density, grain size, shape and orientation, properties of the interstitial water, intergrain attractive forces, temperature, and degree of cementation. Sound velocity will be ultimately controlled by the interactive effect of these properties on the elastic structure of the sediment (Hamilton, 1959; Mayer, 1979; Manghnani et al., 1980). Compressional velocity is

Mean Grain Size 48

1600

5.2

56

o a

'

22

7.6

80

8.4

8.0

9.2

*

ODo

a

0

a

E

5

,+

6.8

a

60°Aoo

2800 -

6.4

o

o

2400}

6.G

0

0 0

3200&a00

0

a

36004000

o

0

moo a

a

O D

CD

-

0

0

4400

dmo

-

4800-

0 0

a

A

'r

d

0

20

30

40

M

60

70

CaCOj Dissolved, Wl.%

Fig. 4-12. A. Relationship between the mean grain size of the sediment and water depth. B. Relationship between mean grain size and estimated amcunt of CaCO, dissolved. Circles= surface of box cores, triangles=middle of box cores, squares= bottom of box cores. OntongJava Plateau. western Equatorial Pacific. (After Johnson et al., 1977, fig. 4. p. 267.)

BULK SEDIMENT

SAND FRACTION

FINE FRACTION

Fig. 4-13. SEM of two sediment samples, showing the effects of dissolution upon mean grain size. A-C: 1613 m water depth, mean grain size = 4.99 9. Abundant whole, sand-sized forams in bulk and sand fraction. and well-preserved coccoliths in fine fraction. D-F: 4441 m dater depth, mean grain size= [email protected] less abundant and fragmented in bulk and sand fraction, and fine fraction primarily contains calcite fragments and poorly preserved or dissolution-resistant coccoliths. Ontong-Java Plateau, western Equatorial Pacific. (After Johnson et al., 1977, fig. 5, p. 268.)

h,

w w

234

O A 0 A'&'

24001 2800 -

5

A

00

0 OA

E

3200-

w

O

0

m

36000

a

0

0

4000 -

d

ma 0

m3

OOA

4400-q,

0

A

-

50

-

40

Ccnlinenlol Terrace Deposits

30 20 10 1600

Y

e

E

-

a

>

-

90 8070-

60 5040

Abyrsol Ploin Turbidiler

a

A

Mean Grain Size

1

,+

A O

U 0

A

0 0

l520'

C

,b

o;

o; 40 o; 6b o; CaCO3 Dissolved,Wt.%

e'o

40

do

Fig. 4-14. A. Relationship between sound velocity in the sediments and water depth. B. Relationship between sound velocity in the sediments and mean grain size. Solid line and dashed line are regression equations for terrigenous deposits on continental terraces and abyssal plains, respectively. C. Relationship between sound velocity and estimated CaCO, lost to dissolution. Symbols same as in Fig. 4-12. Ontong-Java Plateau, western Equatorial Pacific. (Modified from Johnson et al., 1977, fig. 6, p. 271.)

235 defined as:

(Y=

K+4’3pi”2 P

where K = bulk modulus (a measure of incompressibility or resistance to change in volume); p = modulus of rigidity (a measure of resistance to shear or change of shape); and P = saturated bulk density. It is assumed that the medium is isotropic and perfectly elastic. Thus, V, varies directly as the resistance of the sediment to compression and shear and inversely as its bulk density. A point worth emphasizing is that for compressional velocities to increase with burial depth, the rigidity of the sediment must increase more rapidly than the bulk density. Compressional velocities are on the order of 1.56 km/s for carbonate ooze (porosity 80%) on the sea floor (Schlanger and Douglas, 1974; Milholland et al., 1980). Hamilton et a]. (1956) showed that compressional velocity also correlates directly with grain size. Later, Johnson et al. (1977), studying the physical properties of deep-sea carbonates, reported that oozes in deeper water have lower shear strengths and sound velocities than oozes in shallower water. Carbonate oozes at the deeper sites have undergone more dissolution and resulting fragmentation of the foraminifera than oozes in shallow water. This results in the deepest-water oozes having the smallest mean grain size. Consequently, because compressional velocity varies inversely with grain size (Hamilton et al., 1956), these oozes have the lowest velocities (Figs. 4-12 through 4-14). An excellent paper by Mullins et al. (1979) discusses the echo-character of off-platform carbonates. Biotic constituents. Abundant data exist to show that dissolution of biogenic carbonate particles is a major process at and within the first meter of burial in deep-sea carbonates. These data are derived from theoretical considerations, experimental observations, macroscopic box-core information, and scanning electron-microscope studies. A number of factors control the degree and selectivity of dissolution on the sea floor. These include, but are not limited to, sedimentation rate (as it affects the sediments’ exposure time to waters undersaturated in CaCO,), water depth (in respect to the CCD and lysocline), bottom currents, shell texture and size (large tests with dense shell structures are more resistant) and protective organic films (Gardner, 1975; Berger, 1976; Takahashi and Broecker, 1977). As a generality, coccolithophorids and related forms (Riedel, 1963), such as discoasters, are significantly more resistant to wholesale dissolution than any of the carbonate secreting invertebrates (McIntyre and McIntyre, 1971 ; Berger, 1971; Honjo, 1975). Possible reasons include the incorporation of

236 cellulose-like material within the skeletal calcite, organic coatings on the surface of coccoliths, and relative position of the optic axis in the elements of the coccoliths (Bukry, 1971). Dissolution characteristics of coccoliths on the sea floor have been studied by a number of investigators (McIntyre and McIntyre, 1971; Schneidermann, 1973; Honjo, 1975; Roth and Berger, 1975) and these studies suggest that above 4000m water depth the majority of coccoliths are very resistant to solution (Fig. 4-15). Below 4000m water depth (near the base of the CCD) the sediment consists of a limited flora of resistant coccoliths showing solution effects (Fig. 4- 16) (McIntyre and McIntyre, 1971; Roth and Berger, 1975). In Roth and Berger’s (1975) study, an important observation was made: diagenetic low-magnesium calcite overgrowths on coccoliths occur at the sediment-water interface. This overgrowth cement was most prevalent in samples from about 3500-4800 m water depth. Thus, dissolution of coccoliths and reprecipitation of calcite cement on coccoliths can occur within a few centimeters of the sediment-water interface and does not require much burial nor a geologically long time. Both carbonate content and the number of forams per unit weight of the sediment decrease as the CCD is approached (Berger, 1976). Box cores from the Equatorial Pacific well below the regional CCD show foram assemblages that range from delicately spined tests with no apparent dissolution to severely etched individuals (Figs. 4- 17 and 4- 18) (Adelseck and Berger,

Fig. 4-15. Well-preserved assemblage of coccoliths, 2780 m water depth, central Equatorial Pacific. Scale= 1 pm. (Modified from Roth and Berger, 1975, plate 1(5), p. 89; courtesy of Cushman Found.) Fig. 4-16. Poorly-preserved assemblage of coccoliths consisting almost entirely of fragments and isolated placolith shields, 4600 m water depth, central Equatorial Pacific. Scale= 1 pm. (Modified from Roth and Berger, 1975, plate 2(5), p. 90; courtesy of Cushman Found.)

231

Fig. 4-17. Well-preserved planktonic foram with spines. Collected, in net tows. No obvious signs of solution etching are seen, 3000 m water depth, central Equatorial Pacific. Scale= 10 ym. (After Adelseck and Berger, 1975, plate 1(4), p. 75; courtesy of Cushman Found.)

Fig. 4-18. Planktonic forams from sediment-water interface box cores; all scales are 100 ym. 6 =well preserved, 4226 m water depth; (7) =well preserved, 3980 m water depth; 8 =solution etched, 4226 m water depth; 9 =partly dissolved, 4158 m water depth; lO=intensely dissolved, 4226 m water depth; I 1 =intensely dissolved except for the highly resistant keel, 4158 m water depth, central Equatorial Pacific. (Modified from Adelseck and Berger, 1975, plate 2(6- 1 I), p. 76; courtesy of Cushman Found.)

238 1975). Adelseck and Berger concluded that mixed assemblages showing varying degrees of preservation are to be expected if dissolution occurs mainly on the sea floor rather than in the water column. With continued sea-floor exposure, selective dissolution of planktonic species probably takes place (Orr, 1967, 1969; Berger, 1971; Savin and Douglas, 1973), especially if sedimentation rates are low. Experimental work by Hecht et al. (1975) shows that smaller forams dissolve more rapidly than larger ones, which is supported by data from sediment samples (Berger, 1971). Thus, in a general way selective dissolution among planktonic forams affects the relative percent of various species preserved in deep-sea carbonates (Berger, 1970; Hecht et al., 1975). The trend is for dissolution to enrich the resistant forms of planktonic foram and coccolith assemblages. The percentage of enrichment in resistant forms is one measure of the amount of dissolution that has taken place (Berger, 1976). Benthonic forams tend to be more resistant to dissolution than planktonic forams and the ratio of benthonic to planktonic forams has also been used as a measure of the degree that bottom sediments have dissolved (Parker, 1971; Parker and Berger, 1971; Ramsay et al., 1973; Berger, 1976). In a study of cores from the eastern Equatorial Atlantic, Gardner (1975) concluded that dissolution of microfossils was greater during Late Pleistocene glacial stages than during interglacial stages. He attributes this to “either an increase in the production and circulation of Antarctic Bottom Water during glacial stages or the production of a glacial North Atlantic Bottom Water. ..”. Approximately equal numbers of pteropods (aragonitic) and forams (lowmagnesium calcite) settle to the sea floor (Berner, 1977). Thus, about 50% by weight of the CaCO, settling to the sea floor is aragonite, a much larger percent than was previously believed (Berner, 1977). Most sea water below the thermocline, however, is undersaturated with respect to aragonite, and pteropods apparently dissolve rapidly before burial while on the sea floor (< 30 days) (Adelseck and Berger, 1975). Only in areas where the water is sufficiently shallow, can pteropod shells resist dissolution on the sea floor. Average maximum depths for significant pteropod accumulation is 500 m in the Central Pacific and 2800m in the Atlantic Ocean (Chen, 1971; Berner, 1977). Rich benthonic faunas exist in many parts of the deep-sea where carbonate oozes are accumulating. Box cores and bottom photographs demonstrate that bioturbation is intense and pervasive. For example, along the OntongJava Plateau in the western Equatorial Pacific, in water depths of about 1600-4500 m, the sea bottom is thoroughly reworked by organisms digging in, plowing through, or crawling across the sediment (Figs. 4- 19 through 4-21). In the stable deep-sea environment it is generally assumed that only those traces produced last and deepest in the sediment by burrowers will be

239

Fig. 4-19. Echinoid trails photographed on the sea floor in 2320 m water depth, Caribbean Sea. (After Ekdale and Berger, 1978, fig. 3, p. 268.)

preserved (Ekdale and Berger, 1978; Berger et al., 1979). Benthonic organisms can also promote dissolution of calcium carbonate in a variety of ways: (1) respiration by benthonic fauna produces CO,, which

240

Fig. 4-20. Spiroraphe on the surface of a box core from 3383 m water depth, Ontong-Java Plateau, western Equatorial Pacific. (After Ekdale and Berger, 1978, fig. 10, p. 274.)

increases the solubility of calcium carbonate; (2) digestive processes can remove organic coatings from foram and coccolith tests allowing unsaturated waters to reach their surface; and (3) bioturbation keeps the upper 5- 100 cm of the sediments mixed, preventing the forams and coccoliths from being rapidly buried and lost to near-surface dissolution (Paul, 1977). Apparently, microbial filamentous fungae can also affect deep-sea carbonates. These fungae can produce tubular microborings of 1 - 12 pm wide (Zeff and Perkins, 1979; Budd and Perkins, 1980). As Zeff and Perkins

Fig. 4-21. Grid pattern of bioturbation in a vertical face of a box core from 2247 m water depth, Ontong-Java Plateau, western Equatorial Pacific. (After Ekdale and Berger, 1978, fig. 6 , p. 271.)

24 1 pointed out: “infestation of individual skeletal fragments by microborers may be so extensive as to produce heavily bored envelopes resembling those previously reported to occur only under shallow-water conditions”. Shallow-burial phase ( I to

- 2.50 m)

For convenience of discussion, the shallow-burial phase closely parallels the shallow-burial diagenetic realm of Schlanger and Douglas ( 1974). As such, it includes those changes that occur within the upper 50-250 m of burial depth and the inferred processes that appear to prevail in this realm. During shallow-burial, early dewatering as a result of gravitational compaction is the major process for reducing porosity. Dissolution of biotic particles accompanied by reprecipitation as calcite cement occurs, but is probably not the dominant porosity reducing process in this realm. Physical properties. Results from the Deep Sea Drilling Project over the last 13 years have clearly demonstrated that in sections dominated by calcareous sediments, there is a gradual decrease in porosity and increase in bulk density with burial depth (Figs. 4-9-4-11). The magnitude of this porosity decrease is on the order of 15% (e.g., from 80 to 65%) with corresponding changes in bulk density from 1.35 to 1.64 (Schlanger and Douglas, 1974). Other DSDP sections show a similar decrease in porosity and increase in bulk density with depth (e.g., Gealy, 1971; Pimm et al., 1971; Tracey and co-workers, 1971; Cook and Cook, 1972; Davies and Supko, 1973; Schlanger et al., 1973; Matter et al., 1975; Van der Lingen and Packham, 1975). Even though porosity values generally decrease downhole, sharp reversals of increased porosity occur. This is especially true where relatively pure coccolith oozes are interbedded with foram-rich oozes. Foram-rich oozes have large amounts of intrabiotic porosity as well as interbiotic porosity, thus accounting for the sudden porosity increases in some sections (Cook and Cook, 1972). Porosity reversals, where porosity decreases with depth and then abruptly increases, can occur on a scale of only a few centimeters (Gealy, 1971). Also, some porosity reversals seem to reflect differences in compaction and/or cementation (Schlanger and Douglas, 1974). Compressional velocities in ooze are about 1.6 km/s- 1.8 km/s for oozes with porosities of 60-75% (Schlanger and Douglas, 1974). Because interparticle (interbiotic) porosity is a measure of the openness of the sediment texture, it is considered to be an indication of the rigidity ( p ) of the sediment structure (Hamilton et al., 1956; Hamilton, 1959). Compressional velocities and bulk density increase with increased burial depth in deep-sea carbonate sections (Figs. 4-9 through 4-1 l), but because compressional velocity is equal

242 to:

iK+i’3p)’/2 the rigidity ( p ) of the sediment must increase more rapidly than the bulk density ( P )with increased overburden. As discussed below, this rigidity increase in the shallow-burial realm can be accounted for at least in part by incipient cementation by calcium carbonate. Biotic constituents and textures. Texture deals with the size, shape, and arrangement of the component grains of a sediment. Inasmuch as most deep-sea carbonates are virtually 100%biogenic calcite, the textural changes that occur in the transformation of an ooze to a limestone involve the changes that foraminifera and coccolithophorids undergo during progressive burial. Individual coccolithophorids consist of spherical coccospheres, on which there are numerous calcareous plates called coccoliths (Fig. 4-3). Thus, the initial textural shape of these algae are spheroids about 5-10OIpm in diameter. What is generally seen in the sediment, however, is a mass of individual coccolith plates 2-20 pm in diameter because the coccospheres are fragile and the individual coccoliths are easily detached. The coccoliths themselves are composed of individual crystal elements ranging from 0.2 to about 1.0 p m in size (Bathurst, 1971). Foraminifera, in contrast, are much larger, about 50-1000 pm in size (Fig. 4-6). Thus, grain-size analyses of deep-sea carbonates are bimodal or polymodal: a coarse-grained fraction composed of foraminifera, a fraction in the 1-10 pm range representing whole coccolith placoliths, and a third, 0.2-1 pm, fraction composed of coccolith crystal particles (the “micarb” of Wise and Kelts, 1972).

Fig. 4-22. SEM of foram-rich nanno ooze with abundant broken coccolith shields. DSDP sample 32-305-5-3, 50 cm (39.0m burial depth). (After Matter et al., 1975, plate 1, fig. I . p. 908.)

243 Planktonic foraminifera vary from spheroidal to almost tabular or oblate in shape. Thus, a typical calcareous ooze will be a rather soupy mixture of randomly arranged forams and coccoliths with a very high pososity (Fig. 4-22) (Matter, 1974). Once the sediments are buried beyond the zone of active burrowing (5- 100 cm below the sediment-water interface), gravitational compaction moves the grains closer together by repacking, by reorientation of randomly oriented coccolith plates, and by mechanical breakage and dissolution-disintegration of delicate tests as overburden pressures are increased. Gradually, the ooze acquires a grain-supported fabric. A fundamental advance in the understanding of carbonate diagenesis is that the transition of a highly porous (e.g., 80-95%) carbonate ooze to a limestone with low porosity (e.g., 20-30%) does not necessarily require the introduction of massive volumes of calcium carbonate-rich pore waters from some great outside repository of CaCO, to supply the cement. A need for enormous pore-volumes of water (10,000- 100,000) from some unknown outside source to cement one pore (Pray, 1966; Dunham, 1969; Bathurst, 1971, p. 440) was predicated on what now appears to 6e two false premises, at least for lime muds of a deep-sea origin. The first assumption was that gravitational compaction is negligible in carbonate muds (Weller, 1959; Pray, 1960; Zankl, 1969; Bathurst, 1970, 1971), and the second assumption was that the stable low-magnesium calcite of forams and coccoliths is not a suitable donor of calcium carbonate in a dissolution-diffusion-reprecipitation process (Bathurst, 1971, p. 401). That compaction must be an important process in the diagenesis of pelagic carbonates was suspected by Tracey and co-workers (1971), Cook and Cook (1972) and others. Schlanger et al. (1973) and Schlanger and Douglas (1974) provided the necessary documentation that pelagic lime muds do compact and that they contain coccoliths composed of low-magnesium calcite, which provide calcium carbonate for cement. They first suggested that the progressive diagenesis of deep-sea carbonates in the Central Pacific was calcite-conservative. Matter ( 1974), Scholle (1977) and many others have since come to this same conclusion. Although this chapter deals with deep-water carbonates, the reader should be aware that early burial compaction can also be an important process in reducing porosity in shallower-water carbonate sediments (Meyer, 1980). It should be pointed out that deep-water submarine cementation of coarsegrained carbonates on the sea floor is probably enhanced by bottom currents circulating large volumes of sea water through the upper few centimeters of the sediment (Mullins et al., 1980). Evidence for extensive early compaction consists of the fact that close packing of originally highly dispersed particles is often accompanied by only a small amount of total cement and that originally circular burrows are flattened. In addition, some breakage of fossils and disintegration of fragile

244 planktonic forams and large coccoliths into micron-sized crystals (micarb) occur, not so much by mechanical means, but through dissolution along their sutures (Scholle, 1971, 1977; Schlanger and Douglas, 1974; Matter et al., 1975). An important point pertinent to compaction is that chemical disintegration of forams into small grains would effectively remove one of the types of evidence that one normally looks for to support gravitational compaction, i.e., crushed fossil tests. Also, carbonate geologists have long labored under the belief that if gravitational compaction occurs in lime muds or lime wackestones their contained microfossils should be crushed. The general lack of or minor occurrences of broken fossils has been taken as cogent evidence that these sediments did not compact. Should one necessarily expect microfossils to be crushed by gravitational compaction, especially at shallow-burial depths? Definitely not! Shinn et al. ( 1977) experimentally compacted modern shallow-water carbonate wackestones from initial porosities of about 7 0 4 0 % to porosities of 30-45%. Of particular importance is that delicate microfossils such as forams and calcispheres were not noticeably broken. Thus, as they pointed out, “ . .. absence of crushed microfossils in ancient limestones can no longer be considered evidence that limestones do not compact”. A possible explanation for microfossils not being crushed at least during early stages of compaction is that the fossils are being subjected mainly to hydrostatic pressures, whereby the pressures within and outside the tests are essentially the same. Shinn et al. (1977) suggested the same explanation for a lack of broken microfossils. Certainly at the sea floor, for example, at a water depth of 3000 m, where hydrostatic pressures are on the order of 4500 psi, microfossils would be pervasively crushed if the pressures were not equilibrated throughout the tests of the fossils. Also, it seems reasonable to assume that at shallow-burial depths (i.e., 100-200 m) the interstitial fluid pressures will be essentially hydrostatic and will be a combination of the height of the overlying water column plus the burial depth. This pressure will greatly exceed the lithostatic pressure. Schlanger and Douglas (1974) estimated that half of the total porosity reduction takes place in the upper 200m of the deep-sea carbonates in the Equatorial Pacific by gravitational compaction. In this interval, porosity is reduced from about 80 to 60% by compaction (Fig. 4-9). Other deep-sea carbonate sections show similar porosity reductions interpreted to be due to gravitational compaction (Matter et al., 1975; Van der Lingen and Packham, 1975) (Figs. 4- 10 and 4- 1 1). Variable degrees of dissolution occur in the shallow-burial interval. As noted by Matter et al. (1975) detailed studies of deep-sea carbonates exhibit differential preservation of different species as well as a range of preserva-

245

elements of the distal central area distal shield

S

proximal shield ( 2 cycles in Cenozoic)

Fig. 4-23. Cross-section of a coccolith placolith showing major structural features and sites of preferential deposition of overgrowth cement (0)or of dissolution (S). (After Matter et al., 1975, fig. 4, p. 897.)

tion states within each species. Even different parts of individual coccolith placoliths vary in their susceptibilities to diagenetic change (Fig. 4-23) (Matter et al., 1975; Wise, 1977). Matter et al. (1975) presented an excellent discussion of the progressive diagenetic changes that occur in a 640-m-thick carbpnate section in the northwest Pacific. Dissolution effects are seen within the upper 9 m of ooze, but overgrowth cement is negligible and diffuse (Figs. 4-24 through 4-27). With increased burial depth, small coccoliths tend to disaggregate along sutures due to dissolution. This yields large numbers of micron-sized crystals (micarb) that are very susceptible to further dissolution (Fig. 4-28). At a

Fig. 4-24. Proximal view of coccolith with most of the proximal shield and minor parts of the distal shield removed by dissolution. DSDP sample 32-305-1-2, 20 cm (1.70 m burial depth); X4000. (After Matter et al., 1975, plate 3, fig. 1, p. 913.) Fig. 4-25. Distal view of well-preserved proximal shield on coccolith. DSDP sample 32-305-1-2. 20 cm (1.70 m burial depth); X8250. (After Matter et al., 1975, plate3, fig. 2, p. 913.)

246

Fig. 4-26. Proximal view of dissociated distal shield of coccolith. Dissolution along sutures is destroying a large part of the elements. DSDP sample 32-305-4-3, 40 cm (29.9 m burial depth); X7050. (After Matter et al., 1975, plate 3, fig. 3, p. 913.)

Fig. 4-27. Distal view of coccolith showing minor overgrowth of cement on a few strongly overlapping elements of the distal shield. DSDP sample 32-305-4-3, 40 cm (29.9 m burial depth); X6700. (After Matter et al., 1975, plate 3, fig. 4, p. 913.)

247 burial depth of 90- 100 m almost all discoasters show overgrowths of highly euhedral calcite (Fig. 4-29). This same feature is reported by Schlanger and Douglas (1974), Adelseck et al. (1973), and others. The discoaster may eventually contain a volume of calcite several times greater than the original, Large coccoliths also show calcite overgrowths. Thus, as burial depth increases, coccoliths and discoasters show an important donor-receptor mode of diagenesis. Small coccoliths and fragments of coccoliths (micarb) undergo varying degrees of dissolution. Reprecipitation of this calcite takes place on discoasters, on large' coccoliths (Fig. 4-30), within the chambers of benthonic foraminifera, and as interparticle cement (Fig. 4-3 1). Experimental work by Adelseck et al. ( 1973), likewise, shows that discoasters and large coccoliths grow at the expense of small coccoliths under simulated diagenetic conditions (Figs. 4-32 and 4-33). Thus, it appears that discoasters receive the earliest cement, followed by individual groups of coccoliths, and finally the interstitial areas between placolith shields on coccoliths are cemented (Wise, 1973; Matter, 1974). Precipitation of calcite within foram chambers can also be important (e.g., Schlanger et al., 1973). Calcite cement within the upper 200-250 m of burial is the result of dissolution-diffusion-reprecipitation processes. The donors of this calcite are mainly planktonic forams (Fig. 4-34), delicate parts of coccoliths, and micron-size supersoluble grains of coccolith crystals (Fig. 4-35) (Schlanger et al., 1973; Schlanger and Douglas, 1974; Matter et al., 1975).

Fig. 4-28. Foram-nannofossil ooze, loosely packed well-preserved coccoliths. Abundant subhedral calcite crystals produced by the breakdown of coccolith elements and foram tcsts. DSDP sample 17-167-5-3, 132 cm (about 107 m burial depth). (After Schlanger et al.. 1973. fig. 4, p. 412.)

248

Fig. 4-29. Discoaster heavily overgrown with calcite cement, showing considerably thickened rays with discrete euhedral crystal facies. DSDP sample 32-305-1 1-4, 90 cm (96.9 m burial depth); X6400. (After Matter et al., 1975, plate4, fig. 3, p. 915.)

Fig. 4-30. Distal view of coccolith showing heavy calcite cement on elements of the shield and the central area. DSDP sample 32-305-11-4, 90 cm (96.9 m burial depth). (After Matter et al., 1975, plate4, fig. 8, p. 915.) Fig. 4-3 I. Chalk with well-preserved coccoliths. Euhedral pore-filling calcite crystals are shown by an arrow. DSDP sample 23-220-9-2, 60 cm ( I 52.1 m burial depth). (After Matter, 1974, plate 3, fig. 5, p. 449.)

249

Fig. 4-32. C. Distal view of discoaster from original sample. E. SEM distal view after discoasters in ( C ) were subjected to 300°C, 1 Kb for 32 days. Shows heavy overgrowth of calcite cement on discoaster rays. Scale= 1 pm. (Modified from Adelseck et al.. 1973, fig. 1 (C, E), p. 2756; courtesy of Geol. SOC.Am.)

The effect of diagenesis on planktonic formaninifera is a continual downhole decrease in abundance due to: (1) wholesale dissolution (Schlanger and Douglas, 1974), and (2) probably due to chemical disintegration into micron-sized particles no longer recognizable as foram fragments. The ratio of the number of planktonic to benthonic foraminifera is often used as a measure of a sediment’s preservation state. In modern, well-preserved pelagic carbonates on the sea floor, planktonic forams compose 97.5-99.9% of the total foram assemblage (Schott, 1935; Parker, 1954; Thiede, 1972). Thus, a

Fig. 4-33. E. Proximal view of coccolith from original sample. F. Proximal view of both coccolith shields after coccoliths in E were subjected to 300°C, I kb for 32 days. Shows central area of proximal shield extensively dissolved with coarse overgrowth cement o n proximal sides of both shields. Scale= 1 pm. (Modified from Adelseck et al., 1973, fig. 2 (E, F), p. 2757; courtesy of Geol. SOC.Am.)

250

Fig. 4-34. Foram-nannofossil chalk. Outer surface of foram test shows granular texture due to dissolution. DSDP sample 17-157-9-4 (about 225 m burial depth). (After Schlanger et al.. 1973. fig. 7. p. 414.) Fig. 4-35. Foram-nannofossil chalk. Same sample as in Fig. 4-34. Relatively well preserved coccoliths with euhedral grains of calcite formed by the breakdown of coccoliths and forams. and euhedral calcite crystals. DSDP sample 17-167-9-4 (about 225 m burial depth). (After Schlanger et al.. 1973. fig. 6. p. 413.)

small increase in the percent of benthonic forams may be the result of a large loss of planktonic forams due to dissolution. A minimum loss in planktonic forams due to dissolution can be estimated from the equation: L = 100 ( 1 - R J R ) , where L is the loss necessary to increase the insoluble residue R , (original) to R , and by assuming that benthonic foraminifera are the insoluble components. This approach has been used long enough to demonstrate that the estimated value of L is a useful measure for understanding preservation in pelagic foraminifera1 sediments. This assumes, of course, that benthonic forams do not dissolve with burial and that their productivity, averaged over long enough periods of time, remains essentially the same. Some dissolution must occur, but most benthonic forams appear to be very resistant. Schlanger and Douglas ( 1974) reported that L increases with increasing lithification. Loss values ( L ) range from 30 to 85% for oozes, 60 to 90% for Tertiary chalks, and over 90% for Cretaceous and marly limestones. Thus. even well-preserved oozes in the upper 60-75 m of the section may have had more than 50% of the planktonic forams dissolved to account for the increased percentages of benthonic forams. These figures are in

25 1 agreement with those obtained by other investigators (Berger, 197 1 : Berger and Von Rad, 1972). Deep-burial phase (

- 250 to 1000 + m)

This phase includes all changes and processes that occur from about 200 to 250 m burial depth to a depth on the order of 1000- 1200 m. In contrast to the shallow-burial phase where gravitational compaction is a major process in reducing porosity, gravitational compaction is generally considered to play a minor role at greater burial depths. Here the dominant process for porosity reduction is the precipitation of low-magnesium calcite overgrowths on large coccoliths, discoasters, within foram chambers, and on micarb grains. This cement is derived mainly from the dissolution and pressure-solution of small coccoliths, micarb grains, and planktonic forams (Schlanger et al., 1973; Matter, 1974; Schlanger and Douglas, 1964; Matter et al.. 1975: Van der Lingen and Packham, 1975; Scholle, 1977; Wise, 1977). Physical properties. According to Schlanger and Douglas ( 1974). between about 200 and 1200m, porosity decreases from about 60% down to 35-40% (Fig. 4-9). These changes are similar to those at other DSDP sites in areas of deep-sea carbonates. For example, on the Ontong-Java Plateau porosities at 200 m are about 60% and at 950 m they have been reduced to 20-25% (Figs. 4-10 and 4-11) (Van der Lingen and Packham, 1975). Wet-bulk densities at 200m are 1.60 (65% porosity) and increase to 1.95 at 1OOOm (44% porosity) (Schlanger and Douglas, 1974). Permeability is normally very low for all chalks regardless of porosity and distinct relationships exist between porosity and permeability (Fig. 4-36). Fractured chalks, however, have permeabilities 50-500 times higher than those in unfractured chalks (Scholle, 1977). Compressional velocities generally increase (a direct relationship) with increasing bulk density and decreasing porosity. Values at 200 m burial depth range from about 1.65-1.80 km/s in chalks (60-65% porosity) to 2.3-3.6 km/s in limestones at a depth of l000m (25-45% porosity) (Figs. 4-9 through 4-1 1, and 4-37) (Schlanger and Douglas, 1974; Van der Lingen and Packham, 1975; Milholland et al., 1980). Similar velocities have been reported for deep-sea carbonates (Hamilton, 1959; Pimm et al., 1971; Scholle, 1977). Velocity changes with burial depth must involve the interaction of a number of factors. For example, if a carbonate ooze compacts without cementation, the velocity-depth relationship would be smooth and fairly predictable (Fig. 4-38) (Schlanger and Douglas, 1974). Marked deviations from this curve, however, do occur. The fact that a plot for oozes, chalks, and limestones on the same velocity-depth diagram does not follow this

25 2

PERMEABILITY (md.)

Fig. 4-36. Relationship between porosity and permeability (air flow) for chalks and calcarenitic chalks. Dots =pure chalks; crosses =coarser, calcarenitic chalks. Lines are least-square fits to data. (After Scholle, 1977, fig. 4, p. 989; courtesy Am. Assoc. Pet. Geol.)

simple compaction curve, indicates that cementation is also reducing porosity. The effect of cementation is to increase the rigidity factor in the compressional velocity formula. Cement also increases the bulk density. As OOZE ( 0 m)

OOZE TO CHALK (200 m )

CHALK

( 6 0 0 rn)

LIMESTONE ( 1 0 0 0 rn)

glcc

1.35

1.60

1.82

1.95

C o m p r e s s i o n a l velocity, .knlsec

1.56

1.64

2.01

2.33

S h e a r velocity, k m l s e c

0.06

0.42

0.82

1.07

80

65

52

44

.465

.40 0

,366

Density.

Porosity,

x

Poisson‘s ratio. dimensionless

499

S h e a r modulus. k i l o b a r s

0.1

2.8

12.2

22.3

A

%

0.0

2.0

2.9

8.0

Z

0.0

0.3

0.5

6.2

0.0

1.9

2.7

7.9

P’

As, APO.

Fig. 4-37. Geoacoustic model for carbonate sediments. (After Milholland et al., 1980, fig. 10.)

253

Fig. 4-38. Relationship between compressional velocity and depth of burial. The solid line is based on data from Laughton (1954) derived through compaction of Glohigerrnu ooze. Circles= from seismic-refraction work; triangles = interval velocity data from DSDP Site 167. (After Schlanger and Douglas, 1974, fig. 7, p. 136: courtesy of Int. Assoc. Sedimentol.)

shown by Hamilton (1959), however, the rigidity of a lithified sediment increases faster than the density during gravitational compaction and cementation. Thus, the compressional velocity increases even though the bulk density is also increasing. It is important to note that the velocity of sound in rocks can increase with the age of the rocks independent of burial depth. Weatherby and Faust (1935) linked this velocity increase to the increase of lithification with age. Mullins et al. (1980) have clearly demonstrated that lithification can occur TABLE 4-111 Comparison of bulk density ( pB) and porosity ( @ ) with age and burial depth at three DSDP sites in the western Pacific. (After Gealy, 1971, p. 1104.) Site 62

Site 63

Depth Recent Middle Miocene Upper Oligocene

Site 64

Depth

Depth

(m)

PB

P

(m)

PB

P

(m)

FB

v

0

1.50

72

0

1.50

72

0

1.50

72

340

1.75

57

140

1.75

57

300

1.71

59

520

1.90

49

350

1.91

49

560

1.85

51

254 within a few thousand years. Gealy (1971) showed that porosity decrease is related to age as well as burial depth (Table 4-111). Schlanger and Douglas ( 1974) and Manghnani et al. (1980) attribute this to the kinetics of lithification and suggest that below about 250m and, therefore, with more elapsed time for dissolution-reprecipitation processes to operate, age is more important than burial depth in promoting lithification. The present writers agree that the role of age has been underplayed. The focus over the years has been on “depth of burial” with “residence time” being considered almost an incidental factor. Kinetics play a critical role and major diagenetic changes can take place if diagenetic processes are given enough time. Biotic constituents and textures. Below about 200-250 m of burial depth, dissolution of small coccoliths, micarb particles, and planktonic forams is severe. The effect of increased burial depth is to bring all biotic constituents into closer contact such that this increased burial stress promotes alteration of the particles by pressure-solution. It is important to note, however, that for pressure-solution to be most effective, the pore fluids must be undersaturated with respect to magnesium (Neugebauer, 1974). U p until this stage, dissolution takes place under normal hydrostatic conditions in which the pore fluids are undersaturated with respect to low-magnesium calcite. Pressure-solution occurs under nonhydrostatic stress, with dissolution occurring at grain and crystal contacts where the pressure is greater than the surrounding hydrostatic pore fluid pressure. The dissolved calcium carbonate migrates mainly by diffusion through the pore fluids (Berner, 1971; Bathurst, 1975) and precipitates locally on biotic grains at locations of lower pressure. The result of pressure-solution is that the adjacent biotic particles interpenetrate one another to produce sutured contacts (Fig. 4-39). Thus, at this stage porosity reduction is taking place mainly by pressure-solution and by reprecipitation of the dissolved calcium carbonate in void spaces. Matter et al. (1975) noted that at burial depths greater than 600m, proximal and distal shields on coccolith placoliths are commonly entirely welded, and the central areas on placoliths are filled with calcite cement (Fig. 4-40). By this stage only the most robust coccoliths and discoasters and some benthonic forams are preserved. Benthonic forams are generally more preservable than planktonic forams due to their thicker walls and fewer pores. Even these microfossils are subject to alteration, however, and Schlanger and Douglas (1974) noted the following changes: (1) a deterioration in surface luster and transparency; (2) a preferential removal of thinwalled, more porous rotaline species, especially miliolids; (3) a rapid downhole increase in percentage of broken tests; and (4) calcite overgrowth and chamber filling by low-magnesium calcite cement (Fig. 4-41). Solution seams (Scholle, 1977) and stylolites form with increased burial

255

Fig. 4-39. Pressure solution between a coccolith (top of photo) and a rhabdolith (arrow). DSDP sample 32-306-24-1, 140 cm (310.4m burial depth)X8000. (After Matter et al.. 1975. plate 6, fig. 9, p. 91 8.)

Fig. 4-40. Proximal view of coccolith showing calcite cement welding both shields togcther and infilling of cental area. DSDP sample 32-305-64- I , 100 cm (599 m burial depth) X ,3950. (After Matter et al., 1975, plate 7, fig. 5, p. 921.)

depths and advanced 'stages of pressure-solution and grain interpenetration. Stylolites, which are generally considered to be the result of pressure-solution (Bathurst, 1975), differ from grain interpenetration contacts only i n scale. Dunnington (1967) proposed 600-900 m as the minimum depth for the formation of stylolites and noted that in strongly stylolitic rocks volume reductions of 20-35% are common. Beall and Fischer (1969) suggested that pressure-solution effects begin at a burial depth of 250 m in DSDP cores in an area east of the Bahamas. Eventually, the once soupy mush becomes a dense limestone whose planktonic forams and small coccoliths may be virtually absent. Larger coccoliths and discoasters are severely etched, interpenetrate one another, and are overgrown with calcite cement. Most of the coccoliths are unrecognizable as to their true taxonomic affinity (Matter, 1974; Matter et al., 1975). Thus, at burial depths on the order of 1000- 1500 m, the texture of sediments may bear little resemblance to the original texture. The planktonic forams and smaller coccoliths may have long disappeared due to dissolution or are intensely recrystallized (Fig. 4-42), whereas the large coccoliths and discoasters have grown several times their original size due to cement "overcoats" and have a tightly packed mosaic appearance. The most advanced diagenetic changes, as observed in some of the

256

Fig. 4-41. Intrabiotic pore space of foram chambers partly filled with euhedral calcite crystals. DSDP sample 23-223-32-4, 118 cm (568.68 m burial depth). (After Matter, 1974, plate 9, fig. 6, p. 46 1.) Fig. 4-42. Nannofossil chalk. View of broken spherical hollow mass of radially arranged calcite crystals. This structure is interpreted to be the result of intense recrystallization of a planktonic forarn. DSDP sample 17-167-41-4 (about 680 m burial depth). (After Schlanger et al., 1973, fig. 12, p. 416.)

Fig. 4-43. Remnant of coccolith showing effects of pressure-solution. Coccolith immersed in a tightly welded fabric due to pressure-solution. Scale is 5 pm. Upper Jurassic (Tithonian) pelagic limestone, Austria. (After Fischer et al., 1967. fig. 54, p. 90; courtesy of Princeton Univ. Press.)

257

deep-sea limestones of Europe exposed on land, is to produce a “hammerringing” limestone, Grain interpenetration and welding of particles in these limestones is pervasive and the once soupy ooze has been transformed to a

Fig. 4-44. (a) Ameboid and (b) mosaic textures in limestones that have undergone some degree of metamorphism. (After Fischer et al., 1967, figs. 3(b), 4(b), pp. 19, 20; courtesy Princeton Univ. Press.)

“pavement mosaic”( Fischer et al., 1967) of completely interlocking crystals with less than 5% porosity (Figs. 4-43 and 4-44). Quantitative estimates of the relative amounts of the various biotic constituents that have dissolved are difficult to calculate (Matter et al., 1975). For Equatorial Pacific carbonates, Schlanger and Douglas ( 1974) concluded that the decrease in planktonic forams with increased burial depth was due to dissolution and that dissolution and collapse of weakened tests account for most of the porosity loss within the sedimentary column. They further concluded that the amount of calcite dissolved from the planktonic

258 tests almost equals the amount of calcite cement that occurs as: (1) overgrowths on coccoliths, (2) infillings in remaining benthonic foram tests, and (3) interparticle cement between adjacent foram grains. Their conclusion implies that most of the cement is derived from planktonic forams with lesser amounts being derived from small coccoliths and benthonic forams. Although it is very difficult to quantify the amount of dissolution that small coccoliths undergo, there is a general consensus that the diagenetic system is largely calcite conservative. Also, most investigators agree that the major changes in going from ooze to chalk to limestone involves the dissolution and eventually almost complete destruction of the frameworksupporting tests of foraminifera and the etching of small coccoliths and micarb particles. This decrease in the volume of biogenic calcite by dissolution and pressure-solution is balanced by a buildup of calcite as overgrowths on remaining fossils, particularly discoasters, large coccoliths, within the chambers of benthonic forams, and as interparticle cement.

GEOCHEMICAL CHANGES

Geochemical data is now abundant from both oceanic occurrences and onshore exposures of deep-sea carbonates that support the idea that the induration of carbonate ooze to limestone is a gradual process that takes place throughout the sediments’ burial history (Matter et al., 1975). Strontium

The biogenic calcite precipitated by coccoliths, discoasters, and foraminifera contains about 1400 ppm of Sr2+ which substitutes for Ca2+ in the calcite crystal lattice. A number of studies on DSDP cores have demonstrated a general decrease in Sr2+ content in the sediments and an enrichment of Sr2+ in. the interstitial waters with increasing burial depth (Fig. 4-45) (Matter et al., 1975; Sayles and Manheim, 1975; Manghnani et a]., 1980). The decrease in Sr2+ content in the sediment with burial depth, is accompanied by a progressive lithification by overgrowth of low-magnesium calcite cement on biotic particles. When unaltered biogenic calcite dissolves, Sr2+ is released into the pore water. However, Sr2+ is largely excluded from the precipitated cement. Thus, with increasing degree of diagenesis the Sr 2 + content of the pore waters increases and the Sr2+ content in the sediment decreases. As Kinsman (1969) predicted and as Matter et al. (1975), Manghnani et al. (1980), and others have pointed out, Sr2+ content of the sediment thus serves as an indicator of the degree to which the original

skeletal calcite has been dissolved and calcite cement added to the sediment. Schlanger and Douglas (1974) stressed that age and not just burial depth per se is a fundamental factor in the lithification of deep-sea carbonates. Manghnani et al. (1980) have assembled data which in a simple yet eloquent way support this idea (Figs. 4-46 through 4-48). At five DSDP sites, there is a poor correlation ( Y = 0.419) between Sr'+ content and burial depth (Fig. 4-46). The plot of Sr2+ content versus age (Fig. 4-47), however, shows a strong negative correlation ( r = - 0.875). This is consistent with Schlanger and Douglas' model for carbonate diagenesis in which they proposed that the degree of cementation is time-dependent and not simply dependent on burial depth. Thus, in areas of low sedimentation, even though the burial depth may not be great (Fig. 4-48), the biogenic calcite system will be moving toward a lower free-energy state through dissolution of metastable biogenic components and reprecipitation of large crystals as cement. Stable oxygen isotopes

Douglas and Savin (1975) presented an excellent discussion relating diagenetic effects to the isotopic composition of forams and coccolithophorids. Both of these microorganisms grow in the surface layers of the ocean and their initial oxygen-isotope composition yields estimates of ocean-surface temperatures. Benthonic forams secrete their tests on the bottom and, therefore, provide approximations of bottom-water temperatures from the oxygen-isotope composition. Two major assumptions in using the oxygen-isotopic compositions of these microfossils for oceanic paleotemperature analyses are that the isotopic composition of the biogenic carbonate remains: ( 1) stable, and (2) diagenetically unaltered after burial. In all of the deep-sea carbonate sections drilled in the DSDP. there is a downhoIe change from ooze to chalk to limestone. During diagenesis many of the coccoliths and discoasters act as receptors of calcite cement, whereas planktonic forams are donors of calcium carbonate. Thus, the assumption that the sediment is diagenetically unaltered is not true for the vast majority of pelagic carbonates. The other assumption that the unaltered biogenic components remain stable with burial depth and do not reach isotopic equilibrium with the interstitial waters appears to be valid at least for DSDP cores that have not been deeply buried (Lloyd and Hsu, 1972: Douglas and Savin, 1975; Matter et al., 1975). The S " 0 value of bulk samples is determined by the original 6lXOvalue of the unaltered biogenic constituents and the 6"O value of the overgrowth cement on the biogenic particles. The overgrowth cement will have a 6"O value appropriate for the interstitial water temperature at which it was precipitated. Thus at deeper burial depths. where cement is abundant and

260

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Fig. 4-45. Composite stratigraphic section of DSDP Sites 305 and 306; geochemical data. porosities. and sedimentation rates. Solid circles= Site 305; open circles= Site 306. (After Matter et al., 1975. fig. 3, p. 895.)

26 1 -

-chemistry

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Fig. 4-46. Relationship between the Sr content and depth of burial at DSDP Sites 289: 167, 305. 317, and 306. All of the samples which were analyzed by atomic-absorption spectrophotometry are largely composed of carbonate. The low Sr values of 500-600 ppm at shallow depths at Site 306 reflects the 90-100-m.y. age of these sediments and their advanced state of diagenesis. (After Manghnani et al.. 1980. fig. 8.)

1800 r

I

1

.

A 1 1, X

600 X

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SEDIMENT AGE ( m y )

Fig. 4-47. Relationship between the Sr content and age of sediment at DSDP Sites 289, 167. 305. 317. and 306. A least-squares f i t to the data shows the Sr-age relationship to be: Sr ( p p m ) = 1356-8288 (age m.y.); r = -0.875. (After Manghnani et a]., 1980. fig. 9.)

263 most of the planktonic forams have been dissolved, the 6 I S 0 values of bulk samples more closely reflect the geologic conditions during diagenesis rather than paleotemperatures of the ocean water. Only ooze which has been firmly 0

200

400

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Fig. 4-48. Relationship between the sediment age and depth of burial at DSDP Sites 289, 167, 305, 3 17, and 306 in the Pacific. These sediment sequences are largely composed of carbonate. Each site displays a unique age-depth relationship in the sequence drilled. Data was obtained from Initial Reports of the Deep Sea Drilling Project, 17, 30, 32, and 33. (After Manghnani et al., 1980, fig. 7.)

lithified on the sea floor (hardgrounds) may yield S 1 * 0 values that most closely approach the original sea-water temperatures (see also Mullins et al. (1980) for 6I8O and I3C data on hardgrounds). A number of excellent studies have been conducted on the isotope geology of deep-sea carbonates and a few examples are discussed here. In an excellent study of chalk diagenesis, Scholle (1977) synthesized a large amount of isotope data on pelagic carbonates. Figs. 4-49 and 4-50 show that as pelagic carbonates become more deeply buried, their porosity decreases and their 6 ' * 0 values become increasingly more negative. High negative oxygenisotope values can result when carbonates are deposited from fresh water or from waters at high temperatures. A gradual shift toward more negative

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30 40 50 POROSITY (PERCENT)

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70

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Fig. 4-49. Relationship between porosity and oxygen-isotope ratios (relative to PDB standard) for chalks. All data other than DSDP data are from bulk analyses of Upper Cretaceous chalk of England and Northern Ireland. DSDP data are from Tertiary chalk from Leg 12, Site 116, North Atlantic. They have been plotted as being 2.0 per mil more negative than the actual analyses indicated in order to compensate for the roughly 10°C colder depositional paleotemperatures of these Tertiary chalks relative to the Upper Cretaceous onshore samples. (After Scholle, 1977, fig. 10, p. 998; courtesy of Am. Assoc. Pet. Geol.)

oxygen-isotope values with burial depth is usually interpreted to be the result of diagenesis, i.e., when coccoliths and planktonic forams undergo dissolution the dissolved calcium carbonate is reprecipitated as cement from pore fluids at elevated subsurface temperatures. tl

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DEPTH ( M )

Fig. 4-50. Relationship between the oxygen-isotopic ratios of North Sea chalk samples and burial depth. (After Scholle, 1977, fig. 12, p. 1000; courtesy of Am. Assoc. Pet. Geol.)

265 As Scholle' (1977) pointed out, "the only exception to the rule that progressive lithification involves a shift in oxygen-isotope values is in the case of hardgrounds". Hardgrounds are cemented at the sea floor from low-temperature fluids and, thus, have relatively heavy (normal marine) oxygen-isotope values (Fig. 4-49). In fact, hardgrounds that are deeply buried and interbedded with limestones that were cemented at depth can be recognized by their isotope signatures (Scholle and Kennedy, 1974). Matter et al. (1975) demonstrated quite convincingly that most of the increasingly negative oxygen-isotope values with burial depth are consistent with progressive lithification under a normal geothermal gradient. They stated: "Our data show a systematic increase in cementation with depth which is accompanied by an equally systematic decrease in S"0". Their conclusion is further supported by the fact that the Sr2+ values in the sediment decrease with burial depth, thus paralleling the oxygen-isotope profile (Fig. 4-45). 'Some workers on DSDP pelagic carbonates have suggested that abnormally negative isotope values (e.g., in the Coniacian-Santonian carbonates) are the result of igneous intrusions, #which elevate water temperatures (Anderson and Schneiderman, 1973), or a combination of volcanic events reinforced by a global-wide decrease in the 6I8O content of ocean water (Fig. 4-51) (Coplen and Schlanger, 1973). The data obtained by Coplen and Schlanger suggest that this Cretaceous isotopic "spike" can be CENTRAL PACIFIC

CENTRAL P A C I F I C L E G XVII, SITE 167

-6

140J

-5 -4

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CARl0BEAN LEG X V , S I T E 146 '80 Relative t o PDB in */. - 7 -6 - 5 - 4 -3 -2

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Fig. 4-5 1. Variation of oxygen-isotopic composition with age for samples from DSDP Sites 44, 47, 55, 146, 153, 1'67, and 171, western Europe. (After Coplen and Schlanger, 1973, fig. 1, p. 506.)

correlated over long distances and indeed may reflect some fundamental oceanic event . Fluctuations in isotope values with burial depth and age are to be expected. They can be caused by paleotemperature fluctuations, and different localities can have different geothermal gradients or different subsurface pressure regimes. Also, lithification is not a smooth continuous process as there are bed-to-bed variations in cement percentages. This too can yield a downhole scatter of oxygen-isotope values. In spite of all the possible factors than can affect oxygen-isotope values, 6I8O data can be used to reconstruct maximum burial depths on a regional or local basis (Hancock and Scholle, 1975). Materials which do not recrystallize or re-equilibrate may show original depositional isotopic values and be useful for paleotemperature studies. Most deep-sea carbonates, however, undergo a progressive dissolution-reprecipitation process with increased burial depth and at elevated pore-fluid temperatures. Thus, the potential exists for wholesale alteration of the original isotopic values during the change from ooze to chalk do limestone. The newly-precipitated calcite cement will be in isotopic equilibrium with the pore fluids (major factors: composition and temperature) at the time of cementation. It seems, therefore, that one must be very prudent in using oxygen-isotope values from a limestone cemented at depth to determine the oceanic water temperature of the original depositional environment. I t is perhaps more likely that isotopic values in deep-sea carbonates can assist in unravelling the diagenetic history of the sediment. Only those deep-sea carbonates that can be clearly demonstrated to have been lithified at or near the sediment-water interface (i.e., hardgrounds) may be reliable for deducing paleotemperatures of oceanic waters.

DIAGENETIC POTENTIAL

Schlanger and Douglas (1974) introduced an important and very useful concept that provides a basis for understanding: ( 1 ) the progressive diagenesis of deep-sea carbonates, (2) why chalks can be interbedded between oozes, (3) the role of burial depth versus age in diagenesis, and (4)the relation between diagenesis and acoustic stratigraphy. They called this concept the “diagenetic potential” of deep-sea carbonates. In its simplest form, the diagenetic potential of a deep-sea carbonate is a measure of its geochemical-textural-constituent maturity. Conversely, given more time and burial depth, it is the sediment’s potential for further diagenetic change. A sea-floor assemblage with a high (planktonic foram + coccolith) : (discoaster) ratio has a greater diagenetic potential than one with

267 a low ratio. Likewise, given the same composition of sediment, the one that has been buried deeper and is older, will have the lower diagenetic potential. Thus, the diagenetic potential of a sediment is variable but it is highest at the moment of its inception in the surface waters. Depending on a number of oceanographic factors (productivity, depth of the lysocline and CCD, biogenic components), the diagenetic potential of ooze will have been reduced by the time it begins its burial theory at the sediment-water interface. As stated by Schlanger and Douglas ( 1974, pp. 133-134): “Diagenetic potential may thus be defined as the length of the diagenetic pathway left for the original dispersed foraminiferal-nannoplankton assemblage to traverse before it reaches the very low free-energy level of a crystalline mosaic. The concept of the diagenetic potential can be used to explain deviations from a strict depth-of-burial-lithification-dependence demanded by a simple gravitational compaction model of diagenesis. The diagenetic potential remaining to a sediment after it passes through the critical boundary between realms IV and V” (Table 4-11) “will determine how far cementation will proceed per unit time. Thus, if layers of very different diagenetic potential are buried sequentially, the amount of cementation in a layer per unit time will be proportional to the diagenetic potential of that layer so that chalks can form above oozes and limestones above chalks in the sedimentary column. The diagenetic potential (DP) can be expressed as follows: D P =f (water depth, sedimentation rate, temperature (surface), productivity (surface), (foraminifer coccolith) : (discoaster) ratio; size(,ax): size(,,,, ratio, predation rate . ..)

+

The depth of water and the sedimentation rate are very important factors as these affect the degree to which the calcite is dissolved while at the sediment/water interface; calcite dissolved at this interface is the more soluble portion of the sediment, and is not available for further diagenesis in buried strata. The diagenetic potential of a sediment is enhanced in zones of high productivity, high lysocline and low calcite compensation depths which favour the mobilization and redeposition of skeletal calcite (Wise, 1972). . . . Since discoasters appear to be the most favoured receptors of calcite overgrowths, and foraminifera1 tests and coccoliths the most called-upon donors for calcite in the cement, it seems possible that the original (foraminifera coccolith)/discoaster ratio is an important factor in the diagenetic potential because a high content of MgC0,-rich tests in a buried sediment would give it a high diagenetic potential. A sorting measure should also be included since an abundance of very small grains, mixed with larger ones such as

+

268 make up discoasters, appear to promote calcite transfer as discussed by Adelseck et al. ( 1973).. . . We postulate that variations of the diagenetic potential with depth of burial in the sedimentary column are related to original variations in basin depth, the calcite compensation depth (CCD), and the calcareous-plankton productivity of the upper water layers”.

MODEL A N D SUMMARY OF PROGRESSIVE DIAGENESIS

Schlanger and Douglas ( 1974) divided the stratigraphic position of deep-sea carbonates into seven “diagenetic realms” (see Table 4-1). They range from realm I, representing the upper 200m of the water column, where initial production of the biotic components is occurring, to realm VII, located 1-10 km below the sediment-water interface. In realm VII, the diagenetic potential is virtually zero and the once discernable textural components have been altered to a “pavement mosaic” (Fischer et al., 1967) of completely interlocking crystals, the original morphology and origin of which may not be identifiable. Schlanger and Douglas (1974) developed a diagenetic model for the ooze-chalk-limestone transition at DSDP Site 167 (Schlanger et al., 1973). This model accounts for: (1) a porosity reduction from 80% to 40% as the sediments compact and become cemented, with about half of this reduction occurring in the upper 200m of burial, (2) the gradual elimination of virtually all planktonic forams and many small coccoliths, (3) the development of interparticle calcite cement, calcite cement overgrowths on discoasters, and intrabiotic calcite cement in the remaining forams (mainly benthonic), (4) the lack of a source of extra-formational calcite for the cement (i.e., the system is calcite-conservative). The model (Fig. 4-52) accounts for all calcite in the system assuming no loss to expelled waters. If no solids are added to this system and water leaves during lithification, the volume changes during diagenesis can be approximated by the following formula, modified from Moore (1969):

= thickness of limewhere HoOze= original thickness of the ooze, Hlimestone stone derived from the ooze, ,+ , = original porosity of ooze (e.g., 80%),and +limestone = porosity of limestone (e.g., 40%). Thus, using this formula, about 3 cm3 of foram-coccolith ooze (80% porosity) will be reduced to 1 cm3 of limestone (40% porosity) during diagenesis (Fig. 4-52). There is a trend, over long stratigraphic intervals, towards decreasing porosity, increasing lithification and bulk density, and increasing compres-

t

OOZE STAGE

t

CHALK STAGE

LIMESTONE STAGE

WATER

WATER (1.3 c m 3 )

( .7 cm3)

( I - d LIMESTONE) H C H A L K = ( ~ -p C H A L K ) “LIMESTONE.

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DEEP BURIAL

/

(0-200 M E T E R S )

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9

INTER INTRA

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VOLUME (Yo) W E I G H T ( g m s )

COMPONENT

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30

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40

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9 INTRA 9

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= 65 %

9

( 4 INTER = ( 9 INTRA =

5%

= 40 %

INTERPARTICLE POROSITY 1 INTRABIOTIC POROSITY, I E . FORAM CHAMBERS)

Fig. 4-52. Diagenetic model showing the volume-weight relations in going from 3.0 crd of ooze with XO% porosity to 1.0 cm3 of limestone with 40% porosity. System is calcite-conservative. (Modified from Schlanger and Douglas, 1974. fig. 6 , p. 130: courtesy Int. Assoc. Sedimentol.)

w o\

\o

270 sional velocity. Accompanying these trends, the Sr 2 + content and 6 ' * 0 values in the sediment decrease. Although the diagenetic changes that occur, appear to be irreversible, beds in advanced stages of lithification (chalk) can be sandwiched between sediments in the incipient stages of lithification (firm ooze). This is to be expected as envisioned by the concept of diagenetic potential which can strongly influence the degree of lithification per unit of time. Within the upper 50-250 m of burial depth, gravitational compaction is a major diagenetic process. In this diagenetic realm porosity is reduced from initial values of 70-95% (Fig. 4-53) to porosity values on the order of 40-75% (Fig. 4-54).Bulk densities and velocities are commonly about 1.6 and 1.65 km/s, respectively, in this depth interval. Porosity reduction during this stage is mainly the result of compaction, which moves grains closer together by repacking, reorientation, mechanical breakage, and chemical disintegration of tests into smaller particles until a grain-supported fabric is achieved. Although dissolution and overgrowth cement do occur, in general

Fig. 4-53. Unconsolidated foram-coccolith ooze. Porosity is equal to 69% at 15.94 m burial depth. Scale bar= 12 pm. DSDP sample from Leg 30, Site 288. (After Van der Lingen and Packham, 1975. fig. 3, p. 455.) Fig. 4-54. Nanno chalk. Porosity is equal to 60% at 202.2 m burial depth. Discoasters in the center show two stages of calcite overgrowth, the right-hand side one being least overgrown. There are abundant micarb particles, partly overgrown with new calcite. Scale bar= 16 pm. DSDP sample from Leg 30, Site 288. (After Van der Lingen and Packham, 1975. fig. 6. p. 456.)

27 1

Fig. 4-55. Foram-nanno chalk. Porosity is equal to 55% at 382.76 m’burial depth. Coccoliths show corrosion and calcite precipitation; particles are welded together; large euhedral calcite crystals in the center (some having “negative crystals”) are probably all overgrown fossil particles. Scale bar=8pm. DSDP sample from Leg 30, Site 288. (After Van der Lingen and Packham, 1975, fig. 1 I , p. 458.) Fig. 4-56. Nanno chalk. Porosity is equal to 52% at 457.8 m burial depth. Radiating calcite clusters are the nannofossil Sphenolithus moriformis. These and many micarb particles show euhedral crystal faces; some coccoliths have developed extensive calcite growth along their edges (arrow). Scale bar=8 pm. DSDP sample from Leg 30, Site 288. (After Van der Lingen and Packham, 1975, fig. 14, p. 459.)

they appear to play a lesser role in porosity reduction than compaction at these shallow burial depths. In the deep-burial realm, carbonate porosities are further reduced from about 60-75% to 20-40% (Figs. 4-55 through 4-62). In this realm, however, porosity reduction occurs mainly by the dissolution of planktonic foraminifera, small coccoliths, and supersoluble micarb particles, with subsequent reprecipitation of this dissolved calcium carbonate on discoasters, large coccoliths, within robust benthonic foram chambers, and as interparticle cement. With advanced burial, pressure-solution (Figs. 4-39, 4-43, 4-6 1, and 4-63) occurs which further occludes pore space due to grain interpenetration and reprecipitation of dissolved calcium carbonate. Further burial can eventually produce a “pavement mosaic” (Fischer et al., 1967) of entirely interlocking crystals (Figs. 4-4 and 4-64). Porosities of these limestones can be lower than 5%. Bulk densities of sediment in the deep-burial realm range from about 1.6 at 200m to 1.9-2.4 at burial depths of l000m.

272

Fig. 4-57. Nanno-foram chalk-ooze. Porosity is equal to 54% at 657.7 m burial depth. Secondary calcite overgrowth on discoasters has incorporated other fossil fragments (there are also coccolith impressions in calcite). Secondary calcite fills the central parts of coccoliths and bridges the space between placolith shields. There is corrosion of some coccoliths. Scale bar=5 pm. DSDP from sample Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 45, p. 469.) Fig. 4-58. Foram-nanno chalk-ooze. Porosity is equal to 47% at 704.5 m burial depth. Strongly corroded coccoliths; at the same time there is secondary calcite growth in central areas, in between placolith shields, and on top of placoliths. Scale bar=5 pm. DSDP sample from Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 48, p. 470.)

Fig. 4-60.

273

Fig. 4-61. Nanno limestone. Porosity is equal to 19% at 1049.7 m burial depth. Advanced recrystallization of nannofossils; discoaster at left center shows impressions of coccoliths in the secondary calcite. Pressure solution is shown by an arrow. Scale b a r = I0 pm. DSDP sample from Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 64, p. 476.) Fig. 4-62. Nanno limestone. Porosity is equal to 26% at 1231.3 m burial depth. Angular coccoliths showing dissolution and calcite overgrowth; relatively large subhedral granular calcite cement; a fair amount of welding has taken place. There are many rhombic micarb particles (coccolith fragments). Scale bar=5 pm. DSDP sample from Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 78, p. 480.)

Compressional velocities at depths ranging from 200 to 1000m increase from about 1.65 km/s to 2.3-3.6 km/s or greater. The dissolution-diffusion-reprecipitation process is very effective in redistributing calcium carbonate in the sediments of the shallow- and deepburial diagenetic realms. It is also effective in carbonate conservation as the calcium carbonate of dissolved biotic elements is reprecipitated within the sediment pile. Fig. 4-59. Foram-nanno chalk. Porosity is equal to 43% at 999.4 m burial depth. Pronounced overgrowth with secondary calcite of coccoliths, discoasters, and micarb particles. There is interpenetration of calcite crystals and fossil fragments (welding); some wisp-like bridging structures (silica?) are also present. Scale bar=5 pm. DSDP sample from Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 55, p. 473.)

Fig. 4-60. Foram-nanno limestone. Porosity is equal to 30% at 1027.4 m burial depth. There are relatively large euhedral calcite crystals inside foram chambers. Scale bar = 5 pm. DSDP sample from Leg 30, Site 289. (After Van der Lingen and Packham, 1975, fig. 59, p. 475.)

274

Fig. 4-63. Recrystallized coccolith limestone showing preferential grain elongation in “NESW” direction. Some of the transversely cut coccoliths are bent and broken. and part of the elongation appears to have resulted from pressure-solution welding. Scale bar = 5 pm. Middle Cretaceous (Albian-Cenomanian, Laytonville-type limestone, Franciscan Formation, California). (After Fischer et al., 1967. fig. 40, p. 76; courtesy of Princeton Univ. Press.)

Fig. 4-64. Pavement mosaic of clear. relatively coarse grains. some of which reveal traces of sutures attesting their coccolith origin. Fabric suggests recrystallization under thermal stress. Scale bar= 5 pm. Cretaceous Calera-type limestone. Franciscan Formation, California. (After Fischer et al.. 1967, fig. 43, p. 79: courtesy of Princeton Univ. Press.)

275 Any significant amounts of recycling of calcium carbonate out of deep-sea carbonates must await uplift of the sediment to subaerial conditions or subduction into oceanic trenches where it may be absorbed in the mantle. Even when deep-sea chalks and limestones are exposed to vadose and phreatic fresh waters for tens of millions of years, alteration effects are minor (Scholle, 1977). Exceptions to this generality occur at the exposure surface itself, which may be affected by dissolution and formation of karstic surfaces (as in England) or in areas of very low rainfall where caliche crusts may form (as in west Texas and Israel, Scholle, 1977). The basic reasons for this resistance to further diagenesis include the initial stability of the low-magnesium calcite of the biotic elements and the dominance of water movement through fractures rather than through matrix pores (Edmunds et al., 1973; Foster, 1975).

ACOUSTIC STRATIGRAPHY A N D DIAGENESIS

Schlanger and Douglas ( 1 974), Milholland et al. ( 1980), and Manghnani et al. ( 1980) have lucidly described a plausible genetic relation between acoustic signatures in deep-sea carbonates and the diagenetic events that take place with burial and time based on the concept of diagenetic potential. Acoustic impedance gradients and discontinuities have generally been ascribed to cementation effects which greatly change the values of the elastic constants. Slight amounts of cement can markedly increase the rigidity of the sediment and result in a compressional velocity increase. Because acoustic impedance ( p X V ) equals density ( p ) times compressional velocity ( V ) , a difference in compressional velocity between two beds (due to subtle differences in degrees of cementation), without a marked difference in density, will cause a reflection of acoustic energy. Thus, as Schlanger and Douglas (1974) summarized: “The diagenetic potential of the sediment, as discussed previously, is in large part predetermined by oceanographic conditions prior to burial and even prior to the arrival of the sediment at the sea floor. We believe the abundant reflectors, characteristic of carbonate sequences, are related to the degree of cementation and that the degree of cementation is controlled by the diagenetic potential. Thus, since diagenetic potential is determined by palaeo-oceanographic conditions, it follows that an acoustistratigraphic event should correlate with a palaeo-oceanographic event”. Reflectors “C” and “D” of Schlanger and Douglas (1974) correlate approximately with: (1) two major temperature changes in the Pacific, (2) fluctuations in the CCD, and (3) periods of atoll emergence in the Pacific. Other studies have recognized the relation of acoustic stratigraphy to time-stratigraphy (Winterer et al., 1971;

276 Mayer, 1979). Likewise, in the Equatorial North Pacific, Piper et al. (1979) proposed an acoustic stratigraphy for the area, accessed the mappability of this acoustic stratigraphy, and related this acoustic stratigraphy to the lithologic units, as they have been identified in Deep-sea Drilling Project cores (Tracey et al., 1971; Cook, 1972, 1975).

PETROLEUM POTENTIAL

Scholle ( 1977) presented an excellent discussion of hydrocarbon production from chalks and proposed several reasons why primary porosity can be high in chalks even at depths where porosity is usually negligible. Under certain geologic conditions, deep-sea chalks can form major petroleum reservoirs. Notable examples are the Ekofisk and associated fields in the North Sea, which have recoverable reserves of about 2.5 billion barrels of oil and 8.0 Tcf of gas, all in chalks (Tiratsoo, 1976). Less productive discoveries of hydrocarbons in chalks occur in other fields of ttie North Sea, Texas, Colorado, and the Scotian Shelf of Canada. Production depths range from less than 200m to more than 4500m, with most of the North Sea production being from a depth of about 3000m. Reservoir porosities range from a high of 45% in the North Sea to 5% or less in the Gulf Coast of the United States. At a burial depth of 3000m, primary porosity should be nonexistent or very low. As Scholle (1977) pointed out, “how, then, is it possible to have matrix porosities averaging about 30% in thick chalk sections in the North Sea at 3000m or greater depth? ... There clearly must be mechanisms of porosity retention or enhancement which can affect the reservoir properties of chalks”. Mapstone (1975) showed that most of the porosity in the.North Sea chalks is the preserved primary porosity with the major part of the pore space remaining uncemented, i.e., they resemble rocks which have been buried to a depth of only 1000 m (e.g., Fig. 4-60) and not their present depth of 3000 m. High porosities in the North Sea chalks are apparently associated with overpressured zones (Scholle, 1977). In overpressured areas, hydrostatic pressure gradually approaches lithostatic pressfire. Thus, a major part of the lithostatic pressure (pressure due to the weight of the overlying rocks) is supported by the hydrostatic pressure. As a result, the driving mechanism for both gravitational compaction and pressure-solution (due to lithostatic pressure) is greatly reduced, and primary porosity can be maintained to greater depths than under a normal hydrostatic pressure gradient. Other factors also enhance primary porosity retention in deep-buried chalks. Oil entering a chalk early in its burial history will displace water. The presence of this

277 “early” oil in chalks may have prevented further dissolution-diffusion-reprecipitation processes in some North Sea chalk reservoirs (Hancock and Scholle, 1975). Also if a chalk is never buried deeply, its potential for retaining high porosities is favored. Extensive fracturing in chalk reservoirs of the North Sea and Gulf Coast of the United States has greatly improved the permeability of the chalks (Scholle, 1977).

SOME UNSOLVED PROBLEMS A N D NEW RESEARCH DIRECTIONS

With the need to expand our search for energy resources into deeper marine environments, it has become increasingly important to understand the nature and diagenetic evolution of deep-sea sediments. Although the major aspects of deep-sea carbonate diagenesis have been outlined above, many details remain to be solved. For example, what is the influence of a decreasing heat flow gradient on diagenesis with time, as sediments move laterally away from major oceanic spreading centers? What effects do siliceous biota, terrigenous components, organic matter, and water depth have on diagenesis? Also, as Matter et al. (1975) pointed out, “the operative processes and the interactions of the solids with the interstitial fluids remain almost totally unknown”. The concept of the diagenetic potential and its genetic relationship to seismic reflectors and acoustic stratigraphy is in its infancy. More work is needed to understand the relation between diagenesis and seismic reflectors. The correlation of seismic reflectors with diagenetic changes is not always clear. If most seismic reflectors in deep-sea carbonates represent diagenetic events, the question is raised: To what extent are these seismic reflectors isochronous versus diachronous? The degree to which palaeo-oceanographic events control the types and abundance of biotic components settling to the sea floor strongly influences the sediments’ diagenetic potential. This argues favorably for diagenetic seismic reflectors being essentially isochronous. Other factors, however, also affect the initial burial state of the biotic components such as water depths (e.g., its effect on sediment-water interface dissolution and grain size), bioturbation differences, bottom currents and their effect on initial textures and fabrics. All of these can modify the initial diagenetic potential. Thus, even if seismic reflectors always represent diagenetic events, this does not necessarily mean that the reflector can always be attributed to a single palaeo-oceanographic event. Subtle events near the sediment-water interface may significantly modify the diagenetic potential of sediments that began in the surface waters with the same diagenetic potential. More studies are needed that will elucidate whether or not the subtle events near the sediment-water interface are capable of altering an

278 ooze to such an extent that its original palaeo-oceanographic significance is lost or severely modified. Geoacoustic models are being developed that attempt to account for the observed property-depth-diagenesis relationship for the first 1000 m of the ooze-chalk-limestone transition. These models are designed to predict density, compressional and shear velocities, porosity, Poisson’s ratio, shear modulus, and compressional and shear velocity anisotropy (Milholland et al., 1980). Also, as Milholland and his colleagues stressed, experiments are needed where seismic frequencies are varied to determine how reflection profiles change with frequency changes and to determine if the reflectors are interference patterns (also see Mayer, 1979, for information on the acoustic properties of deep-sea carbonates).

CONCLUSIONS

(1) In deep-sea carbonates four major processes modify phySical properties, state of fossil preservation, texture, and geochemistry. These include: ( 1) gravitational compaction (mainly within the first 200 m of burial), (2) dissolution (at the sediment-water interface, shallow-burial depths, and some within the water column), ( 3 )pressure-solution (mainly within the deep-burial phase), and (4)cementation (beginning within the first few centimeters of burial and continuing with increased burial depth and age). (2) Pelagic carbonates are widespread and comprise about 70% of the carbonates deposited worldwide during the past 100 m.y. (3) Deep-sea carbonates consist mainly of planktonic foraminifera, coccoliths, and discoasters whose tests are composed of stable low-magnesium calcite. This is in sharp contrast to the highly reactive aragonite and high-magnesium calcite that comprise shallow-marine carbonates. (4)One of the most unique features of deep-sea carbonates is that they are generally not contacted by fresh water and if they are, it doesn’t occur until very late in their diagenetic history. In shallow-marine carbonates, early cementation, produced in part by contact with fresh water, preserves much of the original fabric and retards compaction in many cases. In deep-sea carbonates, on the other hand, because fresh water is absent and stable low-magnesium calcite is the main carbonate constituent, cementation is usually somewhat delayed and lithification by compaction is the rule within the upper part of the sediment column. ( 5 ) Especially since the advent of the DSDP in 1968, a few clear concepts have emerged regarding the processes involved in going from a soupy mush that flows through your fingers to a material that can be used for building cathedrals and monuments of great stability. Studies on DSDP cores have

279 dispelled longstanding misconceptions regarding the lithification of carbonates. Prior to the mid-1960’s three widespread impressions were: (1) calcite is not precipitated from sea water by inorganic means; (2) cementation of carbonates requires either diagenesis of metastable carbonate minerals in fresh water or burial to great depths; and (3) that compaction in carbonate sediments is virtually nil and, therefore, the very high initial porosities could only be reduced by cementation. This required that gigantic volumes of calcium carbonate-rich pore fluids be introduced from some unknown outside source to supply the requisite cement. All of these impressions have been shown to be incorrect for deep-sea carbonates, i.e., compaction does occur, inorganic cementation in a submarine environment does take place, and the cementation processes are calcite-conservative, with the requisite pore filling cement being derived locally by the dissolution of low-magnesium biogenic components. (6) In some deep-water settings, variable amounts of dissolution of the biotic components occur in the water column during their descent. Most of the pre-burial dissolution, however, occurs at the sediment-water interface, especially in areas where the sea floor is below the lysocline. (7) At shallow-burial depths of 50-200 m, gravitational compaction is the dominant mechanism for reducing porosity from about 80 to 60%, with cementation being a subordinate pore-reducing process. Compaction in this shallow-burial stage takes place by closer packing of grains, some crushing of microfossil tests, and disintegration of delicate planktonic forams and coccoliths into abundant micron-sized crystals (micarbs) through dissolution along their sutures. As suggested by Van der Lingen and Packham (1975). this latter process “might in part provide an answer to the problem, stated by Bathurst (1970), that no obvious signs of compaction (e.g., crushed foraminifera tests) can be detected in many lithified carbonate muds (micrites) and that, consequently, the large amount of calcite cement, filling the pore space, must have come from an outside source”. Scarcity of crushed microfossils could also be explained if the interstitial fluid pressures are essentially hydrostatic. I t certainly seems reasonable to assume that these conditions exist within the realm of shallow-burial (i.e., 100-200m). Hydrostatic water pressure, which acts at right angles to the boundary surfaces (i.e., inside and outside a foram test), is the same in all directions at any point within the fluid, and is the same at all points of equal burial depth. Thus, there would be a general absence of unidirectional or anisotropic stresses necessary to crush the microfossils. Shinn et al. ( 1977) reached the same conclusion after their experiments on compacting modern carbonate sediments resulted in large porosity reductions, with only minor evidence of crushed microfossils. (8) Discoasters appear to be the first organisms to develop secondary

calcite overgrowth. This can occur within the first 35m of burial before gravitational compaction produces a grain-supported texture, which suggests that dissolution-diffusion-reprecipitation processes occur early, while the grains are still dispersed. With increased burial depth, a grain-supported fabric develops and pressure-solution becomes a major process in producing calcium carbonate for cementation, with gravitational compaction being a subordinate process. Discoasters continue to grow by the precipitation on them of euhedral calcite derived from the dissolution and pressure-solution of planktonic forams, small coccoliths, and micarb crystals. Compaction fluids could also supply some of the Ca2+ and C0:- ions for cement. The overgrown discoasters often contain a volume of calcite several times greater than the original discoasters. Next, calcite is precipitated along the edges and in the central areas of coccolith placoliths and on micarb crystals. Coccolith cementation continues and proximal and distal shields of placoliths become connected by calcite cement. Foraminifera1 chambers begin to be filled with secondary calcite. Cementation continues to build up discoasters and coccolith placoliths, and fill benthonic foram chambers. In the most advanced diagenetic state observed in DSDP cores, (a) virtually all fossil grains are covered with subhedral to euhedral calcite cement, (b) the central areas of coccoliths are filled, (c) foram chambers are filled with granular calcite, (d) large amounts of interparticle cement are found, and (e) grain interpenetration and welding is common. Continued diagenesis produces an “ameboid mosaic” or “pavement mosaic” texture of completely interlocking grains (Fischer et al., 1967). (9) The degree to which a deep-sea carbonate becomes cemented probably depends to a large extent upon the diagenetic potential the carbonate sediment had at the time it was buried. Diagenetic potential is a measure of how much more diagenesis a sediment can undergo in the normal course of its history. A sediment’s diagenetic potential is a function of numerous parameters, most notable being: (a) the depth of water in which the sediment was deposited, (b) the fertility of the upper water level as i t affects sedimentation rate, (c) the (planktonic foraminifera small coccoliths) :(discoasters large coccoliths) ratio in the sediment, and (d) predation by plankton, as this affects the state of aggradation of the microfossils and, therefore, the settling rate of these organisms. (10) There is a trend towards a decreasing Sr2+ content and increasing SI8O values (negative) in pelagic carbonates with increasing burial depth. This is to be expected as most deep-sea carbonates are diagenetically altered after burial. With increasing burial depth, dissolution-reprecipitation processes reduce porosity. The Sr 2 + originally present in the biogenic calcite is virtually excluded from the secondary cement derived from the dissolution and pressure-solution of calcite microfossils. The SI8O values of this cement

+

+

28 1 will become increasingly more negative with burial depth, due to precipitation of the cement from pore fluids at increasingly elevated temperatures along the geothermal gradient. Marked irregular 6 ' * 0 values may reflect local events that change the sea-water temperatures (e.g., due to volcanism), or more fundamental global causes may be responsible for the aberrant S " 0 values. (1 1) Acoustic reflectors, which are probably related to the degree of cementation in the sediments, may be predetermined in their broad characteristics by major oceanic events, which in turn largely control the nature (diagenetic potential) of the biogenic material reaching the sea floor. (12) Deep-sea chalks form excellent petroleum reservoirs under certain circumstances: (a) their porosity is retained at deep-burial depths if they are hydrostatically overpressured, which decreases grain-to-grain stress and, consequently, retards pressure-solution processes, (b) early oil emplacement may retard dissolution-reprecipitation processes, (c) lack of deep burial depths, and (d) fracturing to enhance the permeability of the chalk.

ACKNOWLEDGEMENTS

The senior author greatly appreciates the prodigious effort put forth by his co-author and friend R.M. Egbert in searching the literature and compiling an intricate but very workable cross-indexed manual of references. Without his innovative compendium such a diverse and burgeoning subject would have been extremely painful for us to synthesize. Rarely have we had the opportunity to work with such a patient and encouraging editor as our good friend and colleague, George V. Chilingar. His support and gentle prodding made a difficult task enjoyable and provided us the freedom to pursue this paper in a deliberate manner. Our paper has benefited considerably from critical readings and acute comments by Hank T. Mullins of the Moss Landing Marine Laboratories, California, and George V. Chilingar of the University of Southern California. We are indebted for their prompt and painstaking reviews. So much depends on the day-to-day support of a group's secretaries. We heartily thank Theresa A. Coit and Lillian M. Wood who typed numerous drafts of this paper in a cheerful and sustaining manner. Their professionalism cannot be overstated. Special thanks go to Judy Thompson Cook who was most understanding of the piles of books, journals, and wads of paper scattered throughout the Cook household for what seemed an eternity. In addition, her own expertisc on the art and techniques of preparing review articles was most helpful, especially in the early stages of preparation of this chapter.

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289 Chapter 5 MINERALOGY AND PETROLOGY OF BURIAL DIAGENESIS (BURIAL METAMORPHISM) AND INCIPIENT METAMORPHISM IN CLASTIC ROCKS

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HANAN J. KISCH

INTRODUCTION

The last two decades have witnessed a breakthrough in the study of burial diagenesis (or burial metamorphism) * and incipient metamorphism of clastic sediments, and particularly of their mineralogical modification. A major impetus in this breakthrough stems from attempts to recognize, on the basis of these modifications, burial-diagenetic and lowest-grade metamorphic mineral zones and mineral facies, in analogy to the well-established concept of mineral zones and mineral facies in higher-grade metamorphic rocks. The developments that have contributed to making such recognition feasible include the following: ( 1) Regional studies and examination of deep-well sections have supplied a rapidly increasing volume of data indicating a regular sequence of modification of syngenetic clay minerals and clastic minerals, and the development of authigenic layer silicates in clastic sedimentary rocks during burial diagenesis and incipient metamorphism. Attempts have also been made to correlate these regularities with the changes in texture and physical properties of the rocks, which are the basis of the so-called “stages of epigenesis (or catagenesis) and metagenesis” of Soviet authors. (2) Distinctive burial-diagenetic or burial-metamorphic mineral assemblages have come to be recognized, which are characterized by diagnostic zeolites and other hydrous Ca-Al-silicate minerals in tuffs and volcanic clastic rocks. They form the basis of the zeolite, prehnite-pumpellyitemetagraywacke, and pumpellyite-actinolite-schist facies. (3) Thermodynamic data on minerals and experimental studies are increasingly used to establish the physical conditions of stability of diagnostic minerals and the equilibrium conditions of mineral reactions. (4)There has been an increasing awareness of the kinetic factors that may

’ Manuscript completed in February, 1977.

’ Burial diagenesis is used here in preference to “late diagenesis” or “catagenesis” of other authors. The terminology preferred by the author is summarized on p. 308. The editors prefer the term metagenesis for incipient metamorphism.

290 act to prevent the establishment of chemical equilibrium between all the mineral phases present in many clastic rocks modified during burial. In such cases equilibrium may be restricted to small portions (volumes) of the rocks, or affect only part of their constituents. There is also increasing awareness of the effects of compositional factors, such as chemistry of interstitial waters, upon the mineral assemblages. ( 5 )Most useful have been the advances in the knowledge of the progressive changes in the chemical composition and the physical properties of coal, bitumen, and hydrocarbons upon burial, and an awareness of the essentially “metamorphic” nature of these changes. (6) It is increasingly realized by petrologists that many of the phasemodification processes referred to as late-diagenetic or epigenetic (or catagenetic) are “metamorphic” in nature. Hence the distinction between burial diagenesis and lowest-grade regional metamorphism is largely artificial and arbitrary. Unfortunately, the modification of various kinds of sedimentary rocks upon burial has largely been studied by investigators from different sub-disciplines within the earth sciences: (1) changes in mineralogy of clay-rich sedimentary rocks by sedimentary petrographers and clay mineralogists, (2) authigenic mineral assemblages in tuffs and volcanic sediments by metamorphic petrologists, and (3) changes in the constitution of coaly and other organic matter by coal petrologists and organic geochemists. The growing awareness of the overlap in space of the physical conditions and of the chemical influences controlling these phenomena warrants renewed efforts to correlate the changes in various types of sedimentary rocks in order to arrive, ultimately, at an integrated picture of the phase changes in different types of sedimentary rocks upon burial and incipient metamorphism. The present chapter is, inter alia, intended to contribute to such an integrated picture. The author concentrated on the mineralogic changes in clay-rich and feldspathic clastic sediments, and in clastic rocks of volcanic origin, as well as changes in the constitution of the associated coaly matter. Modifications of non-siliceous sedimentary rocks (limestones and dolomites) and evaporite deposits are not covered in this chapter. Major mineralogic changes in burial diugenesis: u preview Several processes of mineral alteration and authigenesis occurring during burial diagenesis (the epigenesis or catagenesis of Soviet authors) have been described in other chapters of this two-volume book. These processes result in the progressive disappearance of clastic, syngenetic, or early-diagenetic clay minerals: (1) initially of montmorillonite and vermiculite, and (2) subsequently of irregular illite-montmorillonite mixed-layers, kaolinite, glauconite, and disordered illite ( l M , polytype). Clastic biotite and inter-

29 1 mediate and calcic plagioclase also disappear. In lithic vitric tuffs and graywackes the characteristic minerals of the zeolite facies may occur: ( 1 ) heulandite, analcime, clinoptilolite, and mordenite in the upper, low-grade part; and (2) the diagnostic Ca-zeolite laumontite in the lower half: clastic plagioclase disappears during the latter stage. By the beginning of what is henceforth arbitrarily referred to as incipient metamorphism all these minerals disappear. The predominant layer silicates of incipient metamorphism are illite-muscovite (predominantly of the 2M polytype and with high crystallinity) and chlorite. Occurring only in restricted lithologies, but of diagnostic importance, are pyrophyllite, paragonite, and the regular paragonite-montmorillonite mixed-layer rectorite or allevardite. Albite is the major feldspar present. In intermediate and basic rocks, the Ca-Al-silicates pumpellyite and epidote may occur, whereas zeolites are entirely lacking. The writer considers these changes as the first manifestations of the effects of increasing temperature and pressure following deep burial that pass by gradual transition to regional metamorphism, in the conventionally accepted sense, having more extreme conditions of temperature, pressure, and deformation. Such a concept has to make allowances for the fact that in the low-temperature range of burial diagenesis and incipient metamorphism, mineral reactions are slow and commonly incomplete. This is evidenced by the prevalence of relic phases: as a result of incomplete equilibrium only some of the mineral constituents of the original rock are affected. Taking these restrictions in consideration, the writer believes that these changes in mineralogy of clastic sedimentary rocks can be approached from the metamorphic point of view: initially incomplete, but progressively closer approximations to the establishment of metamorphic zones and metamorphic assemblages are being attained. The ultimate result of such an approach may be the establishment of mineral facies of burial diagenesis and incipient metamorphism, analogous to those established for regional metamorphism. In order to achieve this objective, the modifications affecting individual primary minerals and the appearance of authigenic minerals and mineral assemblages are reviewed here for clay-rich and volcanic-clastic rocks separately in the first instance, in order to arrive at a mineral zonation for each of these two main groups of rocks. Subsequently, the changes in different rock types interbedded in some sections are related to each other and to the change in degree of coalification of associated coaly matter, in order to develop a scheme for the integrated picture of phase changes. Before attempting to do this, however, it is important to review the earlier suggestions of the metamorphic nature of diagenetic reactions and of the limits between diagenesis and metamorphism. In addition i t is necessary to clarify the nomenclature used for various stages in this transitional realm between sedimentary and metamorphic mineralogy and petrology.

DIAGENESIS AND LOWEST-GRADE METAMORPHISM: SCOPE AND LIMITS

Attainment of chemical equilibrium upon burial

Diagenesis and metamorphism both involve chemical and physical changes of rocks and their mineral phases in response to physical and chemical conditions different from those under which the sediment was deposited, exclusive of the effects of weathering. At the time of deposition, the mineral constituents of a sediment may have been either in equilibrium or out of equilibrium with their sedimentary environment and with each other. The first is more likely to be the case with chemically deposited carbonates, evaporites, and sulphides, and to some extent with chemically deposited cements of clastic sedimentary rocks. Detrital minerals of igneous origin and volcanic constituents, on the other hand, which formed at much higher temperatures than that of the depositional medium, were generally not in equilibrium with their low-temperature sedimentary environment, unless such equilibrium has been lestablished through alteration in the course of their transport and sedimentation. Equilibrium processes in burial diagenesis and metamorphism: similarities and qualitative differences As the sediment is buried, it is subjected to physical conditions and chemical environments progressively departing further from those that prevailed during deposition and early diagenesis. Temperatures and pressures increase, and chemical parameters in the environment, such as the composition of interstitial solutions, change. As a result, many of the initially stable minerals become unstable (or metastable), whereas the stability conditions of many of the second group of minerals are still far from being attained. The mineral reactions taking place in both burial diagenesis and regional metamorphism are a response to this instability, and tend to establish or re-establish equilibrium between the various phases, and between the phases and the environment. The reactions that occur, initially between the minerals and their chemical environment (particularly the interstitial solutions) and, subsequently, between the different mineral constituents may be regarded as (1) progrude in the case of minerals that were in equilibrium with the lower-temperature environment of deposition or early diagenesis; and (2) retrograde in the case of igneous and metamorphic minerals that crystallized at much higher temperatures. The rate at which the mineral reactions proceed depends not only on the extent of instability of a particular mineral or mineral assemblage under given physical conditions and chemical environment, but also on kinetic

293 factors (reactivity), which play an important role. Inasmuch as reaction rates increase with rising temperature, kinetic factors are particularly important at low temperatures. Thus, the reaction rates differ widely for different starting minerals. For example, (1) the “prograde” dehydrations of gypsum to anhydrite and of montmorillonite to illite-montmorillonite mixed-layers, or (2) the “retrograde” zeolitization or hydrolysis of volcanic glass (a particularly reactive material), may be well advanced or essentially completed under conditions at which primary clay minerals such as kaolinite, or detrital minerals of igneous and metamorphic origin such as pyroxenes, amphiboles, K-feldspar, and staurolite, though instable, show only minimal signs of alteration. These differences are particularly apparent where minerals showing such different alteration rates occur in the same rocks or the same sedimentary or clastic volcanic sequences. It must also be noted that most prograde mineral reactions involve dehydration, whereas most retrograde reactions are based on hydration. Lack of water, due to the low porosity or permeability, may thus materially retard hydration of detrital minerals. The presence of unstable or metastable mineral relics in any rock reflects failure to attain complete equilibrium due to slow reaction rates (kinetic factor). Their occurrence in mineral assemblages indicates that the latter are not exclusively a function of the bulk chemical composition and the physical conditions of recrystallization, which is inherent in the concept of metamorphic facies, but also of its previous mineral composition. The extent of persistence of such relics is one of the important qualitative differences between diagenetically altered and regionally metamorphosed rocks. Whenever the reactions have proceeded to such an extent that an equilibrium mineral assemblage encompassing the whole rock is essentially established (i.e., without unstable or metastable relics of primary minerals), the result is generally considered a metamorphic rock. On the other hand, in most diagenetically altered clastic sedimentary rocks, unstable or metastable mineral relics are much more prevalent, reflecting the greater impact of reaction kinetics at low temperatures. It must be pointed out, however, that this difference is only qualitative: relics persist not only in the case of retrograde metamorphism, but are also present in many prograde metamorphic rocks, as zoned crystals of garnet, plagioclase, and amphiboles. Inasmuch as relics are more common in sedimentary than in metamorphic rocks, in diagenetic processes kinetic constraints appear to play a more important role than in metamorphic processes. Thus, diagenetic assemblages are to a greater extent a function not only of bulk chemical composition but also of the mineralogic nature of the primary constituents. The qualitative differences between diagenetic and metamorphic processes, however, are so gradual that it is impossible to establish sharp boundaries. Neither is such a sharp delimitation warranted. Various concepts of diagenesis are discussed next.

294 Concepts of diagenesis

Development of the concept of diagenesis has been reviewed by Dunoyer de Segonzac (1968), and the present section will focus mainly on the nature and definition of the high-grade limit of diagenesis. Diagenesis was established as a defined geological concept by Walther ( 1893- 1894) to encompass: “. ..all those physical and chemical changes that a rock undergoes after its deposition, without the intervention of orogenic compression and volcanic heat”. Since that time, this concept and definition of diagenesis has been implicitly or explicitly understood to exclude the effects of deformation (i.e., slaty cleavage and schistosity) and metamorphism, and forms the basis, with minor modifications, of current definitions of diagenesis by most West European and a large number of American authors (e.g., Correns, 1950; Williams et al., 1954; Pettijohn, 1957; Krumbein and Sloss, 1963). Most of these authors note the transitional nature of the gradation from diagenesis to metamorphism, which is due to the similarity of the processes operating in both domains. For example, Correns (1950) stated that: “The gradation (of diagenesis) into metamorphism is uncertain. This is in the very nature of things, since diagenesis is fundamentally nothing but a form of metamorphism.. .” Diagenesis has also been, and is still being extensively used in a much more restrictive sense, for the earliest stages of evolution of sediments up to lithification. This definition is accepted by many American authors, following Twenhofel’s ( 1926) exclusion of post-lithification processes from diagenesis proper (an exclusion admittedly qualified subsequently by the same author: Twenhofel, 1950, p. 276), and by most Soviet authors following Strakhov (1958 and earlier papers) and Teodorovich (1961). This restriction, however, is not used in Western Europe. In its narrower sense diagenesis is approximately equivalent to the earlv diugenesis of the authors adopting the more inclusive concept of Walther (1893- 1894); i t excludes the late diugenesis of authors such as Williams et al. (1954, p.263), and the equivalent epigenesis (Pustovalov, 1955; Teodorovich, 196 1 ; Rukhin, 196 I), catugenesis (Strakhov, 1958), or epigenesis and metagenesis (Kossovskaya and Shutov, 1958, 1963, 1970) as used in the U.S.S.R., rnetharmosis of Kessler (1922), and epigenetic diagenesis of Packham and Crook (1960). Although this usage has certain advantages, as discussed below, it has not been accepted by most West European and many American authors and is not used in this chapter.

295 Arbitrary nature of the distinction between diugenetir and incipientmetamorphic processes Williams et a]. (1954, p. 263) noted that certain diagenetic reactions “take place long after deposition and constitute what may be called lute diugenesis. Some of these reactions occur after the sediment has been buried for some time, or even while it is being deformed. If they occur at great depth below the surface, the pressures and temperatures are higher than those existing at the surface, and diagenesis then grades into metamorphism”. Moreover, Williams et al. ( 1954, p. 162) stated that “low-temperature metamorphism, of course, cannot be sharply distinguished from sedimentary diagenesis; the same or similar minerals (e.g., albite, quartz, zeolites, “chlorites”, and carbonates) can be formed by both processes”. They (p. 209) conclude that: “Indeed, distinction between diagenesis and incipient metamorphism of clay sediments is arbitrary”. Pettijohn (1957) refers to the reactions that take place between the various components in partly or wholly mechanically accumulated sediments as a result of rise in temperature or pressure of a suitable’ medium as ...diagenetic at lower temperatures and pressures, and metamorphic at more elevated temperatures and pressures”. As far as the attempts at distinction between diagenesis and metamorphism are concerned, however, Pettijohn (1957, p. 648) concluded: “All writers would exclude metamorphic changes from the domain of diagenesis, but as pointed out by Deverin (1924), no distinction between diagenesis and metamorphism is possible. Diagenesis is. in fact, the beginning of metamorphism because it leads to modification of textures, structures, and mineral composition of a sediment. Such modifications are the earmarks of metamorphism according to Grubenmann”. Processes occurring during metamorphism, such as recrystallization and flow, take place in essentially solid systems, in part under the influence of non-hydrostatic stress. Turner (1948; 1968, p. 3) defines metamorphism as: “. ..the mineralogical and structural adjustment of solid rocks to physical or chemical conditions that may have been imposed at depths below the surface zone of weathering and cementation, and which differ from the conditions under which the rocks in question originated”. Inasmuch as minor amounts of supercritical, predominantly aqueous pore fluids participate in most metamorphic transformations, as they do in diagenetic processes, even the essentially solid state of the phases in metamorphic processes is not a major distinctive criterion. The above definition would include most processes operative after cementation, i.e., during burial diagenesis (or epigenesis). Many metamorphic petrologists are, therefore, well aware of the arbitrary nature of the limits set between diagenetic and metamorphic rock transformations. As stated by Turner (1968, p. 2): “...the student of diagenesis and “

the petrologist in the incipient stages of metamorphism may find themselves studying the same or closely related phenomena”. This awareness is to a large extent due to the pioneering studies of Coombs (1954, 1961) and Coombs et al. (1959) on the zeolite assemblages formed upon burial in volcanic clastic rocks, and to studies by Soviet authors on the textural and clay-mineralogical transformation of clastic sediments during burial diagenesis (“epigenesis”) and incipient metamorphism or “metagenesis” (e.g., see Kossovskaya and Shutov, 1961, 1963). Coombs ( 1961) was the first investigator to explicitly include burial diagenesis in his model of metamorphism, as stated in the following paragraph: “Certainly any complete model of the metamorphic process must make provisidn for all changes accompanying rise of temperature and pressure following burial, as well as subsequent events which have left their impact on the mineralogy of the rock. Here then is the point at which, while I am considering metamorphic processes, I propose to draw the line marking the onset of metamorphism. It is at the stage at which mineralogical or textural modification first sets in, in response to a temperature differing appreciably from that under which the sediment was deposited, or at which a volcanic rock has consolidated. Not only is this a point beyond whch all changes should be included in any complete model of the metamorphic process, it is also the only point at which I see hope of drawing any reasonably objective line of demarcation in the restricted sense” (p. 214). Coombs thus excludes from his model diagenetic processes in the most restricted sense of that term, i.e., processes occurring essentially at the temperature of deposition, but before complete consolidation or lithification. For mineralogic and textural metamorphic modification taking place in sedimentary columns without intervention of significant penetrative deformation, and excluding early diagenetic processes, Coombs ( 1961) proposes the name burial metamorphism. This term then is approximately equivalent to the late diagenesis of Williams et al. (1954), the epigenetic diagenesis of Packham and Crook (1960), and the epigenesis (or catagenesis) of Soviet authors. The present writer will use these terms as virtual synonyms. There thus seems to be a fair measure of agreement that there is no essential difference between the nature of the processes which are operative during late diagenesis or burial metamorphism on one hand and incipient regional metamorphism on the other, and that the two domains show a wide range of overlap. As discussed later, the two domains in practice can be separated (e.g., see Winkler, 1965, 1967) only by arbitrarily postulating specific mineral assemblages in rocks of suitable chemical bulk composition, which are believed to be indicative of a specific range of physical conditions, to be diagnostic of metamorphism. The writer believes that such a distinc-

297 tion is a fruitless exercise in semantics. The reference to essentially identical processes as “late diagenesis” or “metamorphism” has rightly been noted by Coombs (1961, p. 213) to be due to the traditional outlook of sedimentary petrologists and clay mineralogists on the one hand, and of metamorphic petrologists and mineralogists on the other:“ ...It is fair to say that in the past most metamorphic petrologists have been preoccupied with the problems of relatively coarsely crystallized metamorphic rocks, whereas most sedimentary petrologists were concerned with the nature and origin of sedimentary rocks as sediments. There has been a hiatus in approach, scope and objectives between workers in the two fields. A cynic could easily say that diagenesis includes the processes of rock alteration studied by those trained as sedimentologists, and metamorphism processes studied by those trained as petrologists of crystalline rocks”. It is a challenge to the student of incipient metamorphism to try and bridge this hiatus and arrive at a synthesis of these points of view. Metamorphic versus sedimentary rocks: petrographic criteria

Apart from referring to processes in a more or less arbitrarily delimited range of physical conditions, the term “metamorphic” is also applied to rocks. The labelling of processes and that of rocks as metamorphic is commonly but not necessarily congruent, particularly in the lowest grades of regional metamorphism. Metamorphic processes as conceived by the writer cover a wider range than do the rocks conventionally referred to as metamorphic. Metamorphic and apparently non-metamorphic rock types are closely associated in many sequences exposed to low-grade metamorphic conditions. Williams et al. (1954, p. 209) mention as examples: “...the development of glaucophane schists in otherwise unmetamorphosed sandstones, diabases and basalts. Fine-grained mudstone, also (a rock sensitive to temperature change), may be metamorphosed to slate when associated sandstones show only general induration.” There is no difficulty in distinguishing between diagenetically altered sedimentary rocks and metamorphic rocks when the degree of diagenesis is low or the degree of metamorphism is high. Whenever the reactions between the phases in clastic rocks have proceeded to such an extent that an equilibrium mineral assemblage regarded as metamorphic is established, and if the rocks have acquired a metamorphic fabric in response to recrystallization, the result will generally be considered a metamorphic rock. It should be noted here that this is not necessarily the case for non-clastic sedimentary rocks such as evaporites, carbonate rocks, and coal. There is a definite range of clastic rocks, however, for which the classification as either diagenetic or metamorphic is not so simple. Whether the

298 product of partial modification of a sedimentary rock by metamorphic processes will be considered a “metamorphic rock” will depend on the adopted petrographic criteria of distinction between sedimentary and metamorphic rocks. Such criteria may be of the following types: ( 1) Extent of modification of primary sedimentary texture and structure due to (a) recrystallization of the major phases or marked reaction between grains of different authigenic or clastic mineral phases (chemical reconstitution), or (b) development of metamorphic structure, including schistosity and slaty cleavage. (2) Metamorphic nature of the minerals or mineral assemblage formed in the rock after deposition, i.e., “metamorphic” mineral facies. Problems arise in the case of conflict between these criteria or when they are indeterminate, for instance: (a) When the metamorphic modification of the primary mineral assemblage is well advanced, whereas the primary sedimentary or volcanic texture has largely been conserved, with no development of slaty cleavage or schistosity. This is the case in many zeolite and in some glaucophanelawsonite-schist facies rocks. (b) When the rock has a largely metamorphic texture, but the mineral assemblage is not diagnostic, i.e., may be either primary-clastic or metamorphic. This is the case in many slates in which chlorite and illite are the main layer-silicate minerals. Before discussing the criteria proposed to characterize a rock as “metamorphic”, a few extreme cases should be excluded from consideration in order to comply with generally accepted geological usage. These include rock types that upon shallow burial or early diagenesis completely lose their primary sedimentary structure and texture through recrystallization. An extreme case of such early complete recrystallization is glacier ice, which would be considered by few geologists to be a metamorphic rock. Chemical precipitates, which were in equilibrium with their depositional environment and formed rocks such as many evaporites and gypsum, may recrystallize completely to a new mineral assemblage upon shallow burial. To some extent, carbonate sediments that underwent dolomitization during diagenesis also fall in this category. All these processes may take place at very low temperatures and in association with clastic rocks that are as yet incompletely lithified (i.e., are still in the stage of early diagenesis). The following discussion is, therefore, restricted mainly to clastic rocks, including tuffs and volcanic clastic rocks. Attention is paid, however, to coal and coaly matter, which largely loses its primary texture during early burial, due to its compactability, with marked changes in its chemical composition due to coalification. The effects

299 of deep burial on the chemical composition and physical properties of coal are discussed in chapter5, Vol. I of this book. Diagenetic versus metamorphic: priority of mineralogical reconstitution or of textural modification in distinguishing diagenetic and metamorphic rocks Common metamorphic rocks are characterized by a metamorphic structure, such as slaty cleavage or schistosity, and by a crystalloblastic texture as a result of recrystallization and mutual reaction of the phases, i.e., the chemical reconstitution. Turner (in Fyfe et al., 1958, p. 215; and in Turner, 1968, p. 264) bases the distinction between deep-seated diagenesis and metamorphism mainly on the degree of chemical reconstitution. He considers changes as metamorphic rather than diagenetic “when the coarse clastic grains also are involved extensively in reaction so that the rock becomes substantially recrystallized”. In accordance with these criteria, Turner classifies reactions that took place independently (i.e., restricted to one mineral and its immediate environment), such as development of analcime and heulandite at the expense of volcanic glass and the incipient albitization of plagioclase, as diagenetic rather than metamorphic. In contradistinction, he considers the reactions leading to the characteristic assemblages of the zeolite facies sensu stricto, such as laumontite-albite-quartz and quartz-albite-adularia-pumpellyite, and the complete albitization of plagioclase to be the result of interdependent though localized reactions and, therefore, metamorphic (in Fyfe et al., 1958, p. 217). Inasmuch as shearing and ruptural deformation reduce grain size and accelerate low-temperature reactions, Turner (in Fyfe et al., 1958, pp. 215-216; 1968, p. 264) accepts schistosity as a criterion of metamorphism. He points out, however, that under favourable conditions chemically unstable sedimentary rocks may become completely converted into the lowtemperature zeolite-facies mineral assemblages without the aid of deformation and, thus, without the development of schistosity. While noting the similarity of these mineral assemblages to those formed during diagenesis, Turner (in Fyfe et al., 1958, p.216) classes such rocks as metamorphic, although later he admitted the view that “. ..whether the resulting mineral facies (in this case the zeolite facies) are termed diagenetic or metamorphic is of no great significance” (1968, p. 264). The present writer is in full agreement with the latter statement. Packham and Crook (1960), in contrast, place the boundary between diagenesis and metamorphism primarily on the basis of fabric, referring to processes taking place until the original fabric is extensively modified as diagenetic. Thus, certain mineral assemblages characteristic of some particu-

lar temperature range may appear in rocks with hornfelsic fabric (contact metamorphism), original clastic fabrics (diagenetic modification), or schistose fabric (regional metamorphism). Packham and Crook (1960, pp. 403-404) object to Turner’s (1958) usage, which “results in the classifying of many typical sedimentary sequences as regional metamorphics, on the basis of their secondary mineral assemblages. Such an extension of the usage of “metamorphic” is not only confusing, but it may obscure the fact that both typically schistose rocks and typically sedimentary unschistose rocks may develop the same mineral assemblages”. One may ask, however, whether this “obscuring” is all that is undesirable. From a mineral-facies point of view, the similarity of the mineral assemblages is a more important attribute than is the extent of modification of the original sedimentary fabric and the objection then appears to be invalid to the writer. Non-schistose rocks, in which diagnostic mineral assemblages have formed as a result of interdependent reactions of the clastic constituents, would thus be referred to as diagenetically altered sediments by Packham and Crook and as metamorphics by Turner. The two delimitation proposals illustrate the consequences of the difference in point of departure in the choice of the primary criterion for metamorphism, i.e., extent of chemical reconstitution or extent of textural modification. Obviously, there are some types of rocks to which one or both of the criteria are inapplicable. For instance, the major minerals of rocks such as quartzites, chlorite- or illite-rich shales, and limestones remain stable under the conditions of lowest-grade regional metamorphism and, thus, do not give rise to diagnostic metamorphic mineral assemblages. On the other hand, in monomineralic rocks such as orthoquartzites and “marmorized” limestones the textures due to diagenetic effects, such as secondary overgrowths and crystallization, without exposure to “metamorphic” temperatures may be indistinguishable from those induced by metamorphism. Coombs ( 196 1 , p. 2 13) includes all mineralogic and textural transformations following burial to depths where temperatures differ from those of deposition as metamorphic. He points out that the metamorphic modification of the texture of a sediment is essentially progressive, though it may proceed in fits and starts. Consequently, he applies the degree and type of textural modification, such as the development of schistosity, only for a loosely-defined subdivision of metamorphism into intergradational categories, i.e., for distinction between burial metamorphism and regional metamorphism. Coombs, however, considers such qualitative criteria as unsuitable for a primary classification of fundamental earth processes and, therefore, rejects Packham and Crook’s suggestion of a textural delimitation of diagenesis and metamorphism as “foredoomed to failure”. Non-schistose

301

TABLE 5-1 Textural zones (mainly in Verkhoyanye, Siberia), major alteration processes in terrigenous rocks, and porosity of sandstones, after Kossovskaya et al. (1957) and Kossovskaya and Shutov (1959, 1970); correlation with “stages of regional epigenesis and metagenesis” after Kossovskaya and Shutov ( 1961, 1970) Textural zones, major alteration processes in terrigenous rocks, and sandstone porosity

“Stages”

I . Zone of unultered cluy matrix. Original texture and mineralogy essentially preserved; clayey matter in matrix of sandstones and in clays inherited from the stages of sedimentation and early diagene sis; gradual disappearance of some unstable clastic minerals (pyroxene, amphibole, intermediate and calcic plagioclase); progressive compaction downwards; porosity of sandstones decreases from 40-20% to 20-8% in lower part.

initial epigenesis

2. Zone of ultered cluy matrix. Hydromica and chlorite are the main components of argillites; beginning hydromicatization of kaolinite and disappearance of montmorillonite in argillites (in deeper stages also as component of mixed-layer clays); chloritization and hydro micatization of biotite; recrystallization of leucoxene to anatase and brookite; porosity of sandstones decreases from up to 20% to 4-5% and less in the lower part.

3. Zone of yuurtzite-like mosuic (und conformul-regeneruted) structures in sandstones, und chlorite-hydromica matrix. Shaly argillites and clay slates; recrystallization of clastic fragments in sandstones (but original texture preserved); sericitization of primary clay is intensive in the matrix of sandstones and virtually complete in argillites; strong hydromicatization of kaolinite in argillites; complete disap pearance of biotite in lower horizons (coarse clastic biotite replaced by chlorite-sericite “pockets”; porosity of sandstones: 2-4%.

deep epigenesis

4. Zone of spiny-like blastic structures in sundstones, and chloritemuscovite mutrix. Phyllitic slates; complete muscovitization and chloritization of clay matrix in sandstones; coarse lepidoblastic muscovite and chlorite in argillites; absence of clastic biotite; porosity of sandstones: (2%.

early metagenesis

5. Muscovite-chlorite suhfucies of the greenschist fucies. Rocks lose their primary-sedimentary appearance; complete recrystallization of detrital quartz (formation of lenses and veins in the cleavage); greenschist mineral composition of sandy-argillaceous rocks; pure quartz sandstones from quartzites; appearance of spessartine in Mn-rich rocks.

late metagenesis

302 rocks showing mineral assemblages of the zeolite, greenschist, and possibly glaucophane-schist facies are thus referred to by Coombs as burialmetamorphic rocks and by Packham and Crook as diagenetically-modified sediments. The “stages of regional epigenesis and metagenesis” of Kossovskaya and Shutov (1963) were also based, in the first instance, on stages of textural modification of detrital sediments. The textural zones of Kossovskaya et al. (1957; see also Kossovskaya and Shutov, 1958, 1961, 1970) with some of their mineral-alteration characteristics and correlation with the “stages” are given in Table 5-1. Incipient metamorphism or “metagenesis” was formerly considered by these authors to begin with textural zone 3 (e.g., Kossovskaya et al., 1957, Table 5-1; Kossovskaya and Shutov, 1958; a similar correlation in: Kossovskaya and Shutov, 1970, p. 16 is probably a printing error), and subsequently with textural zone 4 (e.g., Kossovskaya and Shutov, 1961; 1970, table I). The latter correlation is followed by the writer. Only subsequently have the “stages” been characterized by diagnostic authigenic mineral assemblages or “facies” that are discussed in the next section. Diagenetic versus metamorphic: mineral facies criteria In the above attempts to arrive at a boundary between diagenesis and metamorphism, the change in mineral composition has been included as an important factor. In recent years, however, German, Soviet, and French authors have endeavoured to define the realm of diagenesis mainly or entirely on the basis of mineral facies, in analogy with the conventional metamorphic facies first introduced by Eskola (1915), at least in clastic and volcanic sediments. They consider the textural modification of the sediments during compaction and lithification of secondary importance. The recognition of the zeolite facies in tuffs and volcanic clastic rocks by Fyfe et al. (1958, pp. 216-217) and Coombs et al. (1959) has been a major impetus to attempts to apply the metamorphic facies concept to diagenetic and lowest-grade metamorphic mineral assemblages in sedimentary rocks of different mineral composition, in particular clay-rich sediments. In such attempts the prevalence of metastable mineral relics, local equilibria, and other evidence of non-equilibrium (the result of incomplete reactions, formation of metastable phases, etc.) must be taken into account. As a result, many authors note that the mineral facies reactions rather than the complete recrystallization into an equilibrium assemblage are decisive. The trend of these reactions may help distinguish diagenetic and other prograde processes from weathering, in which reactions often run in a direction opposite from that of diagenesis (e.g., Steinike, 1966).

303 Kossovskaya and Shutov’s ( 1961, 1963) four stages of regional epigenesis and metagenesis are based on diagnostic mineral assemblages or “facies” in each of five compositional groups or “families” of clastic sedimentary rocks (see Table 5-11). It should be noted that the usage of the term “facies” by Kossovskaya and Shutov (1963) for a diagnostic assemblage occurring only in a restricted range of rock. compositions (“family”) is quite different from the concept of Eskola (as modified by Fyfe and Turner, 1966), in which a metamorphic facies refers to an entire natural set of metamorphic assemblages, showing a constant relation between mineral composition and chemical composition. The various “facies” in the volcanogenic graywackes correspond to the diagnostic mineral assemblages of the burial-metamorphic facies of Coombs. Winkler (1964, 1967) has gone even further and proposed to distinguish diagenesis from metamorphism exclusively on the basis of the appearance of some diagnostic metamorphic minerals. He (1967, p. 154) stated that: “Metamorphism has started or diagenesis has ended as soon as a mineral assemblage is formed which cannot originate in a sedimentary ehvironment, or when mineral assemblages restricted only to sediments disappear” ( 1967, p. 154). According to Winkler (1964, pp. 72-73), the appearance of minerals such as paragonite, pyrophyllite, Fe-rich epidote (pistacite), stilpnomelane, pumpellyite, laumontite, and lawsonite, that do not occur as early-diagenetic neoformations, as well as the last disappearance of typically sedimentary minerals such as kaolinite, montmorillonite, glauconite, diaspore, and zeolites such as analcime, heulandite, or its Si-rich variety clinoptilolite, indicate that the limit of the diagenetic range has been exceeded, i.e., that metamorphism has started. Starke (1968, p. 219) adds to these criteria the formation of 2M muscovite, the disappearance of dioctahedral 1M micas, and the replacement of the diagenetic association illite-chlorite by sericite-chlorite. The above changes take place “above 300°C” according to Starke, whereas Winkler (1967, p. 153) stated that metamorphsm sets in at about 200°C. In the discussion of the individual mineral-alteration processes, however, it will become clear that these changes do not take place within a narrow temperature interval, but cover a wide range of temperatures of burial. For instance, the disappearance of analcime and heulandite (including its variety clinoptilolite) to produce laumontite and albite takes place at much lower temperatures than does the final disappearance of kaolinite to give pyrophyllite. The lowest-grade metamorphism of Winkler, first referred to as “burial metamorphism” ( 1965, 1967) and, subsequently, as the “very-low-stage division of metamorphism” ( 1970), incorporates the laumontite-prehnite-quartz facies or zeolitic facies, the pumpellyite-prehnite-quartz facies, the lawsonite-glaucophane(-jadeite) facies, as well as the somewhat lowerpressure lawsonite-albite facies.

w

0 P

TABLE 5-11 “Facies” of regional epigenesis and metagenesis (after Kossovskaya and Shutov, 1970) Stages

Epigenesis initial epigenesis

deep epigenesis

Quartz- kaolini tic rocks

Lithoclastic sandstones

Acid arkoses

Middle arkoses

Volcanogenic greywackes

facies of detrital matter of an inherited composition

facies of vermiculitic formation with relics of femic minerals

facies of heulandite - analcime

newly-formed minerals: various composition of newly-formed minerals inherited mainly from diagenesis stages

newly-formed minerals: vermiculite-like phases at the expense of biotite, solution of femic minerals

newly-formed minerals: transformation of glass of tuffs into analcime and heulandite

structures: detrital

structures: detrital

structures: detrital

quartz - dickite facies

hydromica - chlorite facies

laumontite facies

newly-formed minerals: quartz, dickite, minerals of titanium, hematite

newly-formed minerals: quartz, albite, chlorite, hydromica (hydromicatization of kaolinite and montmorillonite), mixedlayer formations

newly-formed minerals: formation of laumontite, albite, at the expense of vitreous tuffs and feldspars; quartz, chlorite, dioctahedral hydromica

structures: mosaic, conformal-regenerated (convexo-concave contacts)

structures: mosaic, conformal regeneration

structures: mosaic, metasomatic and conformal-regenerated

Metagenesis early metagenesis

quartz-pyrophyllite furies

muscovite-chlorite facies

prehnite-muscovite-chloritefacies

newly-formed minerals: quartz, pyrophyllite

newly-formed minerals: disappearance of primary-clay and mixedlayer minerals: muscovite, chlorite, albite, quartz

newly-formed minerals: prehnite and pumpellyite (replacement of laumontite), muscovite chlorite, albite, quartz

structures: mosaic, conformal-regenerated, blastic

structures: mosaic, blastic, “spiniform” (by differential sliding)

structures: mosaic, blastic, “spiniform”

late metagenesis

epidote - muscovite - chlorite -stilpnomelane facies

newly-formed minerals; epidote, albite, actinolite, stilpnomelane, manganese garnet; wide development of chlorite and muscovite; disappearance of prehnite and replacement of potassic feldspars by muscovite structures: lepidoblastic, ocellate with segregated textures Metamorphism regional metamorphism

kyanite furies

biotite facies

newly-formed minerals: quartz, kyanite

newly-formed minerals: biotite (by interaction of muscovite, chlorite and rutile), garnets

structures: mosaic, blastic

structures: lepidoblastic with segregated textures

306 These facies are equivalent to the laumonite zone of the zeolite facies, and the prehnite-pumpellyite-metagraywacke facies sensu Coombs, and the glaucophane-lawsonite-schist facies sensu Turner ( 1968). The “very-low-stage’’ is thus considered by Winkler ( 1970, pp. 202-205) to be indicated by the following minerals: (1) laumontite (at p H z O 4 about 3 kbar), (2) prehnite and pumpellyite, (3) lawsonite (at p H z O> about 3 kbar), and (4) illite with imperfect crystallinities (Kubler’s values between about 10 and 4 mm, the latter value designating the high-grade limit of the anchimetamorphic zone). Other minerals that are considered to appear in parts of the “very-low-stage” of metamorphism, but to persist to different extents into the higher-grade “low-stage’’ facies, include pyrophyllite, glaucophane, jadeitic pyroxene, and stilpnomelane (Winkler, 1970, pp. 204, 225-226). Winkler (1970, p. 204) also claims appearance of chloritoid in the “verylow-stage’’ of metamorphism. The present writer considers this typical greenschist-facies mineral not to occur at lower grades of metamorphism, with the possible exception of rocks extremely rich in Mn. Unfortunately, with the exception of illite the distinctive minerals listed above appear only in rocks rich in intermediate or basic volcanic material. Strict application of Winkler’s mineral-facies criteria would assign many of the associated pelitic rocks (which, for instance, in the laumontite zone may still abound in sedimentary clay minerals such as montmorillonite) to diagenesis. Such rocks cannot be regarded as metamorphic on the basis of their own mineral facies, but can only be postulated to be such on the grounds of their association with some other rock types containing a diagnostic “metamorphic” mineral assemblage such as laumontite-quartz or pumpellyite-quartz. Winkler’s ( 1970) criteria for the onset of “metamorphism”, therefore, are directly applicable to a very restricted range of rock compositions and only on the basis of association with other rocks, at least in the lower grade of this group of metamorphic facies. Kubler (1967a, 1967b, 1968), followed by other French and Swiss authors (Dunoyer de Segonzac et al., 1968; Artru et al., 1969; Dunoyer de Segonzac, 1969, 1970; Frey, 1970) have introduced the concept of “anchimetamorphic zone” or “anchizone” intervening between the zone of late diagenesis and the epizone of regional metamophism (including the greenschist facies), reviving a term coined by Harrassowitz in 1927. Kubler’s approach makes use of the increase in “crystallinity” of illite during diagenesis and lowestgrade metamorphism, as monitored by the concomitant decrease of the width of the illite (10 A) X-ray diffraction peak at half-peak height. This anchimetamorphic zone is defined primarily by limiting values of the illite “crystallinity” index (Kubler, 1967a, p. 1 1 1; 1968, pp. 392-395). Pyrophyllite may appear, whereas kaolinite-group minerals and irregular illitemontmorillonite mixed-layers are usually absent.

307 Inasmuch as the low-grade limit of the anchimetamorphic zone corresponds closely to the appearance of pyrophyllite at the expense of kaolinite and its polymorph dickite in aluminous sedimentary rocks, which takes place at a more advanced stage of burial than does the alteration of kaolinite and montmorillonite to illite and chlorite (cf. Starke, 1968; K m h , 1969), Kubler (1967a, pp. 110- 11 1; 1968, pp. 394-395) correlates the zone with the “early metagenesis stage” of Kossovskaya and Shutov ( 1961, 1963, 1970). The illite-“crystallinity” method has the advantage of being applicable to a much wider range of clastic sediments than is the low-grade mineral facies method in volcanic sediments. Low potassium contents and abundance of clastic muscovite, however, may cause anomalously high and low crystallinity indices, respectively. The limitations of this method are discussed in greater detail in a later section of this chapter. Inasmuch as the “stage of early metagenesis” and the lower-grade “stage of deep epigenesis” are believed by Kossovskaya and Shutov (1961 and later papers) to encompass Coombs’ prehnite-pumpellyite-metagraywacke facies and laumontite zone of the zeolite facies, respectively, the suggested equivalenke of the anchimetamorphic zone with the “early metagenesis” stage would provide a correlation between the illite-“crystallinity” zones and the low-grade facies in volcanic sediments. The acceptance of such a correlation has apparently led Winkler ( 1970) to include a specified range of illite-“crystallinity indices” as one of the diagnostic criteria for his “very-low-stage metamorphism”. Although Kisch (1974) has shown that the anchimetamorphic zone as based on illite “crystallinities” and the appearance of pyrophyllite, on one hand, and the prehnite-pumpellyite-metagraywacke facies, on the other, are associated with a similar range of high coal ranks, direct correlation studies of diagnostic low-grade zones in the same sequences are lacking (but see Kisch, 1980b). Such studies will be required before this correlation can be considered to have general validity. In fact, unless the diagnostic mineral equilibria are affected in the same way by temperature and pressure, i.e., have the same d T/d P relationships, they may well have different relationships under various geothermal gradients. In the writer’s opinion, until well-defined diagnostic “metamorphic” mineral-transformation stages are firmly established to occur at approximately similar stages of deep burial or incipient metamorphism in various areas (at least in clay-rich clastic sedimentary rocks and volcanic sediments), low-grade mineral facies (though invaluable for the establishment of diagenetic and lowest-grade metamorphic zoning) should not be used as the criterion to label rocks as “metamorphic” or “non-metamorphic”, or to establish a boundary between the diagenetic and the metamorphic realms. In collectively attributing the lowest-grade metamorphic facies to “burial metamorphism”, Winkler ( 1965, 1967) has appreciably modified the meaning of this useful term as proposed by Coombs (1961); he defined it on the

basis of a temperature range and assigned the absence of schistosity, which is a major factor in Coombs’ definition, a second place: “In general, the absence of penetrative movement precludes a schistose structure. Although the original fabric of the rocks may be largely preserved, the mineralogical composition has been changed. Metamorphic changes are hardly visible in hand specimens; only in thin sections can they be clearly recognized” (Winkler, 1967, pp. 4-5). Arguing that in many occurrences the lawsonite-glaucophane-schist facies rocks show no schistosity, Winkler included this facies, as well as the somewhat lower-pressure lawsonite-albite facies, in the “burial metamorphism” stage. This inclusion would also apply to the many occurrences common schistosity of which originally gave rise to the name glaucophaneschist facies, whereas according to Coombs’ concept schistose gluacophaneschists should be regarded as regional-metamorphic and not as burial-metamorphic. Although rocks such as sandstones that are associated with glaucophane-schists may hardly show any textural imprint of metamorphism, it is very doubtful if any mineral assemblage containing the characteristic high-pressure minerals glaucophane, lawsonite, or jadeite, can form as a result of mere burial without the action of tectonically-controlled pressure (cf. De Roever, 1972). In the writer’s opinion, Winkler’s extension of “burial metamorphism” to cover series of predominantly schistose rocks of the glaucophanelawsonite-schist facies is at variance with Coombs’ concept, divests it of most of its usefulness, and should not be used. The proposed distinction between diagenesis and burial metamorphism is arbitrary and applicable directly only to a very restricted range of rock compositions. It is also undesirable in that it reduces the usefullness of the term “burial metamorphism” to cover all late-diagenetic and lowest-grade metamorphic processes up to the development of schistosity. Kubler (1967b) has shown that correlation between schistosity and illite-“crystallinity” indices were unsuccessful: cases are known of anchimetamorphism without development of schistosity and presence of schistosity without anchimetamorphism. Apparently, the dynamic factor which plays an important role in the development of schistosity, has hardly any effect on the development of illite crystallinity. The writer believes that there is no close parallelism between the development of metamorphic fabric and diagnostic mineral modifications. All attempts to attach a mineral-facies interpretation to a texturally delimited stage, such as burial metamorphism, therefore, must be rejected. Terminology

It may be useful, in summary, to present the terminology of burial diagenesis and lowest-grade metamorphism as used in this chapter.

309 The author uses the terms referring to specific stages, mineral zones, or mineral facies only for sequences containing the mineralogic criteria defining these stages, etc., and not for sequences that have been extrapolated from the presence of other criteria to have reached a similar degree of lowest-grade a1teration. Thus, the lowest-grade metamorphc facies are used only for sequences containing the. diagnostic zeolite and hydrous Ca-Al-silicate minerals or mineral assemblages. “Anchimetamorphic” or “anchizone” is recognized only on the basis of illite-“crystallinity” data and the appearance of pyrophyllite or rectorite-allevardite. The “stages of regional epigenesis and metagenesis” are defined on the basis of Kossovskaya and Shutov’s criteria. Usage of general terms in a more specific sense, or one departing from common practice, will be marked by parentheses. Such terms include “diagenetic zone” and “epizone”, as based on illite-“crystallinity” ranges, and the “facies” of Kossovskaya and Shutov. Burial diagenesis and burial metamorphism are used according to Coombs (1961) definition for partial or complete reconstitution, commonly on a regional scale. Characteristically there is no development of a penetrative fabric such as schistosity. There is also no mineral-facies connotations. These terms are virtual synonyms of late diagenesis and epigenetic diagenesis. Incipient metamorphism is used as a general term for the more advanced stages of mineral modification as characterized by the appearance of the attributes of anchimetamorphism. Thus it is approximately equivalent to the “stage of early metagenesis”. This term has very generalized mineral-facies connotations: Of Winkler’s ( 1970) “very-low-stage’’ metamorphism, it includes at least part of the prehnite-pumpellyite-metagraywacke facies and, probably, all of the lawsonite-albite-schist and glaucophane-lawsonite-schist facies. Slaty cleavage or schistosity, at least in pelitic rocks, is a common though not an essential attribute of incipient metamorphism. Lowest-grade metamorphism, finally is used as a general term covering all the above stages and metamorphic facies, up to the onset of the greenschist facies.

MODIFICATION OF CLAY MINERALOGY OF CLASTIC SEDIMENTARY ROCKS UPON BURIAL DIAGENESIS (BURIAL METAMORPHISM) AND INCIPIENT METAMORPHISM

By the appearance of the greenschist facies, which is regarded as signalling the onset of traditionally-recognized regional metamorphism, clastic sedimentary rocks have undergone major changes in mineralogy. Minerals such

3 10

as smectite, kaolinite, irregular illite-smectite mixed-layers, as well as clastic biotite, calcic plagioclase and K-feldspar, have entirely disappeared. Muscovite and chlorite have become the predominant layer-silicates in most clastic metasediments. Although minerals such as pyrophyllite and paragonite may be less abundant, they are highly characteristic constituents. Most of these changes d o not take place suddenly: many take place over wide range of burial conditions, and are in fact largely completed during burial diagenesis (burial metamorphism); others mainly occur in that ill-defined stage between diagenesis and conventional regional metamorphism variously referred to as incipient metamorphism, very-low-grade metamorphism, anchimetamorphism, metagenesis, etc. In this section, the writer discusses these various changes. Disappearance of primary clay minerals and some clastic constituents during burial diagenesis, as amply documented by mineralogic studies in deep oil wells, are discussed first, followed by the mineralogic characteristics of incipient metamorphism. Modification of primary mineralogy of clastic sedimentary rocks during burial diagenesis

The present sections deal with mineral modifications which occur during burial diagenesis for each one of the major primary, early-diagenetic, or clastic silicate mineral. The major alteration processes taking place during this stage include: ( 1) Replacement of smectite through progressive mixed-layering by illite and chlorite. (2) Replacement of kaolinite group minerals by illite and chlorite. (3) Increasing three-dimensional order and modification of the polytype of kaolinite group minerals. (4) Progressive change in the composition and polytypes of chlorite. ( 5 ) Progressive alteration of clastic biotite through vermiculite to chlorite. (6) Progressive albitization of clastic plagioclase. The formation of zeolites and other hydrous Ca-Al-silicate minerals at the expense of calcic plagioclase are discussed separately. Replacement of srnectite upon burial Illite-smectite mixed-layering During burial diagenesis, smectite tends to disappear gradually with increasing depth through formation of irregular illite-smectite mixed-layers with a decreasing percentage of expandable layers. The replacement can be

L O G 6 A B A WELLS ( D O U A L A BASIN, CAMEROON 1 haolinite. dickitr

chlorite

-

!EIGHT LO55 t JOOO C' ( i n % i f initial weiaht I

I

20

10

CARBON RATIO" non-volatile carbon olal organic c a r b o 0.2 0.1, 0.6 0

0

D 0

\

\ i

I

i \

d

\

i

\

. a o

0

0

t

,000

.. .. .......,

u 40 30 20

6

-

30

20

1

0

0.2

0.4

0.6

Fig. 5-1. Mineralogical and geochemical analyses of the Upper Cretaceous sequence in the wells of Logbaba, Douala Basin, southwestern Cameroon. (After Dunoyer de Segonzac, 1969. fig. 10.)

312 monitored by the pro8ressive shift of the glycolated (001) peak, from 17 A in smectite towards 10A in illite, with the latter without expandable layers being the final product. Smectite or irregular illite-smectite mixed-layers do not occur in either the greenschist facies or, as discussed later, in the anchimetamorphic zone or equivalent zones of incipient metamorphism. Progressive increase in illite-smectite mixed-layering with depth was first demonstrated in wells in the Tertiary of the Gulf Coast (e.g., Weaver, 1959; Burst, 1959; Powers, 1959) and in the Carboniferous deposits of Oklahoma (Weaver, 1959). Subsequently this was also demonstrated in the Upper Cretaceous deposits of the Douala Basin, Cameroun (Dunoyer de Segonzac, 1964, 1969, see Fig.5-1); a well in the Tertiary and Cretaceous deposits of the Camargue, southern France (Dunoyer de Segonzac, 1969); wells in the Jurassic “Terres Noires” of southeastern France (Artru and Gauthier, 1968); the Cenozoic of Azerbaydzhan (Teodorovich et a]., 1967); the Jurassic deposits of eastern Ciscaucasia (Teodorovich and Konyukhov, 1970); the Pliocene and Quaternary deposits of the geothermal areas of the Salton Sea, California (Muffler and White, 1969); and the Taupo Volcanic Zone of New Zealand (Steiner, 1968; Eslinger and Savin, 1973a). On a regional scale, the progressive illitization of montomorillonite with increasing degree of incipient metamorphism has been described from the Jurassic “Terres Noires” of the subalpine chains of the westernmost French Alps (Dunoyer de Segonzac et al., 1966; Artru et al., 1969; Dunoyer de Segonzac, 1969). The absence of montmorillonite from deeply-buried sedimentary sections which grade laterally into less deeply buried sections which contain montmorillonite or illite-montmorillonite mixed-layers has been noted by Carrigy and Mellon (1964) in Alberta, by Karpova (1969) in the Greater Donbas area of the U.S.S.R., by Karpova and Timofeeva (1971) in the northern Caucasus, and others.

’

Depth and temperature of illite-smectite mixed-layering. Although the progressive mixed-layering appears to be largely a function of maximum depth of burial, the depth ranges over which the process takes place differ widely between various regions. This is true in the case of those sequences that underwent no significant postdepositional uplift and, thus, may be considered to be at their maximum depth of burial at present. Table 5-111 gives the depth ranges over which the proportion of montmorillonite layers in illite-montmorillonite mixed-layers progressively decreases from 75 to about 20% in a number of wells for which detailed

’

The name “montmorillonite” is used following the original articles, although in many of these papers no clear distinction is made between montmorillonite sensu strict0 and the smectite (montmorillonite group) in general.

TABLE 5-111 Depth and temperature intervals of reduction in the percentage of expandable layers in illite-montmorillonite mixed-layers from 75 to about 20% in six wells Locality or area and age of formation

Depth interval

Near Galveston, Texas (well E) Oligocene

Geothermal gradient

Reference

about 1300 m 35" (approx. 2100-3400 m depth) (85°-900 to 12Oo-125"C)

3I0C/km

Perry and Hower (1970, 1972)

Coastal Louisiana (well C) Miocene

>3OOO m 75 " (approx. 2600-5800 m depth) (80"-155"C)

24"C/km

Perry and Hower (1970, 1972)

Chambers County (Gulf Coast), Texas Tertiary

about 1200 m 35" (approx. 2600-3800 m depth) ( 100°-1350C)

29"C/km

Burst (1959; 1969, pp. 77-79)

32"C/km

Dunoyer de Segonzac (1964, 1969)

60°C/km

Muffler and White (1969)

180°C/km

Muffler and White (1969)

Logbaba, Douala Basin, Cameroun about 2100 m Upper Cretaceous (Approx. 1500-3600 m depth)

Temperature interval

'

50"

(70"-120°C

I)

to 20Oo-210"C ')

Wilson No. 1 well, Salton Sea geothermal field, southeastern California Plio-Pleis tocene

about 2000 m (approx. 1300 m 2-3300 m depth)

110°C (95"-100"

I.I.D. No. 1 well, Salton Sea geothermal field, southeastern California Plio-Pleistocene

about 500 m (approx. 45 m 2 - 5 0 0 m depth)

not known (up to 2OO0-21O0C ')

'

I 35% expandanble layers in illite-montmorillonite mixed-layers; * disappearance of discrete montmorillonite: completion of conversion of illite-montmorillonite mixed-layers to illite; mean geothermal gradient to 5000 f t depth. calculated from Muffler and White (1969. fig. 3).

w c. w

3 14 clay-mineralogic data and well temperatures (geothermal gradients) are available. These depth ranges vary from 500 to more than 3000m, with differences much larger than can be accounted for by differences in the geothermal gradients. The temperature intervals over which the progressive mixed-layering takes place range from 35" to more than 100°C. although Perry and Hower (1972, p. 2017) note a "much closer coincidence of the temperature intervals over which the dehydration takes place (in the two Gulf Coast wells E and C) than there is for the depth intervals". The temperature of the onset of illite-montmorillonite mixed layering at the expense of montmorillonite or of mixed-layers with minor proportions of illite layers varies within a relatively narrow range, commonly 70" to 115°C (cf., Dunoyer de Segonzac, 1970, fig.5; Burst, 1969, fig. 10). Both the temperature interval of progressive mixed-layering and the temperatures at which the mixed-layers retain only about 20% expandable layers show a far greater range of variations. These temperature differences could in part be related to differences in geothermal gradients. The higher water pressures prevailing at greater depths at which a given temperature is attained under a low geothermal gradient may be expected to necessitate higher dehydration temperatures. Such an explanation has been invoked to account for differences in dehydration temperatures in the Gulf Coast wells by Burst (1969) and might also explain the differences in temperature at which 20% montmorillonite content is reached in the mixed-layers of the four wells of the Gulf Coast and the Douala Basin listed in Table 5-111. Such a conclusion, however, is at variance with the high temperatures at which the conversion of illite-niontmorillonite mixed-layers is completed in areas of extremely high geothermal gradients: ( 1 ) 200°C for the Salton Sea geothermal field (Muffler and White, 1969), (2) above 230°C for the OhakiBroadlands, New Zealand, geothermal area (Eslinger and Savin, 1973a), and ( 3 ) about 230°C for the Wairakei thermal area of New Zealand (Steiner, 1968). The lower temperatures found for the advanced stages of mixed-layering in some of the Gulf Coast wells are in accordance with the experimental data of Khitarov and Pugin (1966), who found that the loss of interlayer water from montmorillonite was completed at temperatures of about 130°C at 1000 bars pressure (equivalent to 3.5 km depth at a geothermal gradient of 35"C/km) and 170°C at 5000 bars pressure (equivalent to 18.5 km depth at a geothermal gradient of about 10"C/km). In the presence of potassium, the dehydrated montmorillonite is converted into a 10-A "hydromica". G. Sabatier (unpublished. results, quoted by Dunoyer de Segonzac, 1969, pp. 71-72) found that in a mixture of montmorillonite and kaolinite in an artificial seawater with 0.60 g KCI/I, illitization of the montmorillonite took

315 place at a temperature of 200°C and 15 bars pressure. Experimental results on Louisiana Gulf Coast argillaceous sediment in artificial seawater at 100°C and 7 bars and 200°C and 21 bars by Hiltabrand et al. (1973). showed formation of chamosite or iron-rich chlorite and illite. whereas feldspar, kaolinite, and mixed-layered montmorillonite were destroyed or decreased in relative amount. Montmorillonite showed a shift of the basal peak to 14.7A, probably due to the formation of chlorite-like sheets. The effect of the concentration of cations, particularly of potassium. o n mixed-layering is discussed in a later section. Development of regular illite-smectite mixed-la.vers. In recent years. a tendency towards partial stacking order has been noted in many illite-sniectite mixed-layers from bentonites (e.g., McEwan, 1956: Maiklem and Campbell. 1968), presumably formed at the expense of montmorillonite, and subsequently also in other clastic rocks (see references in Hamilton, 1968. p. 13). The ordering commonly is apparent from the presence of sequences of non-integral(O0 1) spacings associated with a low-angle reflection (at 26-28 A in the air-dried or 29-32 A in the glycolated samples), which disappears upon heating above 4OO0C, with a concomitant reduction of the first basal spacing to 10.0-10.2A (see Fig. 5-2). In some deep wells, for instance those in the Miocene of Sicily and the Triassic of Tunisia studied by Long and Neglia (1968), the partial ordering of the mixed-layers characterize a deeper zone than does irregular illitemontmorillonite mixed-layering and, thus, obviously represents an advanced stage in the illitizaiion of montmorillonite through K-fixation upon deep burial. The partial ordering in these cells occurs in clays that still contain ample kaolinite, but no discrete montmorillonite. The tendency towards stacking order generally is manifested when the proportion of montmorillonite mixed-layers has been reduced to less than about 40% during progressive burial. Mixed-layer illite-montmorillonites with 40-30% expandable layers from the sandstones and coarse siltstones of the Permian succession of the Lower Hunter Valley, N.S.W., show extensive, non-integral 001-spacing sequences (see Fig. 5-3). Fourier transform investigation of these sequences by Hamilton (1968, pp. 12-13) indicated a structure of alternating mica and montmorillonite layers in approximately 2 : 1 proportions, in a layer sequence showing a striking tendency towards 1 : 1 stacking regularity, somewhat similar to that of mica-montmorillonite mixed-layers from Kinnekulle (Sweden) and Woodbury (England) analyzed by McEwan ( 1956). Some silty claystones and siltstones from the Tomago Coal Measures of this area contain 50 : 50 mica-montmorillonite mixed-layers showing peaks at -25A (air-dried) and 28A upon glycolation, with peaks at integral

3 16 Kinnekulle A - 2

1-L

0

I

,

I

I -

02 e

i

Two Medicine

13 1

660

3025

463

2662 2 4 7 0

4

L ,

.____-

/-A

0 b

“2 8

Fig. 5-2. Calculated profiles (upper curves) and X-ray diffraction patterns (lower curves) of ethylene-glycol-treated illite-montmorillonite with maximum 1 M ordering. a. Kinnekulle A-2, Sweden: 32% montmorillonite layers. b. Two Medicine Formation, Mont. ( U S A . ) : 35% montmorillonite layers. (After Reynolds and Hower, 1970, fig. 1 . )

submultiples of this value, indicating regular 1 : 1 stacking (Hamilton, 1968, p. 12; Fig. 5-3, upper XRD trace). The 50: 50 mica-montmorillonite mixedlayers found in a few claystones are less regular. The most common mixed-layer type in the Permian sediments of the Lower Hunt Valley has mica-montmorillonite layer ratios of 60 :40 to 70 : 30. Samples from claystone and clayey siltstone show basically irregular stacking, usually with a weak low-angle peak indicating a slight tendency towards stacking order, particularly in the more silt-rich claystones and

317

Y

?

m

AIRDRIED

-I I

Fig. 5-3. (001) diffraction spacing sequences of ordered and partly ordered illitemontmorillonite mixed-layers in X-ray diffraction patterns of air-dried and ethylene-glycolsolvated oriented-aggregate <2 p m fractions of Permian sediments of the Lower Hunter Valley, New South Wales. I(= kaolinite: M = mixed-layer illite-montmorillonite: Q = quartz. Upper traces: low-angle spacing with integral spacing sequences of regular 50: 50 illitemontmorillonite mixed-layer of a siltstone of the Tomago Coal Measures. Middle and lower traces: non-integral spacing sequences with weak (middle trace) or marked (lower trace) low-angle peak indicating different extent of tendency towards stacking order in respectively a siltstone to silty clay shale and a coarse siltstone or sandstone.

siltstones (Fig. 5-3, middle XRD trace). More striking evidence of stacking order is shown by the 70: 30 mixed-layers from the coarser siltstones and sandstones (Fig. 5-3, lower XRD trace). The tendency towards layer-stacking order thus seems to be related to the grain size of the sedimentary rock, being lowest in the finer claystones, and highest in the coarser siltstones and sandstones. This relationship is ascribed by Hamilton (1968, p. 21) to higher sediment permeability allowing a greater

318 degree of postdepositional alteration of a parental clastic mica-montmorillonite, mixed-layering of which is basically random. Perry and Hower (1970) have shown that in shales from wells E and B of the Gulf-Coast Tertiary deposits the illite-montmorillonite mixed-layers tend to develop an allevardite-like ordering of the interlayering upon the reduction in expandability to about 35%. The subsequent reduction in expandability to 20% montmorillonite layers takes place over a short depth interval, after which the latter phase persists for some 1500m to maximum well depth (see Fig. 5-4). The above-mentioned partially-ordered mixed-layers consistently show predominance of K over Na (e.g., Maiklem and Campbell, 1968, table 11; Hamilton, 1968, table 11; Perry and Hower, 1970, tables 3, 4, and 5 ) , and may be regarded as the K-analogs of the regular mixed-layer rectorite or allevardite, which is characterized by low K/Na ratios (cf. Henderson, 1970, p. 243). A similar K-rich mixed-layer mineral has been obtained in the laboratory by alteration of muscovite (Tomita and Sudo, 1968). A detailed study of diffraction profiles of mixed- layer illite montmorillonites from various localities in the northwestern United States and Sweden by Reynolds and Hower ( 1970) demonstrated that virtually all

Surface

b

a

o Well E

Well C

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-

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$10-

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I

20 30 40 50 60 70 Percent montmorillonite layers

1

80

*oL o * o

; Well C

o u 5

A percent rnontrnorillonite layers 2 5 0 ' d e p t h interval

Fig. 5-4. Relation between burial depth and (a) proportion of swelling layers in mixed-layer illite-montmorillonite and (b) decrease in percent swelling layers in mixed-layer illitemontmorillonites per 250-ft depth interval in two wells (C and E) in the Gulf Coast Tertiary deposits. In well E the geothermal gradient is almost 30% higher than in well C. (After Perry and Hower, 1972, figs. 4 and 6.)

319 illite-montmorillonites with more than 35-40% expandable layers are randomly interstratified, whereas those with less montmorillonite layers almost always have ordered interstratification. The most common ordered interstratification is of the “allevardite-like” type, which involves an IM superlattice, randomly interstratified with additional illite layers. I t is represented also at Kmnekulle (Sweden). An IMII-type superlattice, randomly interstratified with illite-layers, was found only in samples with 5 - 10% montmorillonite layers, particularly in K-bentonites. A tendency towards ordered interstratification has also been noted as a stage in the hydrothermal alteration of mica to montmorilloni te, in hydromica-montmorillonite mixed-layers with 20% swelling layers, the interstratification again becoming disordered at 40% expandable layers (Rateev and Gradusov, in: Rateev and Murav’ev, 1971). Some tendency towards development of regular mixed-layers has also been noted by Artru and Gauthier (1968) in the Jurassic sequence of the well Valvignieres in southeastern France, and by Dunoyer de Segonzac ( 1969, p. 69) at advanced stages of progressive mixed-layering *inthe Logbaba wells in the Douala Basin of Cameroon. From the acid pyroclastic rocks interbedded in the middle of the 4000 m thick Carboniferous coal measures of the Karaganda Basin, Kazakhstan, Kossovskaya et al. ( 1964, p. 525) have described ordered “allevardite-type” illite-montmorillonite mixed-layers, locally accompanied by kaolinite. They formed during the epigenetic hydromicatization of montmorillonite. The subsequent use of the designation rectorite for these mixed-layers by Shutov et al. (1969) has not been supported by the presence of high Na content. On the contrary, the high K contents indicate that they are ordered Khydromica-montmorillonite mixed-layers. These partially or completely ordered K-rich illite-montmorillonite mixed-layers discussed above appear to form at much less advanced stages of diagenesis than the regular Na-rich mixed-layer allevardite (or rectorite) associated with anchimetarnorphism and discussed in a later section. Rate of loss of expandable layers with depth; stability of illite-smectite mixrdlayers. Weaver (1967) was one of the first investigators to notice that the percentage of montmorillonite layers lost per depth interval is not constant over the whole range of mixed-layering in deep wells, and that mixed-layers tend to retain their last 20-10% expandable layers over a great depth range. Perry and Hower (1972, pp. 167- 170; 1972) found that in the Gulf Coast wells C and E the decrease in expandable layers content from 75 to 37% takes place comparatively and almost monotonically, e.g., in a 950 m depth interval from 2100 to 3050m in well E. On reaching the expandable layers content of 37%, the interstratification becomes ordered (see the preceding

320 section). Then there is a rapid further decrease to 20% over a brief depth interval, e.g., only 150 m in well E and less than 900m in well C (see Fig. 5-4), after which the montmorillonite/illite ratio remains constant with increasing depth (an additional 1350 m to the well bottom in well E). In the Chambers Co., Texas, Burst (1969, pp. 77-79) found a rapid decrease in content of expandable layers from 80 to 35% (“second dehydration stage”), within 1250m with no further reduction in expandable layers for an additional 650-m depth interval (“third dehydration stage”) (see Fig. 5-5). In a more detailed study of shales from another Gulf Coast well, Hower et al. (1976) showed a similar decrease in smectite content of the mixed-layers in the < 0.1 p m fraction from 80 to about 20% over a depth interval of 2000

p -

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f D e e p burial w a t e r loss

I

Fig. 5-5. Water escape curyes from montmorillonitic sediments during illitization upon burial. Dashed curve after Burst (1969), solid curve after Perry and Hower (1972). From Johns and Shimoyama (1972. fig. 3.)

32 1 to 3700 m, with “allevardite-ordered” mixed-layering being present below 3500m. They found that the proportion of smectite in the 0.2-2.0 p m fractions was consistently 5-20% higher than in the corresponding < 0.1 pm fractions (see Fig. 5-6). This seems to indicate that illite-smectite mixed-layers in the compositional range of 7 5 to about 30% smectite are relatively unstable. Mixed-layers with more than 75% or less than about 25% smectite appear to be relatively stable over considerable depth intervals at shallow and deep levels. respectively. Illite-montmorillonite mixed-layers containing 20% swelling layers has also been noted by Kossovskaya (1969, p. 344) to constitute “an exceptionally stable phase”, which persists throughout an interval of approximately 3000m in the cement of graywacke in the 5000-m thick Permian coal measures of the Petchora Basin, U.S.S.R. In the lower part of this interval, at about 3500 m depth, in the lean-coal zone, kaolinite disappears both from

1

oc

Fig. 5-6. Percentage of illite layers in illite-montmorillonite mixed-layers in the finest and coarsest clay fractions of the shales from a Gulf Coast well as a function of depth. (From Hower et al., 1976, fig. 3.)

322 the graywackes and the associated argillites (p.345). Below a depth of approximately 4000 m in the sequence, there is a gradual decrease in content of swelling layers to about 10% in the semi-anthracite zone, which may be considered to represent the beginning of “metagenesis” or initial metamorphism. Velde (1969) conducted on experimental study of the phase relations in the compositional join muscovite-pyrophyllite which contains the ideal K-montmorillonite composition at M U ~ ~ Pat~pH,O , ~ , pressures of 1000 to 3500 bars. As temperature increases to 300-3 1O”C, the compositional range of the mixed layers narrows down progressively to about Py,,~,,, the mixed layers usually being disordered or partially ordered (see Fig. 5-7). The mixed-layer mica-montmorillonite phase forming between 300” and 400°C is ordered with a convergence to the composition of Py,,, which corresponds to 30% montmorillonite interlayering. When the bulk composition of the system is more pyrophyllite-rich or more mica-rich than the compositional range of the mixed-layering, pyrophyllite (or kaolinite below 300-310°C) or mica appear as an additional phase. Velde stresses that the reaction temperatures and pressures found may be somewhat higher than those to be expected in natural clay-mineral suites because of the presence of magnesium and iron in natural material, among other reasons. He concluded that within the compositional range of mixedlayering (which is temperature dependent), the composition of natural mixed-layers is largely controlled by the bulk chemical composition of the system in which it is found. This appears to be confirmed by Bystrom (1956) who demonstrated that higher content of mica layers in the illite-montmorillonite mixed-layers in a number of K-bentonites from Sweden, whch formed under the same diagenetic conditions, is essentially controlled by higher K-content of the whole rock. Control of mixed-layer composition by whole-rock chemistry has an important bearing on the question whether the diagenetic alteration of montmorillonite to illite through mixed-layering is a mere one-way reaction, or whether the mixed-layers have a real thermodynamic stability. This question was raised by Zen (1963, 1967) and according to Perry and Hower (1970, p. 174) is “of prime importance to the petrologic interpretation of pelitic sediments”. The effective bulk composition of the system, however, should be considered to exclude unaltered clastic minerals. With decomposition of clastic minerals during diagenesis, the effective bulk composition may be modified gradually. A rapid release of potassium by decomposition of K-rich clastic minerals thus could not only prevent the appearance of an alkali-free phyllosilicate (kaolinite or pyrophyllite above and below about 300”C, respectively, according to Velde) in addition to the progressively less

323 1000 B A R S

Mica

MU

+ PY

I

I

I

I

I

l

l

I

12

25

37

50

60

69

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92

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+ Py

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Mont + K

+Q

I

I

I

I

I

l

l

I

12

25

37

50

60

69

78

92

PY

Fig. 5-7. Schematic phase relations of the compositional join muscovite-pyrophyllite at 1 and 2 kbar P ” , ~ The . arrows represent directions in which reactions were observed to proceed. The h e a v y black dashes at the pyrophyllite composition are the reaction intervals deduced from the experiments. The ideal K-montmorillonite composition corresponds to Mu ?,,Py,,,. M L = ordered mica-montmorillonite mixed-layer phase; K = kaolinite; Q = quartz. (From Velde, 1969, figs. 3 and 4.)

expandable mixed-layer phase, but would permit the latter to maintain its expandable-layer content below the maximum possible at that temperature in the case of K-poor compositions. In view of the experimental evidence for the existence of a two-phase field between mica and illite-montmorillonite mixed-layers, increasing in compositional extent with increasing temperature, rapid release of K might even lead to “stable” coexistence of an ordered mixed-layer and mica.

324 Repkucement of montmorillonite b y chlorite und occurrence of corrensire Considerable amounts of authigenic chlorite form during burial diagenesis at the expense of montmorillonite in a magnesium- or iron-rich environment. The mixed-layers that may appear as intermediate stages in this replacement (e.g.. see Koporulin, 1972), in several areas have been shown to be either swelling chlorite or the regular swelling mixed-layer corrensite, which may be regarded as chlorite-vermiculite, chlorite-“swelling” chlorite, or chloritemontmorillonite. Corrensite was first described by Lippman (1954) from the Upper Triassic deposits of southwestern Germany, in which it has since been found to be common. It is also common in anhydrite rocks and salt-clays of the Zechstein Formation of western Germany (e.g., see Fuchtbauer and Goldschmidt, 1959) and in the hypersaline facies elsewhere (see references in Millot, 1964; Kubler, 1973). I t is regarded as an early-diagenetic intermediate stage in the evolution of detrital illjte and mixed-layers to swelling chlorite and ultimat el y to chlorite. According to Lucas and Ataman (19681, corrensite, occurrence of which is widespread in the Triassic deposits of the French Jura, formed before appreciable burial. Dunoyer de Segonzac (1969, p. 228; 1970, p. 303), however, noted that these Triassic deposits were buried to a depth of 2000m, and that high concentration of magnesium cation in the interstitial solutions with concurrent increase in temperature would favor the formation of the corrensite. This may be true particularly in view of the success of Wyart and Sabatier ( 1966) in synthesizing corrensite from montmorillonite in Mg-rich solutions between 300°C and 600°C. I t appears, however, that if corrensite is not of sedimentary or early-diagenetic origin (as in evaporites), then it forms as a result of burial diagenesis (Kubler, 1973). The chlorite-vermiculite mixed-layers, found as an intermediate stage in the conversion of vermiculite to chlorite in the Logbaba wells of Carneroun (Dunoyer de Segonzac, 1969, fig. lo), are associated with montmorillonites down to a depth of 1500 m and with illite-montmorillonite mixed-layers with more than 30% expandable layers down to a depth of 2500 m. Eckhardt ( 1958) has reported occurrence of corrensite of high-grade diagenetic origin in appreciable quantities in Liassic clays associated with lean-coal rank carbonaceous matter in the cover of the “Bramscher Massif” near Osnabruck, Westphalia. Corrensite occurs below smectite-rich zones in association with laumontite and locally analcime in the tuffs of the Niigata oil field, northern Honshn, Japan (Iijima and Utada, 1971), and associated with laumontite in the volcanic Taveyanne graywackes of the cover of the Pelvoux Massif, French Alps (Kubler, 1970; Aprahamian, 1974; Kubler et al., 1974). In the latter

325 area, it appears to be absent from the higher-grade prehnite-pumpellyite zone. In the Wairakei geothermal area (New Zealand), long-spacing, regular. chlorite-montmorillonite mixed-layers occur below the occurrence of laumontite in the transition zone between high (10.5- 12.5 A) and low (9.9- 10.3 A) values for the do,, spacings of illite-montmorillonite mixedlayers (Steiner, 1968). Kubler (1973) has suggested that the occurrence of corrensite in magnesian carbonate rocks is also restricted to those that have undergone advanced burial diagenesis or even anchimetamorphism. Corrensite has been reported in the anchimetamorphic zone from slightly marmorized dolomitic limestones of the “chiiinons calcaires” of the northern Pyrenees (Kubler, 1967a, pp. 115-1 16). In general, however, it appears to be very rare in the anchizone. The occurrence of regular mixed-layers of dioctahedral chlorite and montmorillonite (tosudite) of epigenetic origin is discussed later in this chapter, because the dioctahedral chlorite appears to be mostly an alteration product of kaolinite. Influence of chemical environment on mixed-layering ’ The illitization of smectite involves not only dehydration, but also fixation of potassium between the layers, which implies an effect of the cation concentrations in the environment on the transition temperatures. Although progressive mixed-layering takes place upon deep burial both in the marine and the non-marine environment, the requisite potassium must be available (e.g., see Keller, 1970), either directly from pore solutions or released by the alteration of K-rich clastic minerals, such as micas or K-feldspar. Composition of interstitial solutions and permeability. Experimental results show the effect of composition of interstitial solutions on the illitization of smectite. Szadecky-Kardoss et al. (1968) converted illite into 12.5 A montmorillonite at temperatures of 200-400°C under water vapor pressures of 100-2000 bars. G. Sabatier (unpublished results, quoted by Dunoyer de Segonzac, 1969, pp. 71-72) found that whereas illitization of montmorillonite took place in a mixture of montmorillonite and kaolinite in artificial sea water with 0.60g KCI/I at a temperature of 200°C and p H L Oof 15 bars, montmorillonitization of kaolinite took place under the same conditions in distilled water. Powers (1959) has proposed that at burial below a certain depth, called the “Mg2+-K+ equivalence level”, K + is adsorbed onto clays in preference to M g 2 + . The depth of this equivalence level is believed not to exceed several hundred feet. Control of progressive mixed-layering by the increase in the potassium content of interstitial solutions with depth has also been found by Long and Neglia (1968) in a chemical study of clays of deep wells from Sicily and Tunisia.

326 Progressive mixed layering appears to be favored by high permeability of sediments and rocks, which presumably facilitates the access of K-bearing solutions. Vlasov (1966) observed that in epigenetically-altered clastic rocks of the same well-core intervals the percentage of expandable layers in illitemontmorillonite mixed-layers ranged from 30% in marl, through 10-20% in mudstone, to nil in siltstone, which reflects the dependence of the extent of epigenetic illitization on the permeability. The content of Fe-chlorite and 1M hydromica, both considered to be authigenic alteration products of biotite, was highest in the siltstone. Koporulin ( 1972) noted that the chloritization of montmorillonite in coarse-grained sandstones of Kamchatka is virtually complete, whereas the associated medium-grained sandstone and siltstone contain chloritemontmorillonite mixed-layers and un-chloritized montmorillonite with minor amounts of kaolinite, respectively. Differences in the extent of crystallization of the mica-montmorillonites of the Permian succession of the Hunter Valley, New South Wales-evidenced by the earlier mentioned tendency for better defined, ordered layer stacking in sandstones and coarse siltstones than in the fine-grained siltstones and silt-rich claystones, and the virtual absence of such layer stacking in the finer claystones (Hamilton, 1968, pp. 12- 13)-could also reflect degree of availability of potassium. Effect of clastic constituents and bulk composition on degree of mixed layering. In many cases, the increase in potassium content with depth and in potassium availability with increasing permeability, reflects the progressive decomposition of clastic K-rich minerals. Alteration of these minerals that may take place during burial diagenesis includes mixed-layering of muscovite, chloritization or vermiculitization of biotite. and decomposition of K-feldspar. The content and rate of alteration of such K-rich clastic minerals are likely to exert a control on the availability of potassium and, thereby, to affect the rate and range of temperature of mixed-layering at the expense of smectite. Teodorovich and Konyukhov ( 1970) have discussed the relationship between the lithology of sedimentary rocks and the depth at which mixedlayering takes place at the expense of smectite in various areas. In formations having polymineralic clastic rocks rich in unstable clastic constituents, such as feldspars and micas (particularly biotite), this depth is relatively shallow, e.g., in the Jurassic deposits of Ciscaucasia and western Siberia. Mixed-layering here takes place during (early) diagenesis and early catagenesis. and is completed at the stage of moderate catagenesis (see also Karpova, 1969, p. 16). This is also the case in the Paleozoic series described by

327 Gavrilov and Aleksandrova (1968) from the southern Urals, and in the Neocomian deposits of the western Siberian Platform (Sakhibgareev and Galikeev, 1971). In sequences of mature, quartz-rich sedimentary rocks poor in such unstable constituents, e.g., the Mississippian and Miocene formations of the North-American platform and the Cenozoic deposits of Azerbaydzhan, mixed-layering takes place at much greater depth in the stage of late catagenesis. Transformation is completed only during the stage of initial metamorphism (or metagenesis). Logvinenko ( 1964) has even suggested that mixed-layers not only may persist in the zone of metagenesis and the greenschist facies, but may even originate there. These differences were considered by Teodorovich and Konyukhov (1970) to be controlled mainly by the availability of potassium supplied by the alteration of clastic micas and the albitization of K-feldspar. The percentage of expandable layers in illite-montmorillonite mixed-layers, therefore, can be used as an indicator for the degree of epigenesis only in sedimentary rocks of similar lithology (e.g., see Sakhibgareev and Galikeev, 1971). The dependence of the extent of mixed-layering on the sedimentary rock type is evidenced particularly by the difference between the percentages of expandable layers of three-layer clays in ( 1) bentonites or K-bentonites derived from vitric-lithic tuffs or ash, and (2) the associated clastic rocks. For instance, in the Permian succession of the Lower Hunter Valley, N.S.W., the bentonitic clays consist of an almost fully expandable beidellite, whereas the associated clastic sedimentary rocks contain mixed-layer illitemontmorillonites with 40 to 30% expandable layers (Hamilton, 1968). The lower Paleozoic K-bentonites or “metabentonites” known from various parts of the eastern and central United States consist of irregular illite-montmorillonite mixed-layers with about 20-30% expandable layers (Lounsbury and Melhorn, 1964; Mossler and Hayes, 1966; Beall and Ojakangas, 1967). The clastic shales associated with these K-bentonites contain illites that are virtually devoid of expandable layers (Mossler and Hayes, 1966).

Direct evidence for decomposition of clastic K-rich minerals during the mixedlayering process. The potassium required for the occurrence mixed-layering may in part be provided by the late-diagenetic decomposition or illitization of clastic muscovite. This has been shown to take place in the same depth range as the progressive mixed-layering of montmorillonite, based on data obtained from deep wells (Dunoyer de Segonzac, 1964; Perry and Hower, 1970). Weaver and Wampler (1970) found a decrease of 40 million years (from 226 to 186 m.y.) in the apparent age (K-Ar) with depth in the ( 2 pm fractions, and of about 100 million years (from 325 to 231 m.y.) in the bulk

328 samples from a 3700-m depth interval of the Mississippi Delta, in which the percentage of montmorillonite layers in mixed-layers decreased from 80 to 20. In contrast, the apparent ages of the > 2 p m fractions remained relatively constant at about 400 million years, with a considerable decrease in the abundance of K-feldspar in this fraction with depth. This demonstrates that the potassium cations which were fixed between the montmorillonite layers in progressive mixed-layering process, have been released from the coarse K-feldspar, and possibly from mica, with increasing depth. Hower et al. (1976) found a similar decrease in percentage of smectite layers in mixed-layers in the < 0.1 p m fractions of shales from a depth interval of 1700m in a Gulf-Coast well, that was accompanied by the decrease of the clastic K-feldspar (but not the albite) content of the shales to almost zero. In a further K-Ar study of a 4000-m section of Miocene Gulf Coast sediments, Perry (1974) found that the decrease in apparent whole-rock ages from 230 to 137 m.y. does not reflect merely the age decrease in the < 0.5 p m fraction (from 164 to 100 m.y.), accompanying progressivk illitization. The apparent ages of the coarser fractions also significantly decrease (e.g., from 372 to 158 m.y. in the > 10 prn fractions). This is due to the greater loss of radiogenic argon relative to potassium upon destruction of detrital illites and K-feldspars, and not to simple degassing. Aronson and Hower (1976) have shown that the < 0.1 p m fraction, which is composed nearly of pure illite-smectite, actually gains radiogenic argon with depth. The K-Ar studies thus confirm that large portion of the shale reflects homogeneization of an initially heterogeneous mixture of detrital phases. The content and rate of decomposition of K-rich minerals during diagenesis thus exert a considerable effect on the rate of reduction in the percentage of expandable layers in mixed-layers with progressive burial, through their control of the effective bulk composition of the system. Mixed-layering in lithologically different deposits. Only when comparing lithologically similar rocks can one expect to find a regular decrease in the percentage of swelling layers upon deep burial. If successive rock units are different, irregular decrease or even inversions may occur. For instance, in the Volga-Urals region, U.S.S.R., Vlodarskaya and Nosov ( 1973) reported higher percentages of swelling layers in the illite-montmorillonite mixedlayers of the Riphean deposits in the deepest zone of epigenesis at a depth of about 3000 m, than in the overlying Middle Devonian and Carboniferous deposits. Polvtype and composition of illite during burial diagenesis

Dioctahedral micas and illites appear in nature in different stacking polymorphs or polytypes, which can be distinguished using the spacings of

329 (hkl) X-ray diffraction peaks in the 20-36°C (Ca K, radiation) region (Yoder and Eugster, 1955; Velde and Hower, 1963; Maxwell and Hower, 1967). The ordered stacking arrangements most common in nature are the monoclinic polytypes 1M and 2M, I . The trigonal polytype 3T and the second two-layer arrangement 2M, are somewhat less common (in the polytype nomenclature, the first symbol represents the number of mica layers per unit cell, whereas the second one gives the symmetry). The lM, form is a disordered one with random layer-stacking (Yoder and Eugster, 1955). The 2M, muscovite polytype, which is the predominant form in igneous and metamorphic rocks, including the greenschist facies, and is a detrital constituent of sediments, has been shown experimentally by Velde (1965a) to be the only stable form. The reaction l M d 1M 4 2M, takes place with increasing time, temperature, or pressure. Loss of potassium from 2M muscovites of igneous or metamorphic origin during weathering and sedimentation may produce degenerated (“stripped”) illites, which commonly have the disordered lM, polytype form. These tend to revert to illite due to the K-fixation in the marine environment during burial or, possibly, on the sea floor (Weaver, 1958, 1959). Weaver (1958, pp. 841-2) showed experimentally that such weathered illite regained the 2M polytype form upon regeneration as a result of saturation with potassium. Although most illites (excluding glauconites and illite-montmorillonite mixed-layers) examined by Weaver (1958; 1959, pp. 164, 170) were of the 2M polymorph type, Velde and Hower (1963) and Hower and Mowatt (1966) showed that disordered form predominates in the < 1 pm size fraction of most non-metamorphic Paleozoic sedimentary rocks. Assuming that 1 M, material derived from a high-temperature progenitor by degradation should reconvert to 2M illite by K-fixation, they concluded that this predominance of 1M, material in the fine-grained fraction indicates a lowtemperature origin. Most authigenic low-temperature illites, including those produced by alteration of clastic K-feldspars, appear to have formed as the l M d form (cf., Weaver, 1967). In the Precambrian Belt Series of Idaho, Maxwell and Hower ( 1967) found an increase with stratigraphic depth from 20-40% to 80-100% in the proportion of the 2M polymorph in the illite of the < 2 p m fraction of argillites from a 11,500-m thick section; the 1M form was not found. Oxygen-isotope thermometry on coexisting quartz and illite from the Belt Series (Eslinger and Savin, 1973b, p. 2557) indicates that 50% of the illite of the (0.5 p m size fraction would have been converted from the lM, to the --f

’

The 2M, polytype (Smith and Yoder, 1956) was referred to by earlier authors as 2M polytype.

330 2M polytype at a temperature of about 290°C. The advanced grade of diagenesis in this section is also evidenced by the absence of illitemontmorillonite mixed-layers and of kaolinite. The transformation from the lM, to the 2M polytype is paralleled by textural reconstitution of the argillites to phyllite and is completed only upon the appearance of biotite. A good inverse correlation was also found to exist between the percentage of 2M polymorph in illites and the IX0/"O ratio of carbonates from the Ordovician limestones in west-central Vermont, which have undergone various degrees of lowest-grade metamorphism (Gavish and Reynolds, 1970). In the progressive burial diagenesis and anchimetamorphism of the Jurassic black marly shales ("Terres Noires") of the subalpine zone of the French Alps (see Fig. 5-17), the 2M polytype replaces the lM, polytype at a very advanced stage of burial diagenesis, corresponding approximately with the disappearance of kaolinite and irregular illite-montmorillonite mixed-layers (Artru et al., 1969; Dunoyer de Segonzac, 1969, pp. 157-160). The 1M polytype, though shown by Velde (1965a) not to be the stable form below 200-350°C for the muscovite composition, as eaflier believed by Yoder and Eugster (1955), may appear as a metastable intermediate form in the formation of 2M muscovite at the expense of kaolinite (see Fig. 5-8). Many 1M illites found in nature can be interpreted either as such alteration products of kaolinite (e.g., Kisch, 1966a; Muller, 1967a) or of feldspars through an intermediate kaolinite stage (cf. examples in Velde, 1965a, p. 447; and Triplehorn, 1967), in part formed during deep-epigenetic alteration.

:::pMd, ;- 1 Ix

I

1M

0

100

Kaol inite 5

10

15

Days

Fig. 5-8. Time-temperature relations of the stages in the reaction: kaolinite+ KOH 2M,muscovite. Runs were made at 2 kbar of water pressure; the conversions were not reversed. Crosses= 1M polymorph present; squares=2M, polymorph present; triangles= 1 M, polymorph present; circle= kaolinite only, no mica present. (From Velde. 1965a. fig. 1.)

33 1 Chemical composition and polytype of authigenic dioctahedral hydromicus Most of the 1M dioctahedral hydromicas found in nature have a chemical composition which differs from that of muscovite, e.g., glauconite or celadonite (cf. Weaver, 1958; Keller, 1963; Wise and Eugster, 1964; Kossovskaya and Drits, 1970). The 1M polytype also predominates in montmorillonites and, as pointed out by Weaver (1959, pp. 164-165, 170, 18l), this polytype tends to be retained in illite-montmorillonite mixedlayers, even when the illite-layer contents attain 90%. For instance, the 1M to 1M, polytype characterizes the mixed-layers of K-bentonites or "metabentonites" mentioned earlier, whereas the associated marine limestones and shales contain 2M, micas, presumably of detrital origin (Mossler and Hayes, 1966). Only in the course of deep burial and incipient metamorphism do dioctahedral 1M mica-type clay minerals, such as glauconite and celadonite, and illite-smectite mixed-layers, transform into 2M, hydromicas.

1

Minerals

Mixed-layer phases

Kaolinite

lontmorillonite

Hydromica

Chlorite

I H = 8-8.5

pH = 9-9.5

pH = 0-9

Stages

c

initial

deep

I Metamorphism

Medium index

1 pH = 8-9.5

PH = 5-6

Fig. 5-9. Schematic representation of the changes in the mineralogy of sandstone cements with progressive postsedimentary alteration in the Carboniferous deposits of the DnieprDonetz depression and the Donbas area, U.S.S.R. NB: catagenesis =epigenesis. Symbols for columns: H - M = hydromica-montmorillonite mixed-layer; Ch - M =chlorite-montmorillonite mixed-layer; I =replacement; l a =replacement of mica; l h =replacement of volcanic fragments; / I =cements (matrix). Correction: in deep epigenesis, the predominant chlorite polytype is monoclinic I I h ( p = 97") and not Zb ( /3 = 97") as shown. (After Karpova, 1969, fig. 9.)

332 Karpova (1966, 1967, 1969; also Logvinenko and Karpova, 1968) have correlated a progressive change in both polytype and composition of the authigenic dioctahedral hydromicas in the clastic Carboniferous rocks of the Donbas (U.S.S.R.)with the intensity of burial-diagenetic alteration. There is a change from the Fe-rich, glauconite lM, or 1M hydromicas in shales and hydromica with lower Fe content in sandstone cements in the “stage of initial catagenesis” ( = epigenesis) in the Dniepr-Donetz depression, through more Al-rich and Fe-poor 1M hydromicas during the “stage of deep epigenesis”, to predominance of 2M, hydromica in the “stage of early metagenesis” in the central and southeastern areas of Donbas (see Fig. 5-9). Kossovskaya and Drits (1970), in a review of the chemistry of dioctahedral micaceous minerals in sedimentary rocks, have noted a concurrent increase in the tetrahedral charge from 0.28-0.60 to 0.65-0.9 and decrease in the octahedral charge from 0.35-0.55 to 0.1-0.4 in illites from environments of “deep epigenesis and metagenesis” as compared to those unaffected by “epigenesis”. Although the epigenetic illites still contain appreciable contents

Percent expandable layers

0

Celadonite

+Inc

oct chg,

Hower and Mowatt (1966, table 4 1

Pyrophylli t e

K1+{ Fe3+(Fe2’; Mg2+)[Si,0,,/(OH)2])”

Fig. 5- 10. Relationship between percentage expandable (montmorillonite) layers-percentages are shown for each point-and tetrahedral-octahedral structural charge in illites and illitemontmorillonite mixed-layers, plotted on the composition triangle pyrophyllite-muscoviteceladonite. (After Hower and Mowatt, 1966, fig. 1 1 . Note: in the original paper the caption of this Figure has been interchanged with that of fig. 12.)

333 of bivalent Fe2+ and Mg’+ in the octahedral layers, these changes do indicate the evolution of the illite towards the muscovite composition prevailing in metamorphic rocks. Similarly, Hower and Mowatt (1966) found that twenty-one illite-montmorillonite and illite samples all contained close to 100% of the lMd polytype in the (0.5pm size fractions that exclude detrital constituents, but that the illites have much lower tetrahedral charges (0.26 to 0.59) than sericite, and that their K and A1 contents do not approach the muscovite vertex of the pyrophyllite-muscovite-celadonite composition triangle (see Fig. 5-10). They concluded that illites differ from the degraded hgh-temperature dioctahedral micas, and remain distinct from the true micas until the beginning of metamorphism. Relationship between host-rock lithology and polytype of authigenic dioctahedral hydromicas During epigenesis marked variations still exist in the polytype of authigenic illites present in different lithologies. In part at least, these variations reflect the more advanced burial-diagenetic alteration in more porous rocks. Vlasov (1966) found, for example, that Devonian siltstones from the eastern part of the Russian Platform contain authigenic 1M hydromica, whereas the associated, less permeable mudstones contain 1M, mixed-layers with 10-2053 swelling layers. The clastic hydromicas are of the 2M, polytype in both cases. During the “stages of initial and deep catagenesis” ( = epigenesis) of the Dniepr-Donetz depression and the marginal areas of the Donbas, the 1M hydromica polytype predominates over the 2M, polytype in the authigenic cements, whereas 2M polytype predominates in less altered clastic clays. Both types of rocks contain montmorillonite showing progressive mixedlayering (Karpova, 1969, pp. 14-16; see Fig. 5-9). A convergence of the hydromica polytypes in the argillaceous and sandy rocks takes place in the central and southeastern areas of the Donbas; the latter area contains anthracite and has reached the “stage of early metagenesis”. Even cements in this zone are characterized by the predominance of the 2M, over the 1M polytype hydromica (Karpova, 1969, pp. 10- 11). In summary, it appears, therefore, that most authigenic dioctahedral three-layer silicates, including “illite”, formed during late diagenesis, have chemical compositions and polytypes different from those of muscovite. The 2M polytype becomes predominant only on approaching the “stage of incipient metamorphism”, as the composition approaches that of muscovite. The presence of major amounts of the lMd and 1M polytypes constitutes an indication that this stage has not yet been attained.

,

334

Alteration and modification of kaolinite during burial diagenesis Kaolinite does not form mixed layers and tends to persist to somewhat greater depths than does smectite. Its disappearance with depth is commonly due to replacement by illite and/or chlorite, provided that sufficient amounts of potassium, iron or magnesium are available. This transformation has been documented in a large number of deep wells, among others in the Cretaceous sequence of the Douala Basin of Cameroun (Dunoyer de Segonzac, 1964, 1969), the Jurassic deposits of the southwestern French Alps (Dunoyer de Segonzac, 1969, pp. 164- 174), the Permo-Triassic sequence of the Paris Basin (Dunoyer de Segonzac, 1969, pp. 147-152), the Mesozoic deposits of western Germany (Fuchtbauer and Goldschmidt, 1963), and the Upper Carboniferous deposits of northwestern Germany (Scherp, 1963; Stadler, 1963; Esch, 1966). In the absence of potassium, magnesium or iron cations, kaolinite tends to persist into much more advanced stages of diagenesis.

Lithological and compositional (interstitial fluids) control of illitization and chloritization of kaolinite Low-temperature experimental studies have shown that at a suitable range of temperatures, illite and chlorite are stable with respect to kaolinite at high K + / H and high (Mg2+ F e Z f ) / H + concentration ratios, respectively (cf. Garrels and Christ, 1965, pp. 363-370). As mentioned earlier with respect to the alteration of smectite, K f , M g 2 + ,and Fe2+ cations must be supplied either by the decomposition of clastic minerals (Shutov and Dolmatova, 1961; Kossovskaya et al., 1963; Kossovskaya et al., 1965), or by interstitial solutions (e.g., Lisitsyn et al., 1969). The replacement of kaolinite by illite and chlorite is obviously favored by high porosity and permeability and, thus, is further advanced and/or completed earlier in the sandy than in the associated argillaceous rocks (e.g., Scherp, 1963; Stadler, 1963; Karpova and Shevyakova, 1965; Long and Neglia, 1968; Triplehorn, 1970). Moreover, it may be arrested by oil migration and consequent saturation of rocks with oil (Kulbicki and Millot, 1963; Fuchtbauer and Goldschmidt, 1963; Prozorovich, 1967). In the presence of required cations, the replacement of kaolinite by illite and/or chlorite is commonly concluded at comparatively shallow depths of burial. For instance, in the predominantly sandy and shaly rocks of the Logbaba wells in Cameroun and the well Contres in the Paris Basin (Dunoyer de Segonzac, 1964; 1969, pp. 57-81 and 147-151) kaolinite disappears at a temperature of 80-90°C at depths of 2000 and 1500m, respectively, long before the disappearance of illite-montmorilloni te mixedlayers (see Fig. 5-1). In three wells in the predominantly marly Jurassic “Terres Noires” deposits of the southwestern French Alps studied by +

+

335 Dunoyer de Segonzac (1969, pp. 164-174), kaolinite persists to greater depths. It disappears only 500-1000m above the zone of complete illitization of montmorillonite, but still long before the appearance of anchizonal illite “crystallinities”. Temperature gradients are not available for these wells. The persistence of both kaolinite and illite-montmorillonite to a depth of 5100m (well bottom) and temperature of 180°C and at least the lowvolatile bituminous coal rank ( Ro,, 2 1.8) zone in the predominant marls and marly limestones of the well Pierrefeu (Camargue) in southern France must be ascribed to the lack of circulation of interstitial solutions due to the extremely low porosity and permeability (Dunoyer de Segonzac, 1969, pp. 135-146). In the Carboniferous deposits of the axial part of the DnieprDonetz syncline, kaolinite and mixed-layer clays appear to persist throughout the “stage of deep epigenesis” to a depth of 5000 m in association with lean coals (Karpova and Shevyakova, 1965; Karpova et al., 1969). The fact that cation availability controls the stage of burial at which the replacement of kaolinite is completed has been shown by Kisch (1969, 1974) to be also reflected in the marked lack of correlation with associated coal ranks in different areas. This rank ranges from approximately 14-15% V.M. ( = volatile matter yield; dry, ash-free basis) in the Permian deposits of the Bowen Basin, Queensland, through less than 9% in the afore-mentioned wells in the Jurassic deposits of the southwestern French Alps and in the Carboniferous deposits of the central Donbas, to approximately 6-7% in the Upper Carboniferous rocks of the Munsterland 1 well, Westphalia, West Germany. The rank thus ranges from lean coal through semi-anthracite to anthracite. In some areas, e.g., in the Tunguska area of eastern Siberia (Gurewitsch and Toporez, 1968), the replacement appears to be completed in association with even lower rank coking coals. Due to the chemical factors, affecting the illitization of kaolinite in addition to depth and geothermal gradient, this replacement cannot be used as an accurate indicator of stage of burial (Dunoyer de Segonzac, 1969, p. 21 1). By the beginning of initial metamorphism, however, all kaolinite disappears. Kaolinite which has not been replaced by illite and chlorite by that stage is replaced by minerals such as pyrophyllite or paragonite. These replacements are considered in a later section. Changes which affect kaolinite during advanced stages of diagenesis in the absence of cations required for illitization or chloritization are discussed next. Change of kaolinite polytypes during burial diagenesis The kaolinite in ( 1) kaolinite-rich argillites ( e g , kaolinite-coal tonsteins), (2) monomineralic quartz-sandstones with kaolinite matrix, and (3) other clastic rocks poor in K-rich and Mg-Fe-rich detrital minerals, may persist

336 into advanced stages of diagenesis in the absence of the cations required for its alteration to illite and/or chlorite. Such kaolinite tends to acquire an increase in degree of stacking order with advancing degree of burial metamorphism. Eckhardt and Von Gaertner (1962) and Eckhardt (1965) studied the kaolinites from kaolinite-coal tonsteins in Westphalia, West Germany, using the degree of coalification of the associated coal seams as a measure of diagenetic temperature. A distinct gradual increase of ordering from disordered (“fireclay mineral”) to wellcrystallized kaolinite was found to exist. On the other hand, Stadler (1971a) noted that kaolinite in kaolinite-coal tonsteins from Westphalian anthracites at Ibbenburen, Westphalia, showed no noticeable improvement (increase in degree of ordering) compared to those in lower-rank, bituminous coals of the same age in the Ruhr area. Kaolinite may persist throughout the stage of “deep epigenesis” (Karpova and Shevyakova, 1965; Karpova and Timofeeva, 1971). This is always true in the case of kaolinite-rich clays which are common both within the coal seams and in the underclays (e.g., Kossovskaya et al., 1965). During deep epigenesis, however, the kaolinite may also be replaced by its monoclinic 2M, polytype, dickite. This replacement is common in monomineralic quartz sandstones with kaolinite matrix (Shutov and Dolmatova, 1961). Kossovskaya and Shutov ( 1963) consider the quartz-dickite “facies” characteristic for such rocks and for kaolinite clays in the “stage of deep epigenesis”, in contrast to the poorly-sorted rocks rich in primary clay minerals, in which hydromicatization is favored. The transformation of kaolinite to dickite appears to proceed gradually through kaolinite-dickite “packets” or mixed-layers, which inherit the structure of kaolinite twins. The proportion of dickite crystals increases towards the base of the deep epigenesis zone. As a result of intense stress, nacrite, the 2M, polytype of kaolinite, may form at the expense of vein dickite in the “stage of metagenesis” (Shutov et al., 1970). In most cases, kaolinite and its polytypes are replaced during initial metamorphism by pyrophyllite, as discussed in a later section. Locally, however, dickite appears to persist or possibly even to form during incipient metamorphism. In the “Schistes de Steige” in the Vosges, France, Clauer and Lucas (1970) have reported dickite, which is associated with illites and shows crystallinities characteristic of the low-grade part of the anchizone. Replacement by pyrophyllite occurs only in the high-grade part of the anchizone. Relationship between permeability and three-dimensional order in kaolinite and dickite. High permeability appears to favor both the increase of order in kaolinite sensu strict0 and the transformation of kaolinite into its polytype dickite.

337 In the Carboniferous deposits of the Dniepr-Donetz depression, kaolinites from sandstone cements are always ordered along the b-axis, whereas those from argillaceous rocks show a low degree of order (Karpova, 1969). The same relationship has been noted by Fiichtbauer and Goldschmidt (1963), who also observed that the increase in kaolinite crystallinity in sandstones was inhibited by oil migration and saturation with oil. The replacement of kaolinite by dickite similarly appears to be favored by high porosity and permeability. Consequently, dickite occurs in relatively porous sandstones of Hassi Messaoud (Sahara), whereas the accompanying finer-grained sandstones and siltstones still contain kaolinite proper (Ferrero and Kubler, 1964; Cassan and Lucas, 1966; Dunoyer de Segonzac, 1969, pp. 69-70). With increasing depth, replacement of kaolinite by dickite has been observed by Kossovskaya and Shutov (1963) to begin in the cement of sandstones. Only deeper in the section does dickitization affect the kaolinite of argillaceous rocks. This relationship with porosity and permeability apparently also holds true for some dickites of hydrothermal origin. For example, in southeastern Kansas, dickite, which is probably related to groundwaters heated by igneous activity, occurs on a regional scale in cavities and interstices in limestones, whereas kaolinite persists in less porous rocks of the same area (Hayes, 1967; Schroeder and Hayes, 1968). The inhibition of the kaolinite-dickite transformation by low permeability could be the main cause of the persistence of kaolinite into advanced stages of diagenesis and even into incipient metamorphism, as noted in the preceding section. Occurrence of dioctahedral, aluminous chlorites In addition to the burial-diagenetic changes noted above, kaolinite has locally been noted to undergo replacement by sudoite and tosudite. Sudoite is a dioctahedral, aluminous chlorite first identified in 1961 (G. Muller, in Von Engelhardt et al., 1962; see also references in Muller, 1963, 1967a,b). Tosudite (Frank-Kamenetskiy et al., 1965) is a regular 1: 1 sudoitemontmorillonite mixed-layer, i.e., the dioctahedral equivalent of corrensite. Sudoite and tosudite, which seem to be mainly of hydrothermal origin, occur with ( 1) hydromuscovite in quartz-porphyry tuffs in the Black Forest, southwestern Germany (Muller, 1963), (2) Na-rectorite (allevardite) in veins in hydrothermally-altered schists in Dagestan (U.S.S.R.)and Savoie (France), and (3) regular K-mica-montmorillonite mixed-layers (“K-rectorite”) and dickite in Tadzhikistan (U.S.S.R.), Japan, and the Crimea (U.S.S.R.) (see Gradusov, 1971 , and the literature quoted therein). Dioctahedral chlorite and its mixed-layers, however, also occur locally as products of late diagenesis. Together with 1M muscovite, sudoite, which occurs in Permian dune-

338 sandstones in Hesse (West Germany), probably developed from kaolinite or feldspar during deep burial in a Mg- and Fe-poor environment (Muller, 1967a). Authigenic sudoite and irregular sudoite-illite-montmorillonite mixedlayers occur with local kaolinite in Middle-Triassic marls and sandstones in Wurttemberg, West Germany (Von Engelhardt et a]., 1962; Muller, 1963). In the equivalent sandstones in three wells in southern Bavaria the above minerals, accompanied by tosudoite, occur at present depths of 1400-2300 m (Kulke, 1969). In the above occurrences there is an inverse relationship between the amounts of sudite and of kaolinite present. Sudoite is considered by Miiller (1963, p. 125) and Kulke (1969) to have formed at the expense of kaolinite upon deep burial during the late Tertiary time. This replacement is most marked in the most permeable sandstones, kaolinite partly being preserved to greater burial depth in the denser rocks. Drits and Shutov (1966) reported the occurrence of tosudite with dickite or nacrite in veins formed as a result of deep epigenesis in the lower part of the coal measures of the Karaganda Basin of Kazakhstan, U.S.S.R. Regular K-rich illite-montmorillonite mixed-layers were subsequently described by Shutov et al. (1969) from the acid pyroclastic rocks of these coal measures. Although the occurrences mentioned above have formed during late diagenesis, the local association of hydrothermal sudoite and tosudite with minerals such as nacrite and Na-rectorite (but not, as far as reported, with pyrophyllite) may indicate that the dioctahedral chlorites could persist into incipient metamorphism or metagenesis. Change in composition and polytype of trioctahedral chlorite During diagenesis and initial metamorphism, the chlorite content of sediments steadily increases at the expense of primary clay minerals such as kaolinite, smectite, and vermiculite. The chlorites forming during early diagenesis are generally rich in Fe, the Mg contents increasing with increasing depth during epigenesis or advanced diagenesis (Eckhardt, 1958; Muravjew and Salyn, 1969). This change in composition continues during initial metamorphism. In southern New Zealand, Carman ( 1965) has found that the chlorite in the prehnite-pumpellyite facies is predominantly optically-negative diabantite and brunsvigite (Ferich), whereas in the greenschist facies it is optically positive pycnochlorite. Bishop (1972b, p. 3184) noted that the chlorites of metagraywackes in the prehnite-pumpellyite-metagraywacke facies tend to be optically negative with normal interference colors (i.e., relatively high birefringence) and are Fe-rich, whereas those in the pumpellyite-actinolite-schist facies and greenschist facies are optically negative Fe-Mg-chlorites with anomalous

339 blue interference colors. Optically positive Mg-Fe-chlorites are restricted to metavolcanic rocks and marbles. Late diagenesis and incipient metamorphism also cause a change in the polytype of chlorites. Of the six semi-random one-layer stacking polytypes possible in chlorite, Bailey and Brown (1962) have recognized four in a study of polytypes of 303 natural chlorites. The most common polytype, IIb ( p = 97"), occurs in metamorphic and igneous rocks and higher-temperature ore deposits, whereas the remaining three polytypes Ia, Ib ( p = 97") and Ib ( p = 90') occur mostly in lower-temperature environments. The IIb chlorites may weather to Ia chlorite, or through mixed-layer chlorite-vermiculite to la vermiculite. Hayes (1970) has found that the authigenic chlorites in sedimentary rocks almost exclusively belong to type I polytype. The IIb polytypes, which occur in many unmetamorphosed sedimentary rocks, are unweathered detrital chlorites. Hayes proposed a diagenetic crystallization and stability sequence from the disordered-stacking type Ib, through Ib ( p = 9 7 " ) to the Ib ( p = 90°) polytype. The Ib, and Ib ( p = 97") polytypek are absent from all samples obtained from depths of more than a few thousand feet. Low-grade metamorphic conditions, with inferred temperatures of 15O-20O0C, are usually necessary to cause conversion of the Ib ( p = 90") polytype to the more stable and most common IIb ( p = 97') structure. Inasmuch as the chemical composition of the chlorites appears to have little influence upon the relative stabilities of the polytypes, and chlorite polytypism is potentially a useful geothermometer. In the Carboniferous deposits of the Dniepr-Donetz depression, the conversion of the Ib ( p = 9 0 " ) to the IIb ( p = 97") polytype is associated with the transition between the "stages of initial and deep epigenesis" (see Fig. 5-9), as manifested by the disappearance of illite-smectite mixed-layers and fat-coal rank. In the eastern Donbas, the chlorite retains the IIb ( p = 97") polytype into early and late metagenesis, but tends to become richer in magnesium (Karpova, 1969, pp. 12- 14; Logvinenko and Karpova, 1960). Disappearance of detrital minerals during burial

Many clastic minerals become unstable upon burial and tend to disappear in the course of late diagenesis. Kopeliovich et al. (1961) proposed a generalized order of disappearance of clastic silicate minerals of igneous and metamorphic origin. Other authors, however, have drawn attention to the differences in the order of disappearance of these clastic minerals during diagenesis in different chemical environments (e.g., Koporulin, 1966), although little detailed data are available.

340 Koporulin (1962) has noted the greater stability of garnet relative to other detrital minerals under conditions of montmorillonitization and chloritization than under conditions of kaolinitization of biotite. Zaporozhtseva (1963) has pointed to the marked enhancement of corrosion of clastic grains in sediments rich in organic matter due to the increased acidity of interstitial solutions. In this section, only the alteration of minerals, disappearance of which may indicate a particular stage of deep burial, particularly detrital micas and calcic plagioclase, in pelitic and non-volcanic clastic rocks is discussed. The behavior of the clastic minerals in volcanic clastic rocks is reviewed in a separate section. Alteration of clastic biotite Particular attention has been paid by Kossovskaya and other Soviet authors to the disappearance of biotite. Early during diagenesis, part of the biotite may alter to montmorillonite, glauconite, kaolinite, or ,vermiculite, depending on the environmental conditions (Kossovskaya et al., 1964; Starke, 1968). At the beginning of epigenesis biotite may alter to chlorite or “mobile chlorite” with vermiculite-like layers; in platform successions, through such phases or through swelling chlorite to montmorillonite or, even, to kaolinite, depending on the environmental conditions (Kossovskaya et al., 1963; Kossovskaya et al., 1965; see also Starke, 1968, p.215). In clastic sedimentary rocks of both geosynclinal and platform areas, detrital biotite is altered during deep epigenesis, mostly to characteristic chlorite - hydromica “packets”. It disappears in the lower part of the “zone of quartzitic texture and chlorite-sericite cement” (table in Kossovskaya et al., 1957; Kossovskaya and Shutov, 1958, p. 66 1 ; Kossovskaya, 1960; Scherp, 1963), or during the equivalent “stage of deep epigenesis” (Kossovskaya et al., 1965, pp. 161-166), before the onset of incipient metamorphism as indicated by the appearance of pyrophyllite (e.g., Dunoyer de Segonzac and Millot, 1962). Alteration of detrital muscovite Detrital muscovite also tends to alter during late diagenesis. Desintegration of clastic micas concurrently with the “progressive” mixed-layering of montmorillonite has been reported from deep wells by Dunoyer de Segonzac (1964) and by Perry and Hower (1970). This alteration may be an illitization process, as reflected by broadening of the (001) X-ray diffraction peak (e.g., Jaron, 1967; Dunoyer de Segonzac, 1964). Such peak-broadening should be measured preferably in the size fraction that excludes fine-grained authigenic illite.

34 1 Alteration of intermediate and calcic plagioclase Detrital plagioclase tends to undergo albitization during late diagenesis, the more calcic plagioclase being most prone to alteration. Epidote (“sauswrite”), calcite, and white mica (“sericite”) are common by-products. Characteristic minerals forming as a result of decomposition of intermediate and calcic plagioclase upon burial include the Ca-zeolites heulandite, thomsonite, scolecite, and laumontite (characteristic of mineral assemblages of the zeolite facies) and, at more advanced stages of burial, prehnite, pumpellyite, and wairakite. Inasmuch as such plagioclase and these hydrous Ca-Al-silicate minerals are more common in intermediate and basic volcanic and feldspathic clastic sediments, the alteration of detrital plagioclase is discussed in the section dealing with the development of lowest-grade mineral facies in those rocks. Summary of mineralogic changes in clastic sedimentary rocks during burial diagenesis As discussed in the preceding sections, the general trend during burial diagenesis is the gradual replacement of smectite, detrital biotite, K-feldspar, and calcic plagioclase. At the more advanced stages of burial diagenesis illite-smectite mixed-layers and kaolinite are also replaced by chlorite, illite, and albite. The various factors that may affect the rates at which these various modifications take place include ( 1) the composition of interstitial solutions, (2) porosity and permeability, and (3) the nature of the clastic minerals present. At the advanced stages of burial diagenesis and the onset of incipient metamorphism, with the initiation of the development of slaty cleavage, smectite, clastic biotite, and calcic plagioclase are absent. Illite and chlorite are the only phyllosilicates present in formerly polymineralic clastic rocks. The occurrence of kaolinite-group minerals is restricted to alkali-poor and Mg-Fe-poor aluminous sedimentary rocks. Illite-montmorillonite mixedlayers only occur in metabentonites. In addition, the phyllosilicate species present at this stage differ from their equivalents formed during early diagenesis or the initial stages of burial diagenesis. Illite tends to belong to the well-ordered 2M polytypes rather than the disordered lM, modification. There is also an improvement in its crystallinity. Chlorite tends to develop first the Ib ( p = 90”) and ultimately the IIb (/3 = 97”) polytype, and to become more Mg-rich. Kaolinite becomes well-ordered, or occurs in the form of its 2M polytype dickite. With further increase in temperature and pressure during incipient metamorphism, with the initiation of the development of slaty cleavage, some

342 additional diagnostic minerals appear in the clastic rocks of suitable composition: pyrophyllite, paragonitic mica, the regular mixed-layer rectoriteallevardite, and authigenic illites having high degree of “crystallinity”. These mineral modifications taking place during incipient metamorphism, which includes anchimetamorphism or the “stage of early metagenesis”, are discussed in the following section. Modification of layer silicates of clastic sedimentary rocks during incipient metamorph ism Upon incipient metamorphism, as somewhat arbitrarily delineated in the first section of this chapter, many of the mineral-modification processes of burial-diagenesis described above are completed, whereas others continue to occur. Thus, in this respect there is no sharp limit. A number of characteristic authigenic phyllosilicates, however, make their first appearance during incipient metamorphism including: (1) Pyrophyllite: Al,[Si,O,,/(OH),]. (2) Na-rich, paragonitic micas and regular Na-rich mica-montmorilloni te mixed-layers (including rectorite-allevardite). (3) Concurrently, some qualitative changes take place in the composition of common dioctahedral mica-type layer-silicates (illite and muscovite), including: (a) further increase in “crystallinity”, and (b) change in polytype. These mineral-modification processes, some of which are considered to be characteristic for the “stage of early metagenesis” of Kossovskaya and Shutov, or the anchimetamorphic zone (or anchizone) of Kubler, are discussed below. Composition and polytype of illite-muscouite in incipient metamorphism As discussed in the preceding section, in the most advanced stages of burial diagenesis or epigenesis and, particularly, during incipient metamorphism or metagenesis, there is a tendency (1) for the 2M, illite-muscovite polytype to replace the 1M and lM, polytypes of authigenic dioctahedral three-layer clay minerals, prevalent in the shallower levels of diagenesis and epigenesis, and (2) towards a convergence of different polytype ratios in different rock types. This transformation and convergence to the 2M polytype may be completed only by the onset of the greenschist facies, e.g., in the zone of “late metagenesis” of the Carboniferous deposits of the southeastern continuation of the Donbas (Karpova, 1969, p. 1 l), or upon reaching the biotite isograd, e.g., in the Belt Series of Idaho (Maxwell and Hower, 1967). Further evidence that during the incipient metamorphism muscovite occurs predomi-

343 nantly or entirely in the 2M form is provided by ( 1 ) the underclays of the northeastern-Pennsylvanian anthracites (Hosterman et al., 1970), (2) the anchimetamorphic pencil slates (“Griffelschiefer”) of northeastern Bavaria (Ludwig, 1973), (3) the lower-Paleozoic tuffaceous slates and phyllites of the southern Urals and northern Mugodzhars (Gavrilov and Aleksandrova, 1968), (4) the anchimetamorphic black marly Jurassic deposits (“Terres Noires”) east of Gap in the French Alps (Artru et al., 1969; Dunoyer de Segonzac, 1969, pp. 1571160; see Fig. 5-17) and ( 5 ) the anchimetamorphic Upper Triassic Quartenschiefer of the Helvetic zone in the Glarus Alps of eastern Switzerland (Frey, 1969a, pp. 101- 102; 1970, p. 273). There is a concurrent tendency for the composition of the dioctahedral micas to approach that of muscovite with increasing degree of metamorphism, as evidenced by the increase in tetrahedral charge and K content, and the decrease in octahedral charge and water content (Karpova, 1969, pp. 9-10; Kossovskaya and Drits, 1970). The ranges in tetrahedral charge (0.65 to 0.9) and in octahedral charge (0.1 to 0.4), and in the proportion of bivalent Fe and Mg ions in the octahedral layers (0.25 to 0.40) in “deep epigenesis - metagenesis” are intermediate between those in “low temperature” illites and the muscovites of metamorphic rocks (Kossovskaya and Drits, 1970, p. 89). Moreover, in the zone of incipient metamorphism !he 2M polytypes also predominate in dioctahedral micas of divergent (i.e., non-muscovitic) composition. The fibrous Mg-rich variety of dioctahedral mica (“Gumbelit”), found as an authigenic mineral in the zone of “late epigenesis” and initial metamorphism in association with semi-anthracitic plant matter in the Jurassic deposits of Dagestan and southeastern Azerbaydzhan, has been shown to have the 2M, polymorph (Konysheva, 1968). Fibrous Gumbelite found elsewhere is also always associated with organic matter of semi-anthracitic and anthracitic rank (Schreyer, 1969; Volkova et al., 1974). Esquevin ( 1969) has noted that during incipient metamorphism aluminous illites convert to the 2M polytype at lower grades than do Mg-rich illites, which tend to acquire phengitic composition. Dioctahedral K-micas of a phengitic rather than muscovitic composition (i.e., intermediate between muscovite and celadonite) are common in the zone of lowest-grade metamorphism of the high-pressure, low-temperature type, i.e., in the glaucophane-lawsonite-schist facies and the lower-grade part of the greenschist facies (Ernst, 1963; Velde, 1965b, 1967). Studies of the polytype of phengitic micas from such terranes, however, indicate that 2M, is the predominant polytype (e.g., Velde, 1967; Beugnies et al., 1969; Ernst et al., 1970; Dunoyer de Segonzac and Hickel, 1972). In a comparative study of blueschist-type low-grade metamorphism in the California Coast Ranges and the Sanbagawa metamorphic belt of Japan, Ernst et al. (1970,

344 pp. 155-160; see also Seki et al., 1969, tables 19-22) noted that the phengites of all the Sanbagawa terrane samples and most of the Franciscan terrane samples were of the 2M, polytype. In the Franciscan deposits, in which the micas tend to be more celadonitic than their Sanbagawa equivalents, only a few samples contained subordinate amounts of the 1M polytype. The phengitic mica in the anchimetamorphic zone of the Upper Triassic and Lower Liassic age of the Helvetic zone in the Glarus Alps of eastern Switzerland, also shows strong predominance of the 2M, over the lM, polytype (Frey, 1969a, pp. 101-102; 1970, p. 273). In the Penninic domain of the western Alps, phengites of the 3T polytype (reportedly overgrowing 2M phengites) appear only in the higher-grade part of the greenschist facies (Liborio and Mottana, 1975; see also Dunoyer de Segonzac and Hickel, 1972). It appears, therefore, that during incipient metamorphism, all K-rich dioctahedral micas, irrespective of their origin or composition during diagenesis, tend towards acquiring muscovitic and phengitic chemical compositions, with predominance of the 2M, polytype. Increase in the crystallinity of illite in the anchimetamorphic zone During advanced stages of diagenesis and initial metamorphism there is not only a change in the polytype and chemical composition of authigenic dioctahedral three-layer silicates, but also a progressive increase in crystallinity of the authigenic illite, i.e., increase in the dimensions of crystallites and regularity of the layers. This increase in crystallinity is due to concurrent dehydration, K-fixation, and rearrangement of ions. Increase in crystallinity is reflected in the shape of the X-ray diffraction peaks, e.g., particularly in the increasing sharpness of the first basal diffraction peak (001) at about 10 A of illite-muscovite. Illites having poor crystallinity show broad diffraction peaks, whereas well-crystallized micas exhibit sharp peaks. Several methods have been used to define this peak sharpness. Weaver (1961), the first author to propose and use illite “crystallinity” I as a parameter of burial diagenesis and incipient metamorphism, used the ratio of the illite (001) peak height at lOA to the intensity at the flank of the peak at 10.5 A (“sharpness ratio”) as the parameter of “crystallinity”. The numerical value of this “sharpness ratio”, which increases with improving “crystal-

’

Since the shape and width of the l O A diffraction peak will be shown in the following sections to be affected by several factors in addition to crystallite size, the word “crystallinity” as based on the shape or width of this peak is strictly speaking a simplification and will be placed between quotation marks.

345 linity”, has been shown by Kubler (1964) to be independent of the absolute peak intensity for peaks higher than 50 mm on the recorder chart. Using this “sharpness ratio”, Weaver ( I96 1) distinguished illite “crystallinity” zones ranging from “unmetamorphic”, through three zones of “incipient” to “weak” metamorphism, to “low-grade’’ metamorphic zone in the Paleozoic deposits of the Ouachita belt of Texas and Oklahoma. Kaolinite and/or abundant mixed-layer clay is restricted to the “unmetamorphic” zone. Kubler (1964, pp. 1107- 11 12) used Weaver’s “sharpness ratio” for a study in a well in an Upper Paleozoic series. In subsequent studies, however, Kubler adopted the width of the illite (001) peak as measured at half height (see Fig. 5-1 l), at specified instrumental conditions, as a better index of illite

Fig. 5-1 1. Crystallinity of illite: measurement of the half-height width of the diffraction pcak at 10 A. (From Kubler, 1967a, fig. 1.)

346 “crystallinity”. He noted that Weaver’s measurement techniques, though excellent for the broad peaks of poorly-crystallized illites, involved intolerable errors for well-crystallized illites (Kubler, 1967a; 1968, pp. 388-389). As opposed to Weaver’s technique, the “crystallinity indices obtained by measurement of the half-height peak width and the scatter of these values (experimental error) decrease with increasing degree of “crystallinity” towards the epimetamorphic domain (Kubler, 1968; Dunoyer de Segonzac et al., 1968; see Fig. 5-12). ”

x tl

‘ I CI

x

c -

-t ra

c YI

L x U

m

t

0

E

C U w

Peak wldth

X 4l

U

-

x

Metamorphic

-

c

-c ra

c

y1

L x

U

cn

G ra

.. L

C u

Fig. 5-12. Crystallinity of illite: schematic variation of the error (the horizontal distance between the two curves) as a function of the “crystallinity” for the “sharpness ratio” technique of Weaver (below) and the “half peak-height width” technique of Kubler. Note that with Kubler’s method-contrary t o the Weaver technique-the error diminishes towards better “crystallinities”. i.e. towards the lower “crystallinity indices” of the “epimetamorphic” domain. (After Kubler, 1968. fig. 2 . )

347 Using primarily the limiting values of illite “crystallinity indices” I . which correspond to Weaver’s “incipient” and “weak” metamorphism, Kubler (1964, pp. 1107- 11 12; 1967~1,pp. 110-1 1 1 ) revived Harrassowitz’ (1927) term “anchimetamorphism” for the range of phenomena and processes occurring between disappearance of some characteristically sedimentary minerals and onset of the greenschist facies. Dependence of the illite “crystallinity index on conditions of nieusurenient The I O A peak width or “crystallinity index” of illite has subsequently been measured and used as a parameter in evaluating stages of advanced diagenesis and incipient metamorphism by Dunoyer de Segonzac in Strasbourg, France, Frey in Bern, Switzerland, Weber in Bochum. West Germany, Aprahamian in Grenoble. France, and by the present author in Beer-Sheva. Israel. In comparing the illite “crystallinity indices” in these various studies i t must be borne in mind that the absolute value of the peak width at half height depends on the experimental conditions adopted. ‘Even using the same radiation (Cu K a was used in all of the above studies). the half-height peak width expressed in degrees angular separation ( A 2 8 ) is dependent on variables that probably cannot be exactly reproduced except on two identical diffractometers, i.e., identical width of the divergence, scatter, receiving slits, filters, monochromators used, etc. The peak width which is commonly expressed in mm. also depends on the relation of the diffraction angle 26 to mm on the recorder chart, i.e., on the scanning speed in “26/min and the recording speed in mm/min. The latter effect can be normalized by expressing the peak width in degrees A28 rather than in mm which, surprisingly, is done very rarely. Alternatively, the ratio of the width of the illite (001) peak to that of some external standard, such as the (100) peak of quartz at 4.255A, can be given rather than the absolute width. Weber (1972a, b) used the “relative peak width at half height”: ”

Hb (001) illite, mm x 100 Hbrel H b (100) quartz, mm The experimental conditions used by various authors, specified in their papers, and the resulting illite “crystallinity” values equivalent to the limits of the anchimetamorphic zone as proposed by Kubler (1967a. 1968), are compiled in Table 5-IV. Flehmig (1973) has proposed an infrared method for determination of

’ 4.0 and 2.5 mm under the experimental conditions adopted.

TABLE 5-IV Experimental conditions adopted by various authors for the determination of illite "crystallinand the corresponding illite-"crystallinity" values ity" using the diffraction peak at 10 equivalent to the limits of the anchimetamorphic zone of Kubler (1967, 1968)

A,

Author

Also used by

Experimental conditions (see notes)

Equivalent of 1'26 on chart (in mm)

*

Weaver (1961) Kubler (1967a)

Kubler (1967b), Artru et al. (1969)

10

Kubler (1968; also in Dunoyer de Segonzac et al. 1968)

Frey ( 1969a, 1970)

13.3

Dunoyer de Segonzac ( 1969)

Dunoyer de Segonzac and Heddebaut ( 1971) Abbas (1 974)

8.61

Chennaux et al. (1970)

6.66

Sagon and Dunoyer de Segonzac ( 1 972)

10

Weber (1972a, b)

20

Ludwig (1972b, 1973)

20

Aprahamian (1974)

(9)

13.3

Barlier ( 1974)

(10)

7.62

(11)

40

Kisch (1978)

Kisch (1980a,b)

Experimental methods: ( I ) No data available. (2) Philips. XY units, 8"/min (no further details provided). (3) Philips. CuK,. Slits lo-0.2 mm-I". T C = 2 or 4. Goniometer 2"/min; chart 1600 mm/h, i.e., 1" 2 6 = 13.3 mm on chart. (4) Philips 1010. CuK,, Ni-filter. Slits 2"-0.1 mm--2". Sensibility 1. lo2 at T C = 10 to 1 . lo3 at TC=4. Goniometer 2" 2 6/min, Hewlett-Packard X-Y recorder set for 1' 2 6 =8.61 mm on chart. (5) Goniometer 2"/min; chart 800 mm/h, i.e. 1" 2 6 =6.66 mm on chart.

349

Low-grade illite"crystallinity" limit of anchimetamorphic zone in mm (and "2 8 ) 2.3

*

High-grade illite"crystallinity" limit of anchimetamorphic zone in mm (and ' 2 8)

Maximum "crystallinity" of micas in mm (and ' 2 8)

12.1 *

4 (0.4" 2 8)

2.5 (0.25' 2 8)

1.8 (0.180 2 8)

7.5 (0.56' 2 8)

4 (0.30' 2 8)

approx. 2.8 (0.21' 2 8)

5.5 (0.64' 2 8)

3.5 (0.410 2 8)

approx. 2 (0.23" 2 8)

(**I 2.7 (0.400 2 8) 2.7 (0.27' 2 8) 2-6 p m fraction or polished stabs approx. 5 (0.25" 2 8) [Hb,,, 150-1551 ***

approx. 3.4 (0.17' 2 8) [Hb,,, 105-1 101 ***

approx. 3 (0.15' 2 8)

4.0 (0.20' 2 8) [Hb,,, 1.251 ****

approx. 3.2 (0.16' 2 8) Hb,,, 1.00 ****

2-6 pm fraction 5.8 (0.290 2 8)

[Hb,, 1.811 ****

<2 p m fraction 7.1 (0.350 2 8) [Hb,,, 2.221 **** 5.2 to 5.3 (approx. 0.4' 2 8)

2.8 to 2.9 (approx. 0.25" 2 8)

4.5

3

15.2 (0.380 2 8)

8.4 (0.21' 2 8)

approx. 4.5 (0.11' 2 8)

(6) Sensibility 2 . lo3, TC=2. Goniometer 2' 2 B/min; chart 1200 mm/h, i.e., lo, 2 8 = 10 mm on chart. (7) Philips. CuK,. Slits ' (divergence)- 1' (receiving). Monochromator, discriminator. TC =4 at goniometer ' 2 8/min and chart 300 mm/h, or TC =2 at goniometer f ' 2 8/min and chart 600 mm/h, i.e., 1' 2 8 =20 mm on chart. For measurement of the (100) reflection of the external quartz standard at 20.85', 2 8 the I' divergence slit was used (all other instrumental conditions unchanged); the half peak-height width of the quartz (100) reflection is 3.2 to 3.3 mm. The 2-6 p m size fraction, or rock slabs polished parallel to the slaty cleavage are measured rather than the < 2 p m fraction used by most workers.

a

350 TABLE 5-IV (continued) (8) As above (Weber, 1972a, b). but T C = 1 at goniometer I ” 2 6’/min and chart 20 mm/min, i.e., I ” 2 0 =20 mm on chart. Uses both the 2-6 p m and the ( 2 p m size fractions. Half peak-height width of the quartz (100) reflection is 3.2 mm. (9) Philips. CUK,, Ni-filter. Slits 1”-0.1”-1”. T C = 2 at scale 4 . lo2 to 2. l o 3 (usually 1 . lo3).Goniometer 1” 2 B/min; chart 800 mm/h, i.e., I ” 2 6’= 13.3 mm on chart. (10) Siemens Co. Monochromator. C G R diffractometer. Sensibility 50 cps at TC=5. Goniometer I”/min; chart 457.2 mm/h, i.e., I ” 2 0 =7.62 mm on chart. (11) Philips. CuK,, Ni-filter. Slits lo-0.2 mm-I”. T C = 2 at scale 1.103 or 2. l o 3 (usually 1 . lo3),or very rarely T C = 4 at scale 4 . lo2.Goniometer 0 . 5 O 2 6’/min; chart 20 mm/min, i.e. 1 ” 2 6‘ =40 mm on chart. * Weaver (1961) expresses the sharpness of the illite 10 A peak as the ratio of the peak height at 10.0 A to the height of the low-angle side of the peak at 10.5 A, the “sharpness ratio”. Weaver recognized two more zones in what is now referred to as the anchizone, “incipient to weak metamorphism” and “weak to very weak metamorphism”, with average sharpness ratios of 4.5 and 6.3, respectively. These values are equivalent to respectively 5 mm and 4.5 mm in the scale of Kubler (1967a, fig. 3). * * Le Corre (1975, fig. I). *** In order to minimize instrumental error, Weber (1972a, b) used the relative half peak-height width Hb,,,, which is the ratio of the half-height width of the illite (001) peak and of the quartz (100) peak at 20.85” 2 8 , multiplied by 100. The Hb,,, values are given in square brackets. **** The Hb,, values used by Ludwig (1972b, 1973) have not been multiplied by 100.

illite “crystallinity” by measuring the extinction ratio of selected absorption bands. This method is claimed to be more exact than X-ray methods in establishing degree of metamorphism in Paleozoic slates from western Germany. Smykatz-Kloss and Althaus ( 1974) have found lack of agreement between “crystallinities” and lOA peak widths of two standard illites as measured by infrared and X-ray methods. Flehmig’s method may well prove to measure differences in chemical composition rather than in crystalline size of illite. Complications in the X-ray determination of the illite “crystallinity index” In evaluating illite lOA peak widths as parameters in studying the lowest-grade of metamorphism, a number of complicating factors should be born in mind, that have been reviewed by Dunoyer de Segonzac (1970) and are only briefly discussed here. ( I ) Effect of detrital micas. The “crystallinity” of detrital micas is generally better than that of authigenic illites from the same or adjoining sedimentary rocks (the 10 A peak widths are consequently narrower). Detrital micas that

35 1 have experienced only weak chemical weathering will show particularly good “crystallinity”, as exemplified by the anomalously narrow 10 A peaks of the Upper Ordovician periglacial sediments in the Sahara (Chamley, 1967, 1968; Dunoyer de Segonzac, 1970, pp. 97 and 116). The mean illite “crystallinity index’’ of samples containing a mixture of authigenic and detrital illitemuscovites will, therefore, be smaller than that of the authigenic illites that indicate the degree of incipient metamorphism. Inasmuch as the detrital micas are generally coarser than the authigenic illites, the effect of size can to a large extent be avoided by using as an index of metamorphism only the lOA peak width of the clay-size ( < 2 pm) fraction which is largely free of detrital micas. This approach is adopted by most authors; however, Weber (1972a, b), followed by Ludwig (1972a, b), used the more crystalline 2-6 p m fraction instead, in order to avoid possible peak-broadening by less crystalline very fine fractions (< 0.2 pm). Although the writer (Kisch, 1980a, 1980b) prefers to measure the 10A peak width of the < 2 p m fraction and use it as a measure of the degree of metamorphism, he also measures the l O A peak widths of 2-6 pm, or the < 6 p m and < 20 p m fractions, in order to estimate the contribution of clastic micas to the “crystallinity” of the finer fraction. Within the anchizone, the lOA peak widths of different illite size fractions tend to converge (Weber, 1972a). (2) Effect of host-rock lithology and K-deficiency. Variability of “crystallinity” of illites from different beds is considered by Kubler (1968, p.391) to be characteristic of the “diagenetic” as contrasted with the anchimetamorphic realm. Here, illite-smectite mixed-layers are commonly still present, the widths of the broad composite (001)/(001) peaks between 10 and 12A giving anomalously broad “illite” peaks. The peak widths of such mixedlayers are affected by chemical treatments. They are reduced through saturation with 0.1 N KCI relative to the untreated samples and to the 1 N MgCl, -saturated samples, while glycolation results in a further reduction in peak width (Biljon and Bensch, 1970; Ludwig, 1972b, pp. 554-556; Kisch, 1980b). The illitization of smectite and the transformation of illite to muscovite. which are accompanied by reduced peak widths, both require potassium cation and, thus, may be retarded by deficiency in K. Authigenic illites formed during burial diagenesis in more porous rocks, such as sandstones, tend to have better “crystallinities” than those formed in shales. The “crystallinity” of illites in evaporitic rocks also appears locally to be better than in the accompanying shales, marls, or limestones (Kubler, 1968, p. 390) apparently due to the high K-concentrations. On the other hand, illites from coaly or bituminous sediments and from pure limestones and dolomites, in which organic matter forms the bulk of

352 the insoluble residue, often have lower “crystallinities” than in the associated rocks (Kubler, 1968, p. 391; Ludwig, 1972b, pp. 557-558). In the case of extreme deficiencies in potassium with respect to sodium, progressive burial may actually result in deterioration of the illite “crystallinity”, i.e., 10 A peak broadening, and subsequently in the evolution of Na-rich regular mixed-layers, which are more fully discussed below, in addition to or even instead of K-micas (Kubler, 1967a, p. 1 I 1 : 1968, pp. 390-39 1). Disregarding these compositional extremes, however, the range of variation of the illite 10A peak width with lithology tends to diminish during anchimetamorphism (Dunoyer de Segonzac et al., 1969; Kubler, 1967a, 1968). (3) Broadening of the illite (001) peak resulting from the presence of Nu-rich phyllosilicates. A number of mica-type layer silicates that may accompany illite in the anchizone or in the epizone have basal reflections close to I O A . For Na-mica paragonite the basal reflections are at 9.6-9.7 A, for the regular paragonite-montmorillonite mixed-layer rectorite or paragonitk at about 12 A, for the Ca-mica margarite at 9.7 A, and for pyrophyllite at 9.2 A. Poor resolution of these peaks from the illite (001) reflection may result in I O A peak-broadening and thus in an anomalously low apparent illite “crystallinity”. Peak broadening due to the presence of rectorite-allevardite has been reported by Kubler (1967a) and by Dunoyer de Segonzac and Heddebaut (1971) from the anchizone of the Basque Pyrenees. Peak broadening related to the presence of paragonite and of margarite was noted by Sagon and Dunoyer de Segonzac (1972). The peak-broadening effects of paragonite in a section from the northern Glarus Alps (Frey, 1970), ascribed by Ludwig ( 1972b) to paragonite-phengite mixed-layer, are shown in Fig. 5-1 3. Kubler (1967a, 1968) ascribed the formation of paragonite and rectoriteallevardite during incipient metamorphism to high Na/K ratios. Such evolution towards allevardite is considered to be particularly common in the presence of carbonaceous matter (see also Henderson, 1970). In such environments, the illite “crystallinity index” may actually increase with increasing depth and in temperature, contrary to the general behavior. Attention should, thus, be paid to the detection of Na-rich hydromicas and mixed layers, and anomalously broad 10 A peaks in such samples should not be considered to represent illite “crystallinity”. ( 4 ) Relationship between aluminium content of illite and “crystallinity”. Even in

the absence of expandable layers there exists a dependence of the development of “crystallinity” upon the chemical composition of the illite. Esquevin (1969) has argued that high Al/(Fe Mg) ratios in the octahedral layer of

+

353 >,

F

5

CLAY MINERALS

CARBONATE AND CRlSTALLlNlTY OF ILLIT€ QUARTZ CONTENT 53 Distribution inthe fraction Wciqht-% of t h e index after Kubler s tz t <2P total rock v)C 0 20 40 60 00 1000 20 40 6 0 00 100, 0

22

MF66O

MF 659 80 m

II

MF65B HF657

I

7

MF656 MF655

I

60m

MF65L

MF653 MF652

40 m M f 651

HF 650

MF649

MF64E

Mf 617 20 m

parag.

MF 6L6 MF66 MF644 MF6L1 MF6LO Om

MF639

chlorite

pyrophyllitc

phrngite

c7 quartz

LL I l l l l l l l l l j 0 20 40 60 80100 0

Fig. 5-1 3. Mineralogy and illite-crystallinity of shales and slates of the anchimetamorphic Lower Liassic from a profile in the northern Glarus Alps at “Guggenegg”. p u . / p h e . = mixedlayer paragonite-phengite. Note the anomalously high illite-“crystallinity indices” due to peak-broadening (indicated by open circles) in most samples containing paragonite. (From Frey, 1970, fig. 4.)

354 illite favor increased “crystallinity” during anchi- and “epimetamorphism”. Using the ratio of the intensities of the (002) diffraction peak at 5 A and the (001) diffraction peak at 10 A as an approximate index of the Al/(Fe Mg) ratio in the octahedral layer, Esquevin found that in some areas the magnesian illites with Im2/Imi ratio below 0.3 from the “diagenetic zone” and the “epizone” show indistinguishable peak widths (see Fig. 5- 14). Esquevin considers the 10A peak width a dependable indicator of the grade of incipient metamorphism only when the intensity ratio I,,/I,, is above 0.3, which indicates a high Al/(Mg Fe) ratio. Characterization of the illites using both the (001) peak width and the intensity ratio I o o 2 ~ 5 ~ ~ / I o o l as ~ l oproposed ~~, by Esquevin (1969), has been successfully used by Artru et al. (1969) and Dunoyer de Segonzac (1969) for a better distinction of zones of incipient metamorphism (see Fig. 5-15). Dunoyer de Segonzac and Heddebaut (1971) and Dunoyer de Segonzac and Hickel ( 1972), however, subsequently noted lack of correlation between the I,,,/I,,, and the Al/(Mg + Fe) ratios in illites from some anchi- and “epi”-metamorphic terranes. In these cases the lOA peak width is uniform and independent of the 1,,,/1,,, ratio. Moreover, Clauer and Lucas (1970) have shown that the 002/001 peak height ratios are affected by cation treatment, tending to be somewhat higher in Mg-saturated than in K-saturated illites.

+

+

Width of (001) peak

I

0 00 0

.... .

25 2-

1

01

*-

. ..

-*

I

I

,

0.2

03

04

~

I002 I001

c

05

Fig. 5-14. Relation between the width (in mm) of the diffraction peak at 10 A and the ratio of the intensities of the peaks 002 and 001 in the Middle Jurassic of the wells Pont d’As 1 and Saint-Faust 2 (closed circles) and Orthez 102 (open circles) near Pau, southern France. The ratios) show anomalously broad magnesian illites from the former two wells (low I,,,,/I,,, peaks. (From Esquevin, 1969, fig. 2 . )

355 Width of (OoOl) Deak at 10 A

A 13

-

12 1110-

90-

76-

54-

3-

2L 1

0.10

‘Jj4-’

BARCELONNETTE

I

I

0.20

0.30

-

I(002) I(OO1) 0.40

Fig. 5-15. Diagram after Esquevin (1969) showing the distributic of illite 10 A peak widths as a function of the intensity ratio of the illite (002) and (001) peaks in the Jurassic “Terra Noires” deposits of a west-east section through the subalpine zone of the French Alps. The localities of the samples are shown on Fig. 5-17. (From Artru et al., 1969, fig. 2.)

Applicability of illite “clystallinity ”. The above-mentioned complications can to a large extent be minimized, however, by taking the precautionary measures indicated. With such precautions, illite 10 A peak width, particularly as determined on several samples from one locality, remains one of the most suitable and generally applicable mineralogical monitors of deep-burial diagenesis and incipient metamorphism in clay-rich sedimentary rocks. Studies of regional increase in illite “crystallinity” Illite “crystallinity” has proved to be a useful index of high-grade diagenesis with increasing depth in deep oil wells, e.g., in ( 1 ) the Upper Cretaceous sequence of the Douala Basin, Cameroun (Dunoyer de Segonzac et al., 1968), (2) the Jurassic shales and marls of the Provence, southeastern France (Dunoyer de Segonzac, 1969, pp. 164- 174), (3) the Permo-Triassic deposits of the Paris Basin (Dunoyer de Segonzac, pp. 148-151), and (4) the Paleozoic deposits of the Polignac Basin, Algerian Sahara (Dunoyer de Segonzac, pp. 86-102).

356 The main value of illite “crystallinity”, however, lies in providing a parameter for the establishment of burial-diagenetic and incipientmetamorphic zonation in lowest-grade areas and, particularly, in extending the scope of metamorphism beyond the classical metamorphic zones. Such zonation was first established for Alpine terranes by demonstrating that illite 10 A peak widths decreased towards the internal zones, i.e., (1) southward in the Cretaceous deposits of the sub-Pyrenean and north-Pyrenean marginal zones of the western Pyrenees (Kubler, 1967a; see also Dunoyer de Segonzac et al., 1968, fig.2; see Fig. 5-16) and (2) eastward in the comparatively uniform (lithologically) Jurassic marly shales (“Terres Noires”) of the external, subalpine zone of the French Alps (Dunoyer de Segonzac et al., 1966; Dunoyer de Segonzac, 1969, pp. 157-160; Artru et al., 1969; Barlier, 1974; Barlier et al., 1974; Aprahamian, 1974; see Fig. 5-17). Illite “crystallinity” zonation has also been demonstrated for Paleozoic deposits, that underwent incipient metamorphism of Hercynian age in (Projection ) ASSON

( Projection )

A R T H E Z d’ASSON

S

I

I

SCHISTOSE ZONE

0

1

I M I C R O F O L D E D ZONE

2

3

4

I

NAY

I

N

MONOCLINE

z

5 km

Fig. 5- 16. Schematic profile showing the southward increase in illite “crystallinity indices” (shown as circles) in the Cretaceous deposits of the sub-pyreman and the schistose northPyrenean marginal zones of the western Pyrenees between Nay and Arthez d’Asson, south of Pau, France. The appearance of slaty cleavage at the north-Pyrenean fault (marked x ) , which bounds the schistose zone on the north side, is associated with illite “crystallinity indices” characteristic of the low-grade part of the anchizone (10 A peak width of 4 mm at the conditions of measurement). (From Kubler, 1967a, fig. 8.)

357

VEYNES

0

SAMPLE LOCALITIES

1

ASPRES

L

p , 5 , l,Okrn

0---

100%1Md 0.

-. 0..

0..

I

CLAY MINERALS

0

’.‘.,

a’

, ;

y=.?$lh - ._

E Blllite-rnontmorillonite

0

0

..Of

!?, 0

Pobnorphr -O-

Kaoiinite

!lUllIl Ailevardite

hChlorite 0 Illite

0

-- - -8.x.o’. .\,\-- - - __ _

-10

145

0

-100% 2 M

0

A

1

Polymorphr

0

-25

I % I Md I

Fig. 5-17. Change in organic chemistry and clay mineralogy in the Jurassic “Terres Noires” through the subalpine chains from Die eastwards through Gap towards the front of the Alpine nappes at Embrun and Barcelonnette, southeastern France. (After Artru et al.. 1969. figs. 3 and 4; also after Dunoyer de Segonzac, 1969, fig. 65).

northern France (Dunoyer de Segonzac et al., 1968, fig.3), the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971), the central plateau of the Moroccan Meseta (Pique, 1975), and the northeastern Rheinische Schiefergebirge (Weber, 1972b), and for the Cambro-Silurian Jamtland Supergroup of the eastern Caledonides in western Sweden (Kisch, 1981a). I n the comparatively uniform lithologies of the Upper Triassic and lower Liassic marls and pelites of the Helvetic zone in eastern Switzerland it has

358 been followed from non-metamorphic to the amphibolite facies (Frey, 1969a, 1970, 1974). Illite “crystullinity and the development of schistosity Kubler (1967a, 1967b) has commented on the relationship of the onset of anchimetamorphism, as defined by illite “crystallinity”, and the first development of schistosity. He noted that the anchimetamorphic zone was first recognized in schistose series. In both the north- Pyrenean “zone cenomanienne schisteuse” and in the Paleozoic deposits of northern France and southern Belgium the beginning of schistosity is associated with anchizonal illite “crystallinities” (l967a, pp. 115-1 16). Kubler (1967a, p. 116; 1967b, pp. 270-272), however, described the development of schistosity 500m ubove the zone of the first attainment of anchimetamorphic illite “crystallinities” in Cretaceous black marls in an oil well in Aquitaine, France. In the “Terres Noires” of the subalpine chains of the western French Alps, the Alpine slaty cleavage begins to develop near Gap, before illite 10 A widths enter the anchimetamorphic range towards the east (Artru et al., 1969; Dunoyer de Segonzac, 1969, pp. 157-160; see Fig. 5-17); a similar relationship holds in the northwestern part of the central plateau of the Moroccan Meseta (Pique, 1975). AIternatively, where the regional metamorphism is thermal rather than dynamic, anchimetamorphic zones may lack any trace of schistosity. Kubler concluded that the degree of deformation, which is essential for the development of schistosity, does not play any role in the increase in illite “crystallinity”, which appears to be mainly controlled by thermal effects (1967b, 1968). I t must be stressed, however, that the convergence of slaty cleavage and anchimetamorphism is the general case in the above-mentioned regional studies, as well as in the Paleozoic deposits of the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971) and of the northeastern Rheinische Schiefergebirge (Weber, 1972b). Deformation in the latter two areas is Hercynian. Dunoyer de Segonzac and Heddebaut (1971) correctly pointed out that the anchizone is developed only in orogenic and suborogenic zones, and does not occur in unfolded sedimentary basins. In the writer’s opinion, this does not necessarily mean that pressure or shearing stress constitute important factors in increasing the “crystallinity” of illite. I t may merely reflect the stronger postburial uplift in orogenic zones. I t appears, therefore, that temperature is the main factor which controls the increase in illite “crystallinity”. I’

359 Further mineralogical characteristics of anchimetamorphism The anchizone as defined by the “crystallinity” of illite has a number of other mineralogical characteristics whenever it has been studied, the most important of which include: (1) the absence of smectite, irregular illitesmectite mixed-layers, and kaolinite; (2) predominance of illite and chlorite as layer silicates in clastic sedimentary rocks; (3) local development of the regular paragonite-montmorillonite mixed-layer rectorite-allevardite and of pyrophyllite, A1,[Si,Olo/(OH),], often in association with each other (see sections on these minerals); (4) development of Na-rich minerals such as paragonitic micas (including paragonite-phengite mixed-layers) and albite (Dunoyer de Segonzac and Heddebaut, 1971; also see Kubler, 1967a, 1968; Dunoyer de Segonzac, 1969, 1970; Frey, 1969a; and others); and ( 5 ) predominance of the 2 M polytype in dioctahedral K-micas (Artru et al., 1969; Frey, 1969a, 1970). Locally, pumpellyite and prehnite are found in lithic sandstones associated with anchizonal illite “crystallinities” in slates (‘Frey, 1970; Weber, 1972b). Another metamorphic mineral found in the anchizone is stilpnomelane (Frey et al., 1973; Sagon and Dunoyer de Segonzac, 1972). The coal and coaly matter associated with the anchizone generally reaches anthracite rank (Kisch, 1974). Characteristic greenschist-facies minerals of pelitic rocks such as biotite and garnet, do not occur in the anchizone. The same ‘is probably true for chloritoid (cf. Kisch, 1974, p. 86), except in unusually iron- or manganese-rich rocks (e.g., Dunoyer de Segonzac and Heddebaut, 1971). Formation of prophyllite Pyrophyllite (Al,[Si,Olo/(OH),]), a dioctahedral three-layer phyllosilicate with a basal X-ray diffraction peak at 9.15A, occurs in alkali- and Mg-Fe-poor aluminous low-grade metamorphic slates. Formerly, it was regarded as characteristic of the greenschist facies (e.g., Winkler, 1965, 1967) and hydrothermal parageneses (e.g., Stadler, 1971b). Lately, however, it became evident that pyrophyllite may appear at even lower grades of metamorphism, i.e., during incipient metamorphism. Dunoyer de Segonzac and Millot ( 1962) interpreted the pyrophyllite present in the Lower Devonian arkosic sandstone of the Armorican Massif as a diagenetic alteration product of feldspar. They mentioned other pyrophyllite occurrences of diagenetic and lowest-grade metamorphic origin in the Silurian-Devonian deposits of the Algerian Sahara. On the basis of the appearance of pyrophyllite at the expense of dickite in (1) the lower horizons of the “stage of deep epigenesis” in the Russian

360 platform (Shutov and Dolmatova, 1961) and (2) quartzitic sandstones of two other areas, Kossovskaya and Shutov ( 1963) recognized a quartz-pyrophyllite “facies”, considered to be characteristic of metagenesis in the compositional ”family of monomineralic quartz sandstones and kaolinite clays.. .”. Subsequently, pyrophyllite was recognized in the “stage of early metagenesis” in other areas of the U.S.S.R., e.g., in the late-Precambrian argillites of the southwestern Yenisey range, central Siberia (Kasanskiy, 1967). Kubler (1967a, 1968) has noted that the formation of pyrophyllite at the expense of kaolinite-group minerals in aluminous rocks takes place near the low-grade limit of the anchimetamorphic zone, e.g., at the southern limit of the sub-Pyrenean zone in the Department Basses Pyrenees, southern France. On the basis of this diagnostic transformation, he correlated the onset of anchimetamorphism with that of metagenesis of Kossovskaya and Shutov (1963). Subsequently, several pyrophyllite occurrences have been shown to be associated with illite 10 peak widths characterizing the anchizone. These include: ( 1) The widespread pyrophyllite occurrences in the clayey-silty Silurian and Devonian deposits of the Tindouf Basin, the Ougarta, and the Grand Erg Occidental, western-Algerian Sahara (Chennaux and Dunoyer de Segonzac, 1967; Dunoyer de Segonzac and Chamley, 1968; Dunoyer de Segonzac, 1969, pp. 113-1 19; Chennaux et al., 1970; for “crystallinities” of the associated illites see the latter two papers). (2) Pyrophyllite-bearing Ordovician slates of the Ebbesattel near Plettenberg, eastern Sauerland, Rheinische Schiefergebirge (Scherp et al., 1968; for “crystallinity” of associated illite see Weber, 1972b), and the Gedinnian (Lower Devonian) slates and sandstones of the Siegener Sattel in the same region (Weber, 1972b). (3) Pyrophyllite-bearing Upper Triassic “Quartenschiefer” and Lower Liassic of the Helvetic zone of the Glarus Alps, eastern Switzerland (Frey, 1969a, 1969b, 1970; see Fig. 5-13 and 5-18). (4) Pyrophyllite in the Devonian of the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971; see Fig. 5-20). ( 5 ) The pyrophyllite in the kaolinite- or dickite-bearing Silurian tuffaceous slate of the Frankenwald, northeastern Bavaria (Ludwig, 1972a; see Fig. 5-20), which may be associated with the transition between anchizone and greenschist facies (epizone). Several additional pyrophyllite occurrences, for which no data on the 10 A peak widths of the associated illites are available, may be assigned with confidence to a zone lower than greenschist-facies incipient metamorphism i n view of the anthracite rank of the associated coals (cf. Kisch, 1974, table I I).

A

36 1

A

0 Pyrophyllite 0 Paragonite

* A

+

Rhodusite Stiipnorneiane Chloritoid Purnpcllyitc

Fig. 5-18. Distribution of some Alpine metamorphic minerals in the Glarus Alps and the cover of the eastern Gotthard Massif. Most of the minerals are restricted to certain stratigraphic units: the pyrophyllite, paragonite, and chloritoid minerals occur mainly in the Liassic and the Quartenschiefer deposits; stilpnomelane occurs in Fe-oolitic or glauconitic horizons of the Middle Jurassic, Lower Cretaceous and Eocene deposits; pumpellyite (and prehnite) in the Lower Tertiary volcanic graywackes. (After Frey, 1970, fig. 10.)

(6) The pyrophyllite-sericite-chlorite tonstein in the Upper Carboniferous anthracite seam “Trois Bancs” in the southern Belledonne near Grenoble, western French Alps (Scherp et al., 1968; Stadler, 1971a). (7) The pyrophyllite- and kaolinite-bearing underclays and pyrophyllitebearing partings of the Pennsylvania anthracite regions (Spackman and Moses, 1961 ; Hosterman et al., 1970). In the case of some other regional occurrences of pyrophyllite, it is not clear whether they have formed under greenschist-facies or under anchimeta-

362 morphic conditions, e.g., in the matrix of Devonian quartzite and conglomerates in the Brooks Range, northeastern Alaska (Reed and Hemley, 1966), and in the Ordovician and Devonian manganese-rich slates of the Venn-Stavelot Massif, west Germany-east Belgium (Kramm, 1973). Some slates of the latter occurrence contain kaolinite, which is thought by Kramm (1973, p. 184) to be disproportionated to pyrophyllite and andalusite, in association with phengite, paragonite, quartz, spessartine, and minor amounts of chlorite. Chloritoid appears in the isofacial mineral assemblages of host rocks with lower 2 Fe,O, X 100/2 Fe,O, FeO oxidation ratios. The occasional persistence of kaolinite to such advanced, presumably greenschistfacies metamorphism is not well understood at the present time; it may well represent disequilibrium or retrograde metamorphism.

+

Conditions of the kaolinite-pyrophyllite-quartz The equilibrium conditions of the reaction: Al,[Si,O,/OH/(OH),] kaolinite

+ 2 SiO,

-

quartz

equilibrium

Al,[Si,0,,,/(OH)2]

+ H,O

pyrophyllite

water

have been the subject of several experimental investigations, the earlier of which indicated equilibrium temperatures of 400-420°C at 2 kbar p H Z 0 . More recent experimental work, however, suggests the following equilibrium temperatures ("C) at the indicated pressures ( pH20):

500 bars 1 kbar

References

2 kbar

Hemley and Jones ( 1964) Velde and Kornprobst (1969) Althaus ( 1969) 340"

380"

Thompson (1970a) *

325" * 20" 345" 2

368"

3 10" 380"

34

kbar

405 " (6.9 kbar) 10" 375" * 15" (4 kbar)

* Equilibrium conditions determined by studying the weight change of quartz crystals in the presence of fine-grained powdered kaolinite and pyrophyllite, and excess water.

Hemley (unpublished data, in: Reed and Hemley, 1966) produced pyrophyllite from kaolinite quartz at 300°C at 1000 bars p H Z O after 1 year, suggesting that even Thompson's temperatures are on the high side (cf. Haas and Holdaway, 1973, p. 462). At p H z O < Ptota,, or at increased silica activities, this equilibrium tempera-

+

363 ture would be further depressed. Increased acidity, which reduces the fugacity of water, has a similar effect: at 2000 bars p H Z O pyrophyllite form from kaolinite quartz in an acid milieu at 260°C, more than 100°C lower than in pure water (Althaus, 1966; see also Scherp et al., 1968, pp. 158-159). At very low vapor pressures of H,O, Fournier (written communication, in: Scherp et al., 1968) succeeded to form pyrophyllite at temperatures below 200°C.

+

Rare occurrence of pyrophyllite As noted earlier, in the presence of high concentration of potassium, magnesium, or iron during diagenesis, kaolinite alters to illite and/or chlorite rather than to pyrophyllite. This, to a large extent, accounts for the relative rarity of pyrophyllite in normal clastic sedimentary rocks. Nevertheless, the writer believes that because of its optic similarity to muscovite, pyrophyllite may often have been overlooked and mistaken for colorless mica. Thus, it may be a more common mineral in slates than would appear from the currently known occurrences. Velde ( 1968) has proposed that the infrequent occurrence of pyrophyllite within its range of stability in low-grade metamorphism, as compared to the common occurrence of kaolinite in pelitic sediments and sedimentary rocks, is due to a change in the oxidation state of iron. He noted that in Recent and low-grade diagenetic pelitic rocks and in the iron-bearing phyllosilicates commonly found in those rocks, i.e., dioctahedral montmorillonite and, to a somewhat lesser extent, vermiculite, hydrobiotite, and Fe-rich chlorite, the iron is predominantly in the ferric state, whereas the chlorite of the low-grade metamorphic (and burial-diagenetic) paragenesis illite-chlorite is richer in Fe2+. Velde related this difference to the general phenomenon of decrease in the oxidation state of iron as pelitic rocks are metamorphosed, ascribing the reduction of iron from the ferric to the ferrous state during low-grade metamorphism to the influence of hydrogen-rich organic gases released by distillation of primary organic matter. The reduction of the iron results in its removal from the mica-like, three-layer phases and in its being forced into the chlorite phase. This necessitates the removal of aluminium from kaolinite and its entrance into the mica and possibly into the chlorite to replace the lost Fe3 . This could explain why pyrophyllite is found only in very Fe-poor aluminous low-grade metamorphic rocks, commonly in association with the regular Na-rich mica-montmorillonite mixed-layer rectorite (or allevardite), which is discussed in the next section.

364 Appearance of paragonitic micas and Na-rich regular mixed-layers Another mineral from aluminous pelitic rocks, formerly thought to appear only in the greenschist facies (e.g., Winkler, 1965, 1967), but currently being reported from deposits exposed to lower grades of metamorphism, in the Na-rich dioctahedral mica paragonite (NaAI,[AlSi,O,,/(OH),]), the sodic analog of muscovite. Optically, the mineral is virtually indistinguishable from muscovite. In order to identify paragonite in the presence of K-mica in X-ray powder diffraction, suitable high-resolution diffractometer settings should be chosen: slow scanning rates and narrow receiving slits. Only then the basal 9.7 A peak of paragonite is resolved from the adjoining 10A peak of illite and K-mica. The appearance of only one broad peak because of non-resolution of these two diffraction peaks may give the mistaken impression of poor illite “crystallinity” in the advanced stages of incipient metamorphism during which paragonite becomes common. Although the second-order peaks of paragonite and K-mica at 4.85 A and 5.0 A, respectively, are better resolved, they are weaker. The appearance of paragonitic hydromica during incipient metamorphism has been reported from deposits of the “stage of early metagenesis” of the central and eastern Donbas (Karpova, 1966; 1969, p. 11 and fig. 9). Paragonite has also been reported by Stadler (1971a) from illitic, kaolinite-free tonsteins associated with meta-anthracites in ( 1) the Upper Carboniferous seam “Trois Bancs” near Grenoble, which locally carries pyrophyllite; and (2) possibly, in the Hagen 1 anthracite seam (Westphalian C) in a North Sea well off the Ems estuary. Moreover, “illites” described in some other incipient-metamorphic terranes may well be paragonite in part. In the underclays from the northeastern-Pennsylvania anthracite fields, Hosterman et a]. (1970, p. C95) noted that illite has a basal spacing of 9.8 A instead of the more usual one of 10 A. This could indicate the possible presence of paragonite, which would be commensurate with the local occurrence of pyrophyllite in the underclays. Anchizonal paragonite is widespread in the external zones of the central and western Alps: (1) in the Upper Triassic and the Lower Liassic shales and slates of the northern Glarus Alps (Frey, 1969a, 1970, 1974; see Fig. 5-18); (2) in the Upper Cretaceous calcareous slates (“marbres en plaquettes”) and the Rhaetian calcareous slates and dolomites of the internal, eastern part of the “zone brianqonnaise” (Dunoyer de Segonzac, 1969, p. 180; Abbas, 1974, pp. 23-25), and (3) in the Liassic deposits in the synclines north of the Pelvoux Massif in the French Alps (Aprahamian, 1974, pp. 8-9). Hercynian anchizonal paragonite is common in the Paleozoic deposits of the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971).

365

IT N

I

I

I

29

25

21

,

I

1

I

1

1

17

13

9

5

1

2 0"Cu Ko

Fig. 5-19. X-ray diffraction pattern of the < 2 pm fraction of a slate
Paragonitic regular mixed-layers (rectorite-allevardite) In the first of the above-mentioned Alpine anchimetamorphic areas the paragonite is almost invariably accompanied by the 6 : 4 paragonite/phengite mixed-layer with strong tendencies towards regular mixed-layering first described by Frey ( 1969b; see Fig. 5- 19) and, in many samples, by pyrophyllite (Frey, 1970, pp. 266-269 and fig. 10; 1974, p. 496). At the beginning of the anchizone, the first appearance of the mixed-layer paragonite-phengite is often accompanied by minor amounts of a regular mixed-layer mineral with a basal reflection at 22-23 A (air-dried) or 25-27 A (glycolated). It is interpreted as a 1 : 1 regular paragonite-montmorillonite or pyrophyllite-montmorillonite mixed-layer mineral similar to rectorite or allevardite (Frey, 1970, pp. 265-269). Frey believes that both this regular mica-montmorillonite mixed-layer mineral and the mixed-layer paragonitephengite represent intermediate stages in the alteration from irregular mixedlayer mica-montmorillonite to paragonite, which tends to replace entirely the paragonite-phengite mixed-layers in the greenschist facies zone.

'

'

Ludwig's (1972b) re-interpretation of Frey's (1 969a. 1970) illite-"crystallinity" data assigns the illites which are associated with the appearance of pyrophyllite and paragonite in the northern part of the Glarus Alps, to the high-grade part of the anchizone.

366 In the earlier-mentioned Jurassic black shales (“Terres Noires”) of the subalpine belt of the French Alps, Artru et al. (1969; see also Dunoyer de Segonzac, 1969, p. 160) found that allevardite appears with increasing degree of Alpine incipient metamorphism east of Gap. It is associated with rocks that already exhibit Alpine slaty cleavage and anchimetamorphic illite 10 A peak widths. Allevardite appears after the disappearance of irregular illitemontmorillonite mixed layers (see Fig. 5-17). The allevardite of the type locality at La Table (northeast of Allevard) in the Jurassic cover of the western margin of the Belledonne Massif, French Alps, is also associated with anchizonal illite lOA peak widths (Aprahamian, 1974, p. 9). Rectorite and allevardite Regular Na-rich mica-type mixed-layer minerals, described as rectorite or allevardite, are characterized by a highly regular sequence of layers. The X-ray diffraction patterns show a sharp basal (001) reflection at (1) 23.825.3 A in the air-dry, natural state, with (002) reflection at 12.2-12.5 A; (2) 22.0-22.7A in air-dry, Na- or K-saturated samples; and (3) ‘26-27A in ethylene-glycol saturated samples (see Fig. 5-20). Upon heating to temperatures between 500” and 9OO0C, the (001) spacing and the second-order (002) spacing collapse to 19.2- 19.6 A and 9.6-9.8 A, respectively. This behavior upon glycolation and heating suggests alternation of one fixed layer having basal spacing of 9.6A and an expandable layer. Rectorite-allevardite is distinguished chemically from the earlier-discussed partially ordered K-illite-montmorillonite mixed-layers by the predominance of Na over K. This difference is also reflected in the “paragonitic” 9.7-A basal spacing versus the ”muscovitic” 10.0-A basal spacing after collapse of the montmorillonite layers upon heating (cf. Sudo et al., 1962). Originally, the structure of allevardite was interpreted as pairs of mica-like layers that were linked by 0.7 (K, Ca) ions per unit cell and separated by layers of water (Caillere et al., 1950; Brindley, 1956), possibly with some exchangeable cations occupying interlayer positions (Sudo et al., 1962, p. 387). Before the detection of the alkali cations, the structure of rectorite was considered as an alternation of one pyrophyllite and one vermiculite unit (Bradley, 1950). Subsequent studies showed that rectorite and allevardite minerals were very similar, consisting of a regular or nearly regular 1 : 1 alternation of swelling montmorillonite-like (Na-montmorillonite- like or “sodiumhydronium mica”) layers with non-swelling 9.7-A mica-like layers (Brindley and Sandalaki, 1963) that are probably paragonite (Kodama, 1966; Henderson, 1970). Brown and Weir (1963) concluded that allevardite and rectorite are identical, the name rectorite having priority over allevardite which is used by most European authors.

367 Mu

Mu

I

I 1

I

7f\\ 3.07 3.24 3.33 3.52

I I

/c 4.44

>GO

I

I

\

5.00

7.0 ~

I I

t

I

I 14 O

II

t\ 23 27

A

Fig. 5-20. X-ray diffractograms of ( 2 p m fractions [untreated ( N ) , glycolated (C), and heated ( C h ) ] of an anchimetamorphic Middle Devonian slate from the Basque Pyrenees (France). A =allevardite; C =chlorite; M u =muscovite (illite); Py =pyrophyllite; Ch = chloritoid; Pa =paragonite. (After Dunoyer de Segonzac and Heddebaut, 1971, fig. 2.)

Rectorite (or allevardite) as an anchizonal index mineral In natural occurrences, rectorite (or allevardite) is commonly associated with pyrophyllite. These two minerals are persistently associated in (1) the Siluro-Devonian argillites and siltstones of the Ougarta, and of the Tindouf, McMahon and Ahnet Basins in the western Sahara (Beuf et al., 1966, p. 375; Dunoyer de Segonzac, 1969, pp. 115-1 19; Chennaux et al., 1970); (2)

368 carbonaceous shales of north-central Utah (Henderson, 1970), (3) the anchizonal Devonian deposits (but not the Ordovician and Silurian) of the Basque Pyrenees (Dunoyer de Segonzac and Heddebaut, 1971; see Fig. 5-20), and (4) the earlier-mentioned Helvetic terrane of the Glarus Alps. There is also ample evidence for the formation of rectorite (or allevardite) under conditions of incipient metamorphism even in the absence of pyrophyllite, which can usually be related to the high alkali contents. Logvinenko ( 1964) regards the occurrence of allevardite and allevardite-like minerals in phyllitic shales, from which kaolinite and irregular illitemontmorillonite mixed-layers have already disappeared, as due to the effects of incipient metamorphism. In the roof of the Bramscher Massiv, Westphalia, allevardite occurs in the Liassic shales of Vehrte, northeast of Osnabruck (“Vehrter schwarzen Kreide”), in association with corrensite (Eckhardt, 1958). The lean-coal to anthracite rank of the organic matter (Stadler, 1971b, p.488) attests to the advanced stage of diagenesis. The main occurrence of Na-rich micas in the anchizone thus appears to be in the form of regular mixed-layers, either with montmorillonite, rather than as a discrete paragonite phase. Origin of rectorite (or allevardite) The above-mentioned rectorite (or allevardite) occurrences, with the possible exception of the last one, can unequivocally be ascribed to incipient metamorphism. In the widely-studied Japanese occurrences (Sudo et al., 1962; Brindley and Sandalaki, 1963) and other areas listed by Dunoyer de Segonzac (1970, p. 309), however, rectopite is of hydrothermal origin, also usually associated with pyrophyllite. In the geothermal area of Broadlands, New Zealand, rectorite-like minerals occur at higher temperatures than the more common illite-montmorillonite mixed-layers (Eslinger and Savin, 1973a). These relationships point towards some similarity between conditions of formation during incipient metamorphism and in the hydrothermal zone (Dunoyer de Segonzac, 1970, p. 312). According to Henderson ( 1970), rectorite in the pyrophyllite-bearing carbonaceous Manning Canyon shales of north central Utah formed from paragonitic mica during late stages of diagenesis as an intermediate metastable phase, which subsequently was altered to pyrophyllite during prograde low-grade metamorphism. The mineral assemblage pyrophyllite-rectoritemuscovite found by Henderson suggests - presumably metastable - compatibility of rectorite under these conditions in alkali-deficient environments not only with muscovite (low Naf content) or pyrophyllite (K deficiency), but with both of these minerals in the presence of excess water. Rectorite (or allevardite) admittedly appears at a more advanced stage of +

369 incipient metamorphism than the earlier-mentioned regular K-micamontmorillonite mixed-layers. Nevertheless, the phase relationships found by Henderson could be analogous to Velde's ( 1969) experimental evidence that in K-poor compositions in the muscovite-pyrophyllite compositional join, pyrophyllite coexists with an ordered K-mica-montmorillonite mixedlayer between 310" and 420" at 1000 barspH20,and between 310" and 400" at 2000 bars p H Z OAbove . 300°C, the composition of Py,,, which corresponds to an ordered 30% K-montmorillonite phase, represents the maximum expandability present in a single phase in this join. At lower potassium contents (Py > 50), this phase was invariably accompanied by pyrophyllite. The regular paragonite-Na-montmorillonite mixed-layer phases (including the 1 : 1 synthetic analogue of rectorite), however, which were widely encountered in a subsequent experimental study of the reaction: 3 Na-montmorillonite

+ 2 albite

+

3 paragonite

+ 8 quartz,

were shown in the rate studies to be all transient and metastable products obtained during the transformation of the albite-Na;montmorillonite assemblage to paragonite-quartz (Chatterjee, 1973). Rectorite is apparently always less stable relative to the assemblage Na-montmorillonite-paragonite. The above reaction took place around 335°C at 2kbar p H l o , i.e., at approximately the same temperature as the reaction: kaolinite

+ quartz

+

pyrophyllite

+ H,O,

at the same p H z O which , is in good agreement with the field data. When p H , O > Ptota,, however, pyrophyllite should appear before paragonite-quartz in a prograde sequence. Appearance of miscellaneous metamorphic minerals during incipient metamorphism of pelitic rocks Appearance of chloritoid Another greenschist-facies index mineral that has been claimed to appear in anchimetamorphic roofing slates, is chloritoid (Kubler, 1967a, p. 11 1; see also Winkler, 1970, p.204). It had also been reported from the "stage of early metagenesis" of the eastern Yenisey Mountains by Kasanskiy (1967). Subsequently, Kubler (1968, p. 395) has argued that the apparent anchimetamorphic illite 10 A peak widths in obviously higher-grade metamorphic chloritoid schists may merely reflect peak broadening due to insufficient resolution of the (001) diffraction peak of muscovite at lOA from those of paragonite at 9.7 A or of pyrophyllite at 9.2 A (cf. p. 352 of this chapter), and that chloritoid only appears in the greenschist facies proper.

370 In the Quartenschiefer of eastern Switzerland, chloritoid only appears in the Urseren-Garvera zone in the quartz-albite-epidote-biotite subfacies of the greenschist facies (Frey, 1969a, pp. 101, 113-1 17; 1970, table I). Chloritoid has admittedly been reported from supposedly anchimetamorphic Paleozoic slates of the Armorican Massif, western France: from the Ch2teaulin Basin by Sagon (1965) and from Ordovician slates of other parts of the Massif by Boudier and Nicolas (1968) and Le Corre (1969, 1975). It is present, however, only in horizons very rich in both aluminium and iron. Moreover, Sagon and Dunoyer de Segonzac (1972) have argued that the 10A peak widths of most of the micas from the Chiiteaulin Basin are “epizonal”. They ascribe the low anchizonal “crystallinities” of some of the chloritoidbearing samples to peak-broadening due to the contribution of the 9.7 A peak of margarite. The lower stability limit of chloritoid is dependent on the partial pressure of oxygen (Hoschek, 1969; Ganguly, 1969; Kramm, 1973). Ganguly (1968) has argued that the curve defining the lower stability limit of chloritoid has a positive slope on a fo2-T plot. Kramm (1973) found that the chloritoid, occurring in the very low-grade greenschist facies (Mn-, Fe- and Al-rich Paleozoic slates of the southern Venn-Stavelot Massif, Ardennes), is unusually rich in Mn. Its appearance near the lower stability limit is restricted to host rocks with oxidation ratios 2 Fe,O, X 100/2 Fe,O, FeO of less than 85-90. The Mn and Mg contents of the chloritoids increase with increasing oxidation ratio. At oxidation ratios of 90 to 100, kaolinite, which is being replaced by pyrophyllite, and andalusite may be present instead of chloritoid in isometamorphic rocks. The appearance of chloritoid in the anchimetamorphic Middle Devonian black slates of the Basque Pyrenees is considered by Dunoyer de Segonzac and Heddebaut (1971) either (1) to constitute an additional marker of the anchizone under reducing conditions, or (2) to result from a thermal metamorphic event subsequent to the Hercynian anchimetamorphism. At the present-day state of knowledge, it is impossible to exclude the possibility of the local appearance of chloritoid at a metamorphic stage lower than the beginning of the greenschist facies. Although the reported occurrences of chloritoid in the anchizone appear to be restricted to rather extreme chemical rock compositions, such as very high Fe/Mg ratios, high Mn contents, and low oxidation ratios in aluminous rocks, chloritoid cannot unequivocally be regarded as an indicator of the beginning of the epizone.

+

Occurrence of stilpnornelane The occurrence of stilpnomelane in the anchizone is better established than that of chloritoid. This mineral is common not only in the glaucophane - lawsonite - schist facies, but also in the prehnite - pumpellyite -

37 1 metagraywacke facies (see references in IOsch, 1974, pp. 100-101, and p. 468, this chapter). Glauconitic horizons appear to be particularly favorable to the formation of stilpnomelane during incipient metamorphism. In Haute-Savoie, stilpnomelane appears in the Cretaceous glauconitic sandstones associated with “Gres de Taveyanne” metamorphosed in the lowestgrade part of the prehnite-pumpellyite-metagraywacke facies (Martini and Vuagnat, 1970, p. 58; Martini, 1972, p. 262). In the Helvetic zone of the Glarus Alps of eastern Switzerland, in which Frey (1969a, 1969b, 1970) earlier described the anchimetamorphism of Upper-Triassic and Lower-Liassic shales, Frey et al. (1973) have studied the effects of the progressive increase in grade of the low-grade Alpine metamorphism in the Cretaceous and lower Tertiary glauconitic horizons. The initial replacement of glauconite by stilpnomelane in the central Glarus Alps marks the beginning of zone 11, which covers the middle- and high-grade part of the anchizone (see Fig. 5-2 1). Appearance of stilpnomelane in the Glarnisch area is associated with the illite l O A peak widths of about 6.0 mm in the accompanying marly shales (Frey et al., 1973, fig. 6 ahd p. 212), i.e., well beyond the beginning of the anchizone, which is marked by illite 10A peak widths of 7.5 mm under the experimental conditions adopted. In the southern Glarus Alps in the southern part of zone 11, stilpnomelane has entirely replaced the glauconite. Further south, in the autochthonous cover of the eastern Aar Massif (e.g., in the Kistenpass area), stilpnomelane is accompanied by brown biotite and the illite 10A peak widths are largely narrower than 4 mm, i.e., “epizonal”. Similarly, Durney ( 1974) reported a mid-anchizonal illite “crystallinity” for a slate associated with the first appearance of stilpnomelane in the parautochthonous Mesozoic deposits of the Morcles area (southwestern Switzerland). The occurrence of stilpnomelane extends well beyond the high-grade limit of the anchizone. Occurrence of stilpnomelane has also been reported from quartzite-like sandstones of the Russian Platform that have reached the “stage of initial metagenesis”. Thus, in most areas studied stilpnomelane appears within the anchizone; however, locally it has been reported in lower - grade zones within the prehnite - pumpellyite metagraywacke facies (p. 468). Other minerals appearing during incipient metamorphism In the glauconite-bearing horizons of the Helvetic zone of eastern Switzerland mentioned above, the alkali - amphiboles magnesioriebeckite (“rhodusite”) and riebeckite locally appear in the anchizone (e.g., Frey, 1970). At the Urnersee and in the Glarnisch these amphiboles appear already during diagenesis, before the beginning of the anchizone. According to Frey et al. (1973, pp. 195 and 212), virtually all of these occurrences are

372

0 Glauconite 8 Glauconite

+ Stilpnomelane Stilpnomelane 0 Stilpnomelane + Biotite A Alkali-amphibole( Riebeckite) 70

Fig. 5-2 1. Distribution of lowest-grade metamorphic index minerals in glauconitic horizons of Cretaceous and Tertiary age in the Helvetic zone of the Glarus Alps, Switzerland. Zone I corresponds to part of the “diagenetic” zone and the low-grade part of the anchizone, zone I1 to the middle- and high-grade part of the anchizone, and zone 111 to the low-grade part of the greenschist facies on the basis of illite “crystallinity” (Frey, 1970). Distribution of some Alpine metamorphic minerals in other rocks in the same area is shown in Fig. 5-18. The numbers indicate K-Ar ages of glauconites in m.y. (After Frey et al., 1973, fig. 3.)

373 restricted to one particularly Na-rich horizon and are related to slickensides and joints. Coal ranks associated with anchimetamorphism The coal ranks associated with various stages of late diagenesis and incipient metamorphism have formed the subject of a number of papers by the writer. Only the main conclusions of a recent study of coal ranks in the anchizone (Kisch, 1974), however, are presented here. Exclusively anthracitic and low-rank meta-anthracitic coal and coaly matter (i.e., with less than 8% volatile matter (d.a.f.) and more than 2.5% have been found in association with anchimetamorphic illite “crystallinities”, pyrophyllite, and rectorite-allevardi te of the following anchizonal occurrences : (1) The Jurassic “Terres Noires” of the Gap-Embrun and Barcelonnette areas in the subalpine belt of the French Alps (Dunoyer de Segonzac et al., 1966; Artru et a]., 1969; see Fig. 5-17). (2)The Silurian and Devonian deposits in the well Assejmi 1, Tindouf Basin, western Sahara (Dunoyer de Segonzac, 1969, pp. 113- 119). (3) Underclays of Upper-Carboniferous coals of the Pennsylvania anthracite regions (Hosterman et al., 1960; Spackman and Moses, 1961). (4)Shales and sandstones from the roof of the Cretaceous “Bramscher Massiv” pluton near Osnabruck, Westphalia, Western Germany (Stadler, Teichmuller, and Teichmuller, unpublished, in: Teichmiiller and Teichmiiller, 1966a,b; Schreyer, 1969). ( 5 ) Tonsteins in Upper-Carboniferous coal seams in the southern Belledonne area near Grenoble, subalpine belt of the French Alps (Scherp et al., 1968; Stadler, 1971a). In some cases, anthracitic rank is reached in the higher-grade part of the “diagenetic” zone (as delimited by illite “crystallinities”, in which irregular illite-smectite mixed-layers and kaolinite are absent. Examples of such occurrences include (for details see Kisch, 1974): (1) The Jurassic “Terres Noires” in the deepest intervals of the wells Grand Luberon and Montagne de Lure (Dunoyer de Segonzac, 1969, pp. 167-170), and in the Gap area in the subalpine belt of the French Alps (Dunoyer de Segonzac et al., 1966; Artru et al., 1969). (2) The Upper Triassic Quartenschiefer at Quarten, Helvetic zone of the northernmost Glarus Alps (Frey, 1969a, 1970).

’

’

References given pertain to the descriptions of the anchizonal mineralogy. For references to the coal-rank determinations see Kisch (1974, particularly table 11).

374 (3)The Paleozoic deposits of the northeastern part of the Rheinische Schiefergebirge where the coaly matter, which is associated with anchimetamorphic illite “crystallinities” and pyrophyllite (Weber, 1972b; also Scherp et al., 1968), is exclusively high-rank anthracite, with Ro,,(vlt,,nlte) of 4%, and meta-anthracite (Wolf, 1972). (4) Jurassic cover of the southern termination of the “zone houillere Briangonnaise”, the Carboniferous coals of which are even higher-rank anthracites (e.g., Robert, 1971); both anchizonal illite “crystallinities” (Abbas, 1974) and high-rank anthracite (Chateauneuf et al., 1973) have been reported here. Thus, one can conclude that the appearance of anchimetamorphic illite “crystallinities”, pyrophyllite, and rectorite (or allevardite) takes place within the anthracite range of coal ranks. Conversely, the attainment of anthracite rank is associated with lower than anchimetamorphic illite “crystallinity indices” and, in many cases, with the occurrence of kaolinite ( e g , in underclays and tonsteins). Identification of the anchizone: a summary Although the anchizone is mainly defined by a range of illite “crystallinity indices”, it shows a number of additional characteristics that allow its recognition, even where i t might otherwise be overlooked in the absence of illite lOA peak widths or in the case of anomalous “crystallinity” data. Some distinctive minerals, including pyrophyllite, paragonite and, probably, stilpnomelane, are common to both the anchizone and the “epizone”. Their occurrence only indicates that the low-grade boundary of the anchizone has been exceeded. The former two minerals are particularly important in view of the tendency towards abnormally low illite “crystallinities” (i.e., 10 A peak broadening) in rocks containing these minerals. The high-grade limit of the anchizone in pelitic rocks is somewhat more difficult to establish in the absence of illite “crystallinity” data. I t can be considered to have been exceeded upon appearance of typical greenschistfacies minerals, such as biotite, andalusite, and, probably, also margarite and Mg-rich chloritoid. Negative characteristics for attainment of the anchizone must be used with caution. Inasmuch as absence of irregular illite-smectite mixed-layers and of kaolinite could be due to sedimentary control, it can be considered to be due to postdepositional processes only if it can be argued that the rocks were originally rich in smectite or kaolinite content (e.g., bentonites, underclays, and tonsteins). Anthracite and low meta-anthracite rank may also be regarded as primafacie indication for anchizonal environment, although there is an overlap

375 with both the “diagenetic” and the “epizonal” realms sensu Kubler. Finally, appearance of slaty cleavage may also serve as an indication that the anchizone has been reached, despite the possibility of its local development at lower-grade conditions. Towards a scheme of lowest-grade mineral facies for clastic rocks

As described in this chapter sequential changes in the mineralogy of clastic rocks upon deep burial can be used to define grades or stages of burial and incipient metamorphism in clastic sedimentary sequences. Establishment of generally applicable stages of this kind is likely to contribute greatly to various aspects of the study and understanding of sedimentary basins (e.g., in the exploration for fossil fuels and the reconstruction of burial history and tectonic development) and of marginal orogenic zones. Consequently, correlations must be established between the several sequential changes in mineralogy discussed above: ( 1) the decreasing content of swelling layers and the appearance of stacking order in illite-smectite mixed-layers, (2) the polytypes and ultimate disappearance of kaolinite group minerals, and (3) coal rank. For deeper levels and more advanced stages correlations must be established between illite “crystallinity”, polytypes of illite and chlorite, and the appearance of pyrophyllite and rectoriteallevardite. The changes in a given sedimentary mineral after burial, however, are controlled not only by temperature and pressure, but also to a major extent, by the chemical environment. The latter is a function of lithologic variables, such as the mineralogy of the clastic constituents, the composition of pore solutions, and the ability of pore solutions to circulate (i.e., permeability). The chemical factors affecting the stage of burial at which a particular discontinuous mineralogic change begins or is completed, or a continuous change attains a certain defined stage, are so numerous that no uniform relationships among various processes can be expected except by specifying reductions in the number of lithological-chemical variables. Thus, studies of the relationships among specific mineral-replacement processes should be restricted to a small number of well-defined ranges of lithology, e.g., siltstones rich in unstable K-rich detritus and minimal carbonate cement content, which could adversely affect permeability, or K-, Fe- and Mg-poor claystones having low permeability. An attempt in this direction was made by Kossovskaya and Shutov (1963) by developing a scheme for “regional epigenesis and metagenesis”, which allows for major differences in primary lithology by stating diagnostic mineral assemblages (“facies”) in each “stage” for each of five compositional “rock families”. An amplified and modified version of their scheme is

w

TABLE 5-V

4 o\

Schematic mineral distribution during the "stages" of regional epigenesis and metagenesis in clastic sedimentary rocks after Kossovskaya and Shutov (1961, 1963) and the approximately equivalent illite-"crystallinity" zones of Kubler (l967a, 1968) and their correlation with mineral distribution and metamorphic mineral facies (after Coombs et al., 1959 and Coombs, 1960, 1961) in intermediate volcanic graywackes and tuffs Schematic. mineral distribution'.

"stages" Of

cplgenesls and metagenesis

and "facies"

of regional eptgenesis and metagenests in d i f f e r e n t "families " of clastic rocks2 rock54

-

Quartz kaolinltc

Mineral zones of KUBLER (1967. 1968 i 3 major CharacteriSticS

Schematic mineral d i s t r i b ~ t i o n ~ and . metamorphic mineral facies after COOMBS (1960,1961)5, In Intermedmte volcanic graywackes and tufts

"L~thoclastlc sandstones" and "acid arkoses"

HEULANDITE-ANALCIME ZONE of t h e ZEOLITE FACIES (Coombs et al ,1959)

INITIAL EPlGENESlS

"heulandtte -analcimc l a c ~ 5 " (Kossovskaya and Shutov 1961.1963) "DIAGENETIC" ZONE i l l i t e crystallinities '7.5 [>4 irreaular illite/montmortllonite m-l kaoltnitc-group minerals

01

DEEP EPlGENESlS

EARLY METAGENESI!

Absent' pyrophyllite, rectorite, paragontte

ANCHMETAMORPHIC ZONE illite c r y s t a l l i n i t y index 7.5- 4.0 [ 4 0 - 2 . 5 1 pyrophyllite, r e c t o r i t e paragonite. phengitc

LAUMONTITE ZONE of the ZEOLITE FACIES (Coombs ct a1 ,1959) "laurnontlte facles" ( KOSSoVskaya and Shutov ,1961,1963)

PREHNITE-PUMPELLYITTE-METAGRAYWACKE FACIES (Coombs et al , 1 9 6 0 ) "prehnitc-muscovite-chlorite tacies" (Kossovskaya and Shutov. 1963)

Absent. irregular Illlte/

LATE METAGENESI!

Low-grade part of the GREENSCHIST FACIES "

E PI ZONE "

illite c r y s t a l l m t y index <4.0 [
51

-

"epidote -muscov!te-chlorite stilpnomclane faccles" (Kossovskaya and Shutov, 1963)

Mineral distribution after Kossovskaya and Shutov (1963, 1970), with additions and modifications from Karpova (1969), Dunoyer de Segonzac (1969), and Kisch (1969, 1974). Stippled extensions of mineral distributions downwards indicate common persistence of minerals as metastable relics. “Facies” of regional epigenesis and metagenesis and “families” of clastic rocks after Kossovskaya and Shutov (1963, 1970). Rock families: “Quartz-kaolinite rocks”: monomineralic quartz sandstones in association with kaolinite clays. “Acid arkoses”: acidic arkoses or quartz-feldspar sandstones with plagioclase up to An,, and/or K-feldspars, associated with hydromicamontmorillonite and hydromica clays. “Lithoclastic sandstones”: polymict lithoclastic sandstones; detritus-mainly low-grade metamorphic schists, siliceous rocks, etc. The associated clays have polymineralic composition with a consistent presence of hydromica, montmorillonite, and chlorite. Correlation of mineral zones of Kubler with the “stages” of regional epigenesis and metagenesis after Kubler (1967a, 1968) and Kiscb (1974, 1975). Illite “crystallinity indices” of Kubler (1968, fig. 3; Dunoyer de Segonzac et al., 1968, fig. 3) were also used by Frey (1969a, 1970). The illite “crystallinity indices” of Kubler (1967a), which were also used by Artru et al. (1969), are presented in square brackets. Schematic mineral distribution for an intermediate-pressure type terrane (see this chapter). Extension of some minerals (e.g., heulandite and laumontite) into deeper zones in silicic tuffs is not indicated. Heulandite and analcime could persist into the laumontite zone in very impermeable, “sealed” beds; analcime is also present in quartz-free rocks. Clastic medium and calcic plagioclase may persist in the prehnite-pumpellyite facies in carbonate-rich rocks. “Facies” designations of Kossovskaya and Shutov (1963) are also given. The pumpellyite-actinolite-schist facies (Hashimoto, 1966) is not indicated separately; it probably occupies the lowest-grade part of the “low-grade part of the greenschist facies”.

w

4

4

378

presented as Table 5-V. Data are lacking, however, for several diagnostic mineralogic features, such as the extent of illite-smectite mixed-layering, the development of stacking order in mixed-layers, the appearance of rectoriteallevardite and corrensite, and illite “crystallinity”. Moreover, some of the correlations presented by Kossovskaya and Shutov between assemblages in different “rock families”, particularly with the zeolite and prehnitepumpellyite facies assemblages of “intermediate arkoses” and “vulcanogenic graywackes”, are based on very meager data and appear to be imprecise (cf. Kisch, 1974). Another way of reducing the number of variables is to correlate stages of mineral alteration in different lithologies against a continuously-varying parameter of burial-metamorphism, which is either independent of the chemical environment, such as coal rank, or is believed to reflect physical burial conditions when certain extreme chemical environments are excluded, e.g., illite “crystallinity” excluding ( 1) K-poor lithologies and environments, and (2) magnesian (phengitic) illites. Subsequently, one should try to correlate the differences in relative depth or temperature, or even of the order of occurrence of various specific mineral changes, with differences in lithology. Such correlations are few, and much more work on the subject will be required in order to arrive at a firm mineral-facies scheme of burial for clastic sedimentary rocks. As more results become available, no one scheme of burial-metamorphism and incipient-metamorphism may be universally applicable. With further study, the order of disappearance or appearance of diagnostic minerals may prove to be different in different areas and geological settings (such as P/T regimes), even in lithologically similar sequences and under chemically identical environments. These differences might reflect the influence of different pressuretemperature gradients of burial in various geological situations and, thus, be analogous to the metamorphic facies series of higher-grade regional metamorphism. Such facies series have been recognized in lowest-grade metamorphic volcanic graywackes and tuffs. This is the subject of pp. 441-454 in the following sections.

LOWEST-GRADE METAMORPHIC MINERAL LITHIC-FELDSPATHIC SEDIMENTS

FACIES

IN

VOLCANIC A N D

Many sequences of pyroclastics and volcanic clastic sediments are characterized by the widespread development of zeolites and other hydrous Ca-Al-silicates upon deep burial. Particularly common are Ca-zeolites (heulandite and laumontite), alkali-zeolites (analcime, clinoptilolite and

379 mordenite), and the non-zeolitic Ca-Al-silicates prehnite and pumpellvite. which develop either at the expense of volcanic glass or calcic plagioclase. or as cements of the clastic framework. Such minerals have given rise to the recognition in the late fifties and the sixties of burial-metamorphic facies: (1) zeolite (Turner. in Fyfe et a].. 1958. p. 216; Coombs et al., 1959), (2) prehnite-pumpellyite-metagraywacke (Coombs, 1960), and (3) pumpellyite-actinolite-schist facies (Hashinioto. 1966). The characteristic mineral assemblages of these facies. their progressive development upon deep burial, and the chemical controls of their formation are presented here.

Recognition of lowest-grade rnetarnorpl?isrn In classical metamorphic petrology relatively little attention was paid to lowest-grade metamorphism. The use of the petrographic microscope as ;I main tool restricted study and choice of metamorphic index minerals to those which are coarse and distinctive enough for optical identification i n thin section. The lowest-grade metamorphic facies conventionally recognized, i.e., the greenschist facies and the glaucophane-schist facies, were recognized by the appearance of such diagnostic minerals as actinolite. epidote, or glaucophane in basic rocks. and of biotite or chloritoid i n argillaceous rocks. It was realized by some authors that ( 1 ) the chloritemuscovite assemblage of the lowest-grade zone or subfacies of the greenschist facies also occurred in many “non-metamorphic” sedimentary rocks, and (2) that the formation of slaty cleavage, which was regarded by some as heralding the onset of metamorphism, often took place within the chlorite zone ( e g , Hutton and Turner, 1936, in New Zealand). Although the occurrence of diagenetic zeolites in sedimentary rocks had previously been recognized by various authors (see references in Coonibs, 196 I), the characterization of mineral assemblages and the recognition of mineral modification in lowest-grade metamorphic rocks remained elusive until the general introduction of X-ray powder diffraction. As a complement to the polarization microscope, this technique made the mineralogy of lowest-grade metamorphic rocks accessible for detailed study. Beginning in the early sixties, there has been an upsurge in the attention paid to mineral assemblages in lowest-grade metamorphic rocks. These studies, as far as carried out by metamorphic petrologists, have mainly been concentrated on volcanic rocks and volcanic clastic sediments. Inasmuch as rocks and sediments have a high content of constituents that are very unstable and reactive under burial, such as volcanic glass and calcic plagioclase, they are among the first to develop distinctive post-depositional mineral assemblages. These mineral assemblages, however, commonly show

380 qualitative differences with conventional metamorphic facies that are important enough to warrant mention at the outset: (1) Burial-metamorphic assemblages are commonly developed only in a restricted range of rock types. Their extensive development as a rule requires presence of acid and intermediate pyroclastic rocks, or clastic sediments rich in unstable, generally volcanic detritus. Associated, less reactive clastic sedimentary rocks such as argillaceous shales, quartzose arenites, or arkoses, may lack any obvious effects of burial metamorphism. (2) Burial metamorphism is commonly achieved-at least in its lower grade, i.e., in the zeolite facies-without the aid of penetrative deformation. Zeolitic assemblages commonly form without the accompanying development of typical penetrative metamorphic fabrics of higher-grade metamorphism in sedimentary rocks: most zeolite-facies rocks entirely lack schistosity. (3) Mineral-equilibrium reactions at the low temperatures prevailing during burial metamorphism proceed at a slow rate, and the recrystallization of the rocks, particularly of their clastic framework, may be very incomplete. As a consequence, evidence of disequilibrium abounds in burial-metamorphic rocks: the establishment of the low-grade mineral assemblages does not of necessity involve the whole rock, but may be restricted to particularly reactive clastic constituents, such as volcanic glass or calcic plagioclase and their immediate environment, to vesicles in lava, or to the cement of the rocks. Establishment of mere local equilibrium is contrary to the notion of metamorphic facies (e.g., Fyfe and Turner, 1966), in which it is implicit that the mineralogic assemblage within a facies is a function of the bulk chemical composition only. In view of the important role played by the mineralogical composition of the rocks in burial metamorphism, Coombs et al. ( 1959, p. 54) recommended to consider zeolitic rocks in terms of Eskola’s ( 1920) mineral facies rather than metamorphic facies. (4) The burial-metamorphic reactions appear to be favored by high porosity and permeability of the rock: in many sequences i t has been noted that lavas are generally much less altered than are the associated, more porous pyroclastics, and that zeolitization is often more advanced in coarser than in finer tuffs and volcanic clastic rocks. Thus, in addition to control by mineralogic composition, there is also a control by rock texture. In the following sections, the mineralogic changes upon burial metamorphism, their zoning, mineral facies, and chemical and physical controls are described. Finally, the rank of the associated coaly matter and the clay mineralogy of the associated non-zeolitic clastic sediments are discussed, in an attempt to arrive at a correlation with the diagenetic changes described earlier.

38 1 Recognition of the zeolite facies

Zonation with depth in the occurrence of diagnostic zeolites and other Ca-Al-silicates was first described from a very thick stratigraphic section (more than 10 km) of Triassic eugeosynclinal volcanic graywackes, tuffs. and siltstones at Taringatura, Southland, New Zealand, by Coombs ( 1954) (see Fig. 5-22). In the upper half of the section. the Ca-zeolite heulandite or its Na-Si-rich variety clinoptilolite, and, less commonly, analcime quartz are characteristic alteration products of intermediate to acid volcanic glass. Locally, they occur as interstitial cements and replacement products of detrital calcic plagioclase. Smectite, quartz, chlorite, chalcedony. and sphene are associated newly-formed minerals. In the upper part of this heulandite-analcime zone, detrital minerals-including calcic plagioclaseare generally not strongly altered. The alteration of calcic plagioclase increases downwards in the lower part of the heulandite-analcime zone. Below a depth of about 5000 m, in the Taringatura section, analcime quartz is replaced by albite, whereas heulandite is replaced by the less hydrous Ca-zeolite laumontite. Calcic plagioclase is largely replaced by albite and laumontite, although relic andesine persists locally. In the lower 3000 m of the section, the volcanic graywackes are thoroughly reconstituted, common mineral assemblages being ( 1) quartz-albite-chlorite-laumontitesphene and (2) quartz-albite-chlorite-prehnite-calcite-sphene. On the basis of these data, Turner (in Fyfe et al.. 1958) defined a zeolite facies of metamorphism as being characterized by the mineral assemblage laumontite-quartz, assigning the heulandite-quartz and analcime-quartz assemblages of the upper half of the Taringatura section to the zone of diagenesis. A similar type of zonation, with an upper zone characterized by clinoptilolite-heulandite and/or analcime, and a deeper, higher-grade zone by laumontite, has since been recognized in clastic sequences of several other

+

+

Plogioclose

No%

0%

PurnDe.1~ ite Epidote -Pr_en;ite

1 - - - - _

LoJrnontite Heulondite in mperrneoble roc1(s AnolclrneI PsetidpmorpPic 1 25 000 20 000 30 000

- - - - - - - - - in c o a r s e - g r o ~ dsondtones -

-- - - - - - - -

A

15 000

i

10 000

'res?

5 000

-

,0

Present stmtigrophic depth (feet 1 Delow n gnest be05 exposed

Fig. 5-22. Burial metamorphic mineralogy in tuffs and graywackes of the Triassic sequence of Taringatura, Southland. New Zealand (slightly modified after Coornbs. 1954. fig. 2. and Coombs et al., 1959, fig. 2).

382

Purnpellvite Prehnite Laurnontit e

----- ------Htulandre Analcime

__

-A. Triassic New Zealand (Coombs et al.,1959)

B. Cretaceous Puerto Rico (Otalora ,1964)

Lalrmontite Analcime

-

Albitized plagioclase

7

content by extrapolation

.................................

- -- ............................. C. Late Paleozoic New South Wales (Packham& Crook 1960) .

.

D. Triassic Buttle Lake, British Columbia ( Surdam ,1967)

itized Dlaaioclase , I

E. Miocene Tanzawa Mtns, Japan (Seki et al ,,1969)

I

0

I

5

I

I

I

I

10

15

20

25

3

Thousands of feet below top of measured section

Fig. 5-23. Vertical zonation of zeolites, prehnite, pumpellyite, and albitized plagioclase in burial sequences of lowest-grade clastic rocks and volcanics (from Liou, 1971a, fig. 9).

383 areas. These include (1) the Late Paleozoic sequence of New England, New South Wales (Packham and Crook, 1960; Whetten, 1965); (2) the Permian to Cretaceous sequence of the Lena Basin in Yakutia, Siberia (Zaporozhtseva et al., 1961; Zaporozhtseva et al., 1963); (3) the Cretaceous of east-central Puerte Rico (Otalora, 1974); (4) the Miocene “Green Tuff” Formation of the Shinjo Basin, northern Honshii (Utada, 1965; Ijima and Utada, 1966; see Fig. 5-31) and of the Tanzawa Mountains in the southern “Fossa Magna” of central Honshii (Yoshitani, 1965; Seki et al., 1969a; Shimazu et al., 1971; Seki, 1971); ( 5 ) the Neogene tuffs of the Niigata area, northern Honshu (Ijima and Utada, 1971); (6) Cretaceous volcanites of coastal Angola (Portugal Ferreira et al., 1971); and (7) the Late Cretaceous sandstones of Vancouver Island and the adjoining Gulf Islands, British Columbia (Stewart and Page, 1974). Vertical sequences in some of these areas are given in Fig. 5-23. In many of these areas the boundary between the upper and the lower zone of the zeolite facies is reasonably sharp, but other zeolitic sequences show appreciable overlap between the range of occurrence of laumontite on one hand and either heulandite or analcime (or both) on the other. Such overlap is found in the Taringatura section itself (cf. Boles and Coombs, 1975), in the Late Cretaceous pyroclastics and livas of east-central Puerto Rico (Otalora, 1964; Jolly, 1970), in the Sustut Group of south-central British Columbia (Read and Eisbacher, 1974), and in the ”Green Tuff” of the Tanzawa Mountains (Seki et al., 1969a, tableI; Shimazu et al., 1971, table 2), where analcime even persists throughout the laumontite zone until the appearance of prehnite and pumpellyite. In the Triassic Karmutsen Group of Vancouver Island both analcime and laumontite extend well into the prehnite-pumpellyite-bearing zone. Clinoptilolite-mordenite and heulandite-analcime zones Apart from the above described major mineral zoning within the zeolite facies, a zeolite mineral zoning is present within the lower-grade zone in some Cenozoic pyroclastic sequences in Japan and the western United States. These include the Miocene pyroclastics (the “Green Tuff Formation”) of the Shinjo Basin and the Niigata Oil Field of northern Honshu (Utada, 1965: Ijima and Utada, 1966, 1971, 1972; see also Mizutani, 1967; Miyashiro and Shido, 1970; cf. Fig. 5-31), and other areas in central and southern Honshii (e.g., Yoshitani, 1965; Seki et al., 1969a; Utada and Minato, 1972; see also Miyashiro and Shido, 1970), and the Upper Tertiary tuffs and volcanic clastic rocks of southwestern Nevada (Moiola, 1970). These and similar sequences show that the siliceous alkali-zeolites clinoptilolite and/or mordenite ( = ptilolite), commonly accompanied by cristobalite or opal, occur in higher zones than the less siliceous analcime and

384 heulandite. Laumontite appears only in the lower part of the heulandite zone. In some active hydrothermal areas, for instance, in the wells of Onikobe, northern Honshu (Seki et al., 1969b), mordenite occurs immediately above the laumontite zone, without an intervening heulandite zone (Fig. 5-24). The existence of discrete mordenite-cristobalite and heulandite-analcime zones appears to be related to chemical factors that are discussed in a later section of this chapter, rather than to the increase of temperature and pressure upon burial. Temp. (OC)

d

100 I

150 I

200 I

Mordmite Yugawaralite Laumontite Waimkitc 100

200

E

Y

c S

8

300

a

400

500

Fig. 5-24. Geothermal gradients and zeolite distributions in four drill holes in the Katayama geothermal area, Onikobe, northern Honshu, Japan (from Seki et al., 1969b, fig. 4).

385 Occurrence of a wairakite zone Similar zeolite-facies mineral zoning has been found in several active geothermal areas, e g , Wairakei, New Zealand (Coombs et al., 1959; Steiner, 1968), and Onikobe, northern Honshti (Seki, 1966; Seki et al., 1969b: see also Liou, 1970). In addition to the zeolites mentioned above, wairakite, the Ca-analogue of analcime, appears here in the deeper part of the laumontite zone, generally persisting to much greater depths and at much higher temperatures. The mineral zoning of the Onikobe geothermal area is given in Fig. 5-24 (after Seki et al., 1969b). Wairakite and wairakite-analcime solid solutions, however, are also common in some low-grade sequences of more regional extent as replacements of plagioclase or volcanic glass, for instance, in the high-grade part of the laumontite zone and in prehnite-pumpellyite zone of the metamorphosed Miocene lavas and pyroclastics (“Green Tuffs”) in the southern Fossa Magna area of central HonshB, e.g., in the Tanzawa Mountains (Seki et al., 1969a; Shimazu et al., 1971) and in the Fujikawa area (Seki, 1971). Wairakite also occurs in laumontite-bearing lavas in the Karmutsen Group of Vancouver Island, British Columbia (Surdam, 1973), persisting downwards into the prehnite-pumpellyite zone to greater burial depth than does laumontite (cf., Fig. 5-23). Inasmuch as the metamorphism of the wairakite-bearing “Green Tuff” sequences in the southern Fossa Magna area and elsewhere in Japan is invariably related genetically to the thermal action of intrusive quartz diorite and diorite bodies, it appears that the regional occurrence of wairikite is restricted to areas having comparatively high geothermal gradients (e.g., Liou, 1970; Seki, 1971, 1973b). Its presence is, therefore, of considerable importance in the definition of facies series of lowest-grade metamorphic sequences. In view of its persistence into the higher-grade prehnite-pumpellyitemetagraywacke facies, however, wairakite is not by itself a diagnostic zeolite for purposes of definition of the zeolite facies, as long as it is not accompanied by heulandite, laumontite, or pumpellyite. Pumpellyite facies The prehnite-pumpellyite-metagraywacke facies In the lower part of the Taringatura section below a depth of about 17,000 ft, pumpellyite and, subsequently, prehnite begin to appear together with, or at the expense of laumontite (cf. Fig. 5-22, after Coombs, 1954; Fig. 5-26, after Coombs, 1961). Moreover, as first noted by Hutton (1937) and subsequently investigated extensively by Coombs et al. ( 1959), these minerals occur on a regional scale in the lower-grade of the four textural subzones

386 Chl. 1 to Chl. 4 in the partly-reconstituted graywackes (“semi-schists”) of the chlorite zone of the Otago schists (or Haast schists), in the southern part of the New Zealand geosyncline. These subzones-defined on textural grounds by Hutton and Turner (1936) and Hutton (1940)-reflect the increasing intensity of deformation towards the central greenschist-facies area of the Otago schists, both from the zeolite-facies areas of Southland in the south and southwest, and from Canterbury in the northeast (see Fig. 5-25, after Landis and Bishop, 1972). Prehnite and pumpellyite are the predominant Ca-Al-silicate alteration products, with orthoclase still being a common detrital mineral in the metagraywackes (“semi-schists”) of the Chl. 1 subzone, whereas pumpellyite, actinolite and stilpnomelane-but apparently not prehnite-are predominant in those of the Chl. 2 subzone and locally of the Chl. 1 subzone (see also Landis and Coombs, 1967; Bishop, 197213). Pumpellyite is still abundant in the Chi. 3 subzone of eastern Otago, but it is absent in the Chi. 4 subzone. The latter contains abundant Fe-poor epidote and is fully reconstituted to greenschist-facies mineral assemblages (Brown, 1967; see also Turner, 1968, pp. 270-,74). In 1959, Coombs et al. suggested that it might be convenient to accommodate the pumpellyite-bearing rocks of subzones Chl. 1 and 2, which were formerly included by Turner (in: Fyfe et al., 1958) in the quartz-albite-muscovite-chlorite subfacies of the greenschist facies, in a new prehnite-pumpellyite facies of the greenschist facies. Packham and Crook (1960, p. 403) reported the appearance of prehnite, and-some 1200 m lower in the section-of pumpellyite and epidote, in the Carboniferous Parry Group of the Tamworth Trough, New South Wales. They tentatively assigned the lower, laumontite-free part of the Parry Group, in which the detrital plagioclase is albitized, to a prehnite-pumpellyite facies of “epigenetic diagenesis”. Almost simultaneously, Coombs (1960, p. 342) formally defined a prehnite-pumpellyite-metagraywacke facies “to include tho>\: assemblages produced under physical conditions under which the followin;: are commonly formed: quartz-prehnite-chlorite or quartz-albite-pumpellyite-chlorite, without zeolites and without the characteristic minerals of the glaucophane schist facies, jadeite and lawsonite”. Inasmuch as the pelitic greenschist-facies assemblage quartz-albitemuscovite-chlorite is stable also at least in the higher-grade part of the prehnite-pumpellyite-metagraywacke facies, it is not diagnostic. Consequently, Coombs (1960, p. 342) re-defined the onset of the greenschist facies by the critical assemblage quartz- albite- muscovite- chlorite- epidote, without pumpellyite, lawsonite, or prehnite, such as are found in subzones Chl. 3 and 4 of the Otago schists (Coombs, 1960; Brown, 1967; Bishop, 1972b; see also Turner, 1968, pp. 270-274).

387

1

0

1

1

1 0 0 Miles 1

100 Km

Fig. 5-25. Metamorphic zonation in the Wakatipu metamorphic belt (New Zealand geosyncline). z =zeolite facies; p p = prehnite-pumpellyite facies; p u = pumpellyite-actinoli te facies; ez=chlorite zone of greenschist facies; hz= biotite zone of greenschist facies; u = amphibolitc and epidote-amphibolite facies; h =hornfels and some migmatitic rocks; large dots = lawsonite-albite-chlorite and blueschist facies (including narrow flanking zones of prehni tcpumpellyite, pumpellyite-actinolite, and zeolite facies); A F= Alpine Fault; MTL = Median Tectonic Line; T=Taringatura Hills; H = Hokonui Hills (after Landis and Bishop, 1972, fig. 4).

388 The prehnite-pumpellyite-metagraywacke facies has subsequently been recognized in many areas, often with comparatively well-defined boundaries or relatively narrow transition zones from the laumontite zone, e.g., in the east-central and south-central areas of Puerto Rico (Otalora, 1964; Jolly, 1970). the eastern Akaishi Mountains of central Honshii (Matsuda and Kuriyagawa, 1965; see also Hashimoto, 1966, fig. 5 ) , and the Sanbagawa terrane of the central Kii Peninsula, southern Honshu (Seki et al., 1971). In many areas, however, there is a distinct telescoping of the laumontiteand pumpellyite-bearing zones, with transition zones of appreciable width: ( 1 ) in Southland, New Zealand itself (Coombs, 1954; Coombs et a]., 1959; Landis, 1974; see Fig. 5-22), (2) the k a m a district, eastern New South Wales (Raam, 1968, p.329), (3) the Aldrich Mountains, Oregon (Brown, 1961: Brown and Thayer, 1963), (4) the Tanzawa Mountains, central Honshn (Seki et al., 1969a; Shimazu et al., 1961), and ( 5 ) Vancouver Island, British Columbia (Surdam, 1973). In these areas, laumontite and pumpellyite tend to occur in different rock types. In other areas, however, e.g., in the “Gres de Taveyanne” (Taveyanne Graywacke) in the Helvetic zone of central Switzerland (Martini and Vuagnat, 1970, p. 56), they have been reported to occur together. Coombs (1960; 1961, fig. 4C; see Fig. 5-26) regards this transitional laumontite-prehnite-pumpellyite zone as part of the laumontite zone of the zeolite facies, in accordance with the requirement that the prehnitepumpellyite-metagraywacke facies must be free of zeolites. This usage is followed by the present writer, with the exception of laumontite restricted to silicic vitric tuffs within prehnite-pumpellyite sequences, as discussed later. Pumpellyite has been reported locally to occur in sequences that have apparently been subjected to even lower-grade zeolite-facies metamorphism, for instance, in the Kyeburn Formation of central Otago, New Zealand. In this formation, pumpellyite-bearing sandstones are associated with biotiteand oligoclase-bearing tuffs, the groundmass of which is largely altered to niontmorillonite or heulandite-clinoptilolite, and coals that have only reached subbituminous rank with 43-44% volatile matt& (Bishop, 1972a). T h e pumpelbite-actinolite-schist facies In their suggestion of a prehnite-pumpellyite zone, Coombs et al. (1959, p. 68) noted the regional occurrence of actinolite, stilpnomelane, and epidote in typical pumpellyitic semi-schists of the Chl. 2 subzone of Otago. They suggested that the appearance of actinolite and stilpnomelane, and the disappearance of prehnite represent possible isograd markers within the prehnite-pumpellyite zone. In his formal proposal of the prehnitepumpellyite-metagraywacke facies, Coombs ( 1960) drew attention to some additional regional occurrences of pumpellyite with actinolite, but without prehnite, e.g., in the Sanbagawa metamorphic belt of Honshu (Seki, 1958,

A'2°3

Kaolinite Montmorilloniter

A

(Fe, Mg)O

Calcite

( Fe, Mg )O

Colcite

A1203

A'2°3

C

Calcite

ont mor il lonitcs

B

Montmoril loniter

r

(Fe.MgI0

D

Calcite

(Dolomite) Actlnolite

(Fe,Mg)O

Fig. 5-26. Schematic AI20,-Ca0-(Fe. Mg)O diagrams showing some observed mineral assemblages ( X ) in low-grade rocks from New Zealand containing excess quartz. sodic and potassic phases. A. Heulandite zone. B. Laurnontite zone. c. Laumontite zone, high-grade part (pumpellvite present). D. Higher-grade part of the prehnite-purnpellyite-metagraywacke facies (or pumpellyite-actinolite-schist facies): stilpnomelane rnav also be present. (From Coombs. 1961.)

390 1961) and on Celebes (De Roever, 1947) (see below, pp. 391-392). Though including these pumpellyite-actinolite zones in the prehnitepumpellyite-metagraywacke facies, Coombs ( 1960, pp. 341-342) considered that in New Zealand two zones or subfacies were distinguishable: “(a) A lower grade quartz-prehnite zone in which some combination of the following is found: quartz, albite, prehnite, pumpellyite, chlorite, calcite, sphene, orthoclase, muscovite. (b) A higher grade zone with some combination of quartz, albite, chlorite, sphene, actinolite, muscovite, calcite, stilpnomelane, pumpellyite, and epidote.” Subsequently, a pumpellyite-actinolite-schist facies was proposed by Hashimoto ( 1966) to accommodate the generally prehnite-free pumpellyite-, actinolite-, and commonly epidote-bearing zones (cf. Seki, 1972). These occur between the prehnite-pumpellyite zone and the pumpellyite-free greenschist facies in various areas in Japan, such as zone 111 in the central Kii Peninsula, southern Honshii (cf. Seki et al., 1971), referred to earlier. The petrography of the pumpellyite-actinolite schist zone succeeding the prehnite-pumpellyite zone in the Wakatipu metamorphic belt (Landis and Coombs, 1967) of southern New Zealand has since been exhaustively described by Bishop (1972b) from the northern margin of the Otago Schist (or Haast Schist) belt, where it is about 9 km thick. The distribution of minerals in this metamorphic sequence is given in Fig. 5-27, as presented by Bishop ( 1972b). Successive appearance of prehnite-pumpellyite and a prehnite-free pumpellyite-actinolite zones was found in the Greenlaw-Mooseleuk Lake area of northern Maine by Coombs et al. (1970), who recognized the equivalence of Hashimoto’s pumpellyite-actinolite schist facies to the hlgher-grade zone or subfacies of the prehnite-pumpellyite-metagraywacke facies in New Zealand earlier suggested by Coombs ( 1960). This and some other areas on the west side of the Appalachian metamorphic belt locally show the transitional assemblage of prehnite + actinolite, often unaccompanied by pumpellyite. In such cases, particular attention needs to be paid to the mode of occurrence of the actinolite. If it occurs only 3s a replacement of igneous hornblende or pyroxene (cf. Zen, 1974a, p. 213), such actinolite may not be part of the metamorphic assemblage, and thus does not define an actinolite isograd (cf. Hashimoto, 1972, p.25). On the other hand, Schermerhorn ( 1975), noting the persistence of irregular prehnite ( W m e n t l y not associated with pumpellyite) throughout the pumpellyiteLlctinolite zone of the Hercynian terrane of the Iberian pyrite belt almost up to the pumpellyite-out isograd, proposed to regard the prehnite-pumpellyite and punippellyite-actinolite zones as the subfacies of single pumpellyite facies as originally proposed (De Roever, 1947, 1950, see below).

39 1

Authigenic

-

Prehnite-pumpellyite facies

Pumpellyite actinolite facies

Greenschist facies

Albite Prehnite Purnpellyite Act in01ite Fe ep idot e Clinozoisite Chlorite Muscovite St ilpnornelant

-

Detrital Plagioclasc K-feldspar Muscovite Biotile Garnet Epidote Amphibole

- - -_--- -

c- - - - - _ -

I

--_-_-_

I

----__

I

Fig. 5-27. Distribution of minerals with respect to metamorphic grade in metagraywackes of the Dansey Pass area, Otago schist belt. southern New Zealand geosyncline (after Bishop. 1972, fig. 5). The actinolite in the pumpellyite-actinolite facies zone occurs mainly as fringing needles on detrital amphibole.

In contrast, both prehnite and pumpellyite disappear upon the appearance of actinolite in the Tamworth Trough and the Molong Geanticline of New South Wales (Packham and Crook, 1960; Smith, 1969), and in the Tanzawa and Akaishi Mountains of southeastern central Honsha (Seki et al., 1969a; Matsuda and Kuriyagawa, 1965, see also Hashimoto, 1966. fig. 5 ) . Thus, no pumpellyite - actinolite zone intervenes here between the prehnite pumpellyite zone and the greenschist facies. The succession of the Tanzawa Mountains has formed under the influence of a large quartz-diorite intrusion (see also Seki, 1971, 1973a; Iijima and Utada, 1972, pp. 77-78); the high geothermal gradients are also indicated by the presence of wairakite in the laumontite and prehnite-pumpellyite zones. Pumpellyite in glaucophanitic terranes In his studies of the glaucophanitic terranes of eastern central and southeast Celebes, De Roever (1947, 1950) was the first to note the common occurrence of pumpellyite without glaucophane or other critical minerals of the glaucophane-lawsonite subfacies, e.g., in the mineral assemblage albitepumpellyite-pale actinolite. He (1 947, p. 162) suggested that this pumpellyite might characterize “a separate metamorphic facies. containing among others colorless amphibole (i.e., actinolite-HJK), albite, quartz, and carbonate as other typical minerals, and which is a kind of precursor to the

392 lawsonite-glaucophanite subfacies”. Subsequently De Roever ( 1950, pp. 1458-1460) referred to it as the pumpellyite facies and considered it to be a very low-grade equivalent of the glaucophane schist facies. The appearance of pumpellyite, with or without actinolite, at a lower grade of metamorphism than lawsonite and glaucophane has also been reported from New Caledonia and Corsica (see references in Seki, 1961, p. 413), and from several glaucophanitic terranes in Japan. In the glaucophane-bearing Sanbagawa terrane of the Kant6 Mountains, central Honshn, Seki (1958, 1961; see Fig. 5-28) has found the incompletely recrystallized rocks of metamorphic zone I1 to be characterized by the assemblage pumpellyite-actinolite, with common stilpnomelane, but virtual absence of prehnite. Sodic amphibole, though common both in mafic rocks of this zone and of higher-grade pumpellyite-epidote-actinolite stage (zones I11 to V), rarely occurs in association with pumpellyite. The present writer is inclined to assign such pumpellyite-actinolite zones (with or without glaucophane, but without lawsonite), which have since been

Zones

Ia

Lb

Chlorite Stilpnornelane Sericite Purnpellyite Actinolite Sodic amphibole Epidote Lawsonite Jadeite Garnet Piernontite Stage

I

Pump Pump- epi - act -act

I

‘pi -act

Fig. 5-28. Ranges of occurrence of metamorphic minerals in the Sanbagawa metamorphic terrane of the Kant6 Mountains, central Honshii (modified after Seki, 1958, 1961). Note that the minerals indicated in each zone do not necessarily occur in the same rocks. ( I ) Rare prehnite occurs in this zone. (2) The association pumpellyite-sodic amphibole is rare: sodic amphibole occurs in mafic rocks. (3) Sodic amphibole is very rare; pumpellyite is absent. (4)Zone IV represents the lawsonite-pumpellyite-epidote-glaucophane sub-facies of the glaucophane-schist facies (Seki, 1961, pp. 409-41 l), or the glaucophane-iawsonite-schist facies of Fyfe and Turner (1966; after Winkler, 1965).

393 found in several other glaucophane-schist terranes in Japan (e.g., Hashimoto and Kashima, 1970; Hashmoto, 1968), to the pumpellyite-actinolite facies of Hashimoto (1966). An exception should be made for zones in which lawsonite and/or jadeite are associated with glaucophane, e.g., zone IV in the Kanta Mountains, which properly belong in the glaucophane-lawsonite-schist facies. In the Kanta Mountains, pumpellyite appears at a lower grade metamorphic stage than actinolite. However, even in the lawsonite and jadeite-free glaucophanitic terranes of the Chichibu metamorphic belt of western Shikoku (Hashimoto and Kashima, 1970; Hashimoto, 1972) and the Sangun metamorphic belt of the Katsuyama area, western Honshii (Hashimoto, 1968) (see Fig. 5-29 for metamorphic zoning in these terranes), appearance of glaucophane (with abundant pumpellyite, actinolite, and epidote) is preceded by a pumpellyite with subordinate actinolite. At least in the former area, this actinolite exclusively replaces relic clinopyroxene and is not associated with pumpellyite or epidote, whereas prehnite is only a minor mineral (Hashimoto, 1972, p. 25). When both prehnite and actinolite are rare, such pumpellyite zones cannot be considered to be transitional between the prehnite-pumpellyite and the pumpellyite-actinolite facies. Instead, they must belong to the general pumpellyite facies (of which they could possibly be considered to constitute an additional, third subfacies). In the French Alps the appearance of lawsonite in association with pumpellyite takes place at a lower metamorphic stage than that of glaucophane and actinolite (Bocquet, 1971, p. 75, figs. 2 and 5 ; Guitard and Saliot, 1971; Saliot, 1973). The metamorphic mineral zonation of the French-Italian Alps between the Pelvoux and Dora Maira massifs is schematically shown in Fig. 5-30. Similar pumpellyite-bearing lawsonite-albite zones with little or no actinolite or glaucophane are developed extensively in the lower-grade metamorphic zones of glaucophane-lawsonite-schist terranes of ( 1) the Franciscan metamorphic belt in California, (2) central Turkey, and (3) Calabria. They belong to the lawsonite-albite facies rather than to the pumpellyite-actinolite-schist facies. As re-defined by E.W.F. de Roever (1972, p. 147) from Winkler’s (1965, pp. 143, 148) original definition, the lawsonite-albite facies is distinguished from the higher-grade glaucophanelawsonite-schist facies by the absence of jadeitic pyroxene and the absence or extreme rarity of glaucophane-rich amphibole, but not necessarily of ferric sodium-pyroxene or ferric sodium-amphibole. A somewhat divergent definition was given by Winkler (1967, pp. 160, 164-165). This lawsonitealbite facies is intermediate between the lower-pressure zeolite and prehnitepumpellyite facies and the higher-pressure glaucophane-lawsonite-schist

394

Purnpellyite zone Prehnite Purnpellyite Epidote Actinolite

-

-1)-

Act i nolit e zone

-

----- - -1)- -

Glaucophane Stilpnomeiane

L

I

It

Prehnite-pum- Glaucophane pellyite facies schist fades Epidote Prehnite Purnpel lyite Act inolite Glaucophane St i I pnornelane

III Greenschist facies

------2)

- - -2)--

Fig. 5-29. Schematic metamorphic mineral zoning in basic and intermediate rocks of the lawsonite- and jadeite-free glaucophanitic terranes of the Oozu-Nomura district, Chichibu belt, western Shikoku (Hashimoto and Kashirna, 1970; Hashimoto, 1972), and the Katsuyama district, Sangun metamorphic belt, western Honshu (Hashimoto, 1968). (1) Actinolite only as a replacement product of relic clinopyroxene; not associated with purnpellyite or epidote. Prehnite found only in about 2% of examined thin sections of greenstone. (2) Actinolite and prehnite are rarely associated.

facies. The facies, however, appears between the prehnite-pumpellyite zone and a higher-grade epidote-pumpellyite-actinolite zone in the western Southland-Lake Wakatipu area, southern New Zealand (Landis, 1974; Kawachi, 1974), in a terrane containing virtually no glaucophane at all. Disregarding, for the moment, the latter occurrence, one can schematically distinguish three facial successions of pumpellyite-bearing zones in glaucophanitic terranes: ( 1) Prehnite-pumpellyite or pumpellyite (with very minor actinolite and epidote)-pumpellyite-actinolite-glaucophane (with epidote; no lawsonite or jadeite).

Zxternal o r Dauphinoise

Internal or Penninic zone Subbrian- Carboniferous of Internal Briahsonnais konnais Brianpnnais zone (Vanoise zone) and i t s Mesozoic sedimentary cover; Schistes zone and i t s Mesozoic sedimentary cover lustre%

Corrensite ( 1 1 Laumontite (2) no data Prehnite Pumpelly ite Epidote Act ino1ite Lawsonite (3) Glaucophane (31 Stilpnomelane Zoisite Garnet lsograds of Salliot (: 73) Iillite crystallinities diagenesis ant anchizone in SW,anchizone and epizone in NE part (4) Coal rank vitrinite

1

I

?

I

Let&-

only_incalc-schists

I 2.5- 3.5mm**22-2.7;

Occurrence

-

of paragonit. ( 5 )

25.1 mm (6) ==33-31mm (6)

3.7*/.(7) 5 % ( 8 ) 5.5 -1. ( 8 ) (anthracite t o meta-anthracite )

Fig. 5-30. Schematic low-grade metamorphic mineral zonation in the French-Italian Alps from the Pelvoux massif eastwards. Metamorphic mineral ranges based on data by Bocquet (1971), Guitard and Saliot (1971). and Saliot (1973). Width of zones not drawn to scale. ( I ) From Kubler (1970); (2) laumontite being replaced by prehnite (Saliot. 1973); (3) persistence of lawsonite beyond isograd 4 is thought by Saliot (1973) to reflect an older Alpine metamorphic phase; the same may be the case for all the glaucophane; (4) Aprahamian (1974); see also Kubler (1969); (5) data of Dunoyer de Segonzac (1969, pp. 176-180): low-grade limits of the anchizone and the epizone are 5.5 and 3.5 mm; (6) Abbas (1974); crystallinity limits as above; (7) Chateauneuf et al. (1973): (8) Robert (1971).

w

\o ul

396 (2) Pumpellyite (no actinolite; little or no prehnite) -pumpellyiteactinolite (glaucophane, but very rarely with pumpellyite; no epidote or prehnite)-pumpellyite-epidote-actinolite and epidote-glaucophane (glaucophane very rarely with pumpel1yite)-as above with jadeite and lawsonite (very rarely with actinolite). (3) Lawsonite (with pumpellyite; little or no actinolite; no glaucophane or jadeite) -lawsonite-glaucophane ( jadeite; no actinolite). It seems reasonable to assume that these three types of facies successions represent progressively higher-pressure gradients of low-grade metamorphism.

*

Control of zonal distribution of diagnostic burial-metamorphic zeolites und other Ca-AI-silicate minerals by primary petrographic and mineralogic composition In lowest-grade metamorphic sedimentary sequences, there is a considerable overlap or “telescoping” between the ranges of occurrence of various zeolites and related minerals diagnostic of the main mineral zories or facies. The overlap between heulandite or analcime and laumontite, and between laumontite and pumpellyite has already been mentioned. Moreover, in areas with zoning within the low-grade part of the zeolite facies, there may be telescoping of the clinoptilolite-mordenite zone and the analcime or heulandite zone, or even between the occurrence of the siliceous alkali-zeolites and of laumontite. Understanding of the extent, causes and controls of these apparent inconsistencies in zeolite-facies sequences is indispensable in utilization of these minerals for the elucidation of burial-metamorphic facies. In many of these zones of overlap there is fortunately a common tendency for a different mode of occurrence of the respective “incompatible” diagnostic zeolites and other Ca-Al-silicates. They are largely restricted to rocks of demonstrably different lithology, chemical bulk composition, or mineralogic composition. When occurring in the same hand-specimen, they either formed at the expense of different primary minerals, show mutually replacive relationships, or can otherwise be shown to represent non-equilibrium. Examples of each of these types of “overlap” and discussion of their dependence on various chemical and physical controls during deep burial are presented in the following sections. Occurrence of burial-metamorphic zeolites in volcanic elastic rock types Considering the wide range of chemical compositions covered by the zeolite group, zeolite minerals of distinctly different chemical composition expectedly tend to favor rocks of different chemical compositions. For instance, the siliceous alkali-zeolites clinoptilolite and mordenite tend

397 to prefer rocks of alkali-rich, siliceous composition, such as acid volcanics (mainly rhyolites and dacites). Analcime, heulandite, laumontite, and the less common wairakite tend to favor more calc-alkaline compositions, being, together with stilbite, the only zeolites common in basic tuffs. They also occur in a wide range of other rocks, mostly in deeply buried strata. The first three zeolites are very common as cements in sandstones, analcime being particularly common in acid tuffs and even in coal, whereas laumontite and wairakite are widespread in geothermal areas (cf. Iijima and Utada, 1972). A tendency for various diagnostic zeolites to form under the same physical conditions of burial in rocks of different bulk composition has been documented in several burial-metamorphic volcanic clastic sequences. Selective replacement of different primary mineralogic-petrographic constituents: telescoping of zeolite and Ca-AI-silicate zones In the case of telescoping of zeolite and Ca-Al-silicate zones, control of the burial-metamorphic mineralogy by bulk chemical composition of the system-at similar physical conditions of burial-is implicit in the notion of metamorphic facies. The restriction of the diagnostic authigenic zeolites and other Ca-Al-silicate minerals to specify primary constituents within the rocks-rather than encompassing the whole rock-suggests, however, that this apparent preference of different zeolites for a particular range of rock types mainly reflects their tendency to form at the expense of different primary constituents at given physical burial conditions. When such diverse primary constituents occur in distinctly different lithologies (which are either interbedded within one volcano-sedimentary sequence or are even present in the same polymict rocks), the resulting “telescoping” of the zonal ranges of occurrence of different zeolites- though appearing to be related to chemical bulk composition of the host rock-essentially reflects differences in primary mineral composition. Some examples are given below. Preferential replacement of silicic pyroclastics by alkali-zeolites and analcime. The role of selective replacement in the overlap of the distribution ranges of zeolites is exemplified by the preferential occurrence of the alkali-zeolites clinoptilolite and mordenite in rhyolitic and dacitic tuffs in the Miocene “Green Tuff” Formation in HonshU, Japan: Umegaki and Ogawa (1965, pp. 483-489) reported formation of laumontite around plagioclase and filling of interstices by laumontite in tuffaceous sandstones and basaltic-andesitic tuff in western Honshii. In interbedded rhyolitic tuff and tuff breccia, mordenite and clinoptilolite both occur around plagioclase and replace and vein volcanic glass in the matrix. In the

398 heulandite and laumontite zones of the Karasawa-zawa tuff member in the eastern Tanzawa Mountains, Japan, where heulandite replaces matrix glass of andesitic pyroclastics, whereas laumontite and analcime replace matrix glass and intermediate plagioclase of basaltic or andesitic pyroclastics, Yoshitani (1965) has noted the preferential replacement of glass in the matrix and pumice of more sodic, dacitic pyroclastic rocks. On a smaller scale, replacement of clastic constituents of markedly different mineralogic and chemical composition by different zeolite species may take place in the same hand-specimen, for instance, in polymict volcaniclastic rocks. Yoshitani (1965, pp. 210-210) described lapilli tuffs in which the predominant basic-andesitic pyroclastics are laumontitized, whereas the glass of some dacitic fragments has been altered to analcime or mordenite. Utada (1965; see also Iijima and Utada, 1965) stressed the replacive relationships of the authigenic zeolites in the Neogene pyroclastic sequence of the authigenic zeolites in the Shinjo Basin, northern Honshu. He showed that most of the zeolite minerals have a more extended zonal distribution range as replacement of volcanic glass and as fillings of interstltial voids than as replacements of plagioclase (Utada, 1965, fig. 9; see Fig. 5-31 in this chapter). Selective zeolitization may thus result in the formation of two zeolite mineral species at the expense of different clastic constituents in the same rock. For instance, in the rhyolitic to dacitic vitric and pumaceous tuffs of the analcime-heulandite zone I11 of the Shinjo Basin sequence, analcime and, only very locally, heulandite replaced the glass shards and filled pumice vesicles, whereas only heulandite has formed along the cleavages and around albitized oligoclase (Utada, 1965, p. 198; Iijima and Utada, 1966, pp. 336338). In some sequences, rhyolitic tuff is in part replaced by analcime. In the 2500 m thick Carboniferous Currabubula Formation of the Werrie Basin in the New England Geosyncline, New South Wales, heulandite replaced matrix and shards in coarse-grained acid to intermediate volcanic wackes, whereas clinoptilolite and, near the base of the sequence, laumontite occurs as a replacement and a vesicle filling in rhyolitic crystal-vitric tuffs. Analcime occurs in both types of rocks in the upper part of the section (Wilkinson and Whetten, 1964; Whetten, 1965, tables 1 and 2). In central Hokkaido, heulandite formed in altered andesitic glass, whereas the associated rhyolitic tuff was altered to analcimic rocks (Iijima and Utada, 1966, p. 338). One could still maintain that the occurrence of the alkali-zeolites (mordenite and clinoptilolite) merely reflects the presence of an unstable form of silica (see pp. 510-512) and, therefore, is not restricted to the upper analcime-heulandite “zone” of the zeolite facies. In some sequences, however, alternation of volcanic clastic rocks of widely different composition

399 Depth ( k m 1

0

2.8 3

2

0.9

I

ZONE

I

m

II

Y 1

Clinopti lolite Mordenite Analcitt Heulandite La um ont ite

Low c r ist oba I i t c Chalcedonic quartz Quartz Albite Adularia Epidote Montmor iI Ion it. ”Intermediate clay”

I

42=4b

Chlorite Ser i c i t e Calcite

p 2

---

t-

----------

4

Fig. 5-3 I . Zonal distribution of diagenetic minerals in the Neogene pyroclastic formations (“Green Tuff Formation”) on the eastern margin of the Shinjo Basin, Yamagata Prefecture, northern Honshii. I =replacing volcanic glass; 2 =precipitation in interstitial voids; 3 =replacing plagioclase; 4 =replacing mafic minerals. (From Iijima and Utada, 1966, fig. 6, as modified from Utada, 1965, fig. 9.)

may produce a telescoping of zeolites and other Ca-Al-silicates considered characteristic for different zones of burial metamorphism. Such is the case when analcime and/or heulandite persist into the laumontite zone: In the Brothers Peak Formation of north-central British Columbia, heulandite and analcime-quartz are restricted to vitric tuffs, mostly replacing glass shards, the plagioclase being unaltered and still in the intermediatehigh structural state; whereas laumontite is restricted to the cement of the

400 interbedded non-tuffaceous sandstones (Read and Eisbacher, 1974). The occurrence of laumontite in the lower formation of the Sustut Group, to the exclusion of heulandite and analcime, reflects either the gross mineral zoning with depth or the absence of vitric tuffs. Analcime and laumontite have even been reported from the same andesitic and basaltic lapilli-tuffs and tuff-breccias (Yoshitani, 1965). On the other hand, in many cases telescoping of diagnostic burialmetamorphic minerals seems to be unrelated to differences in primary mineralogy, as exemplified by the overlap in vertical distribution of (1) analcime and laumontite in the Barranquitas tuff member of the east-central Puerto Rico (Otalora, 1964); (2) analcime, heulandite, and laumontite over several kilometers in the Triassic of the Hokonui Hills of Soutland, New Zealand (Boles and Coombs, 1975); and (3) laumontite and prehnite over more than 3000m in the Karmutsen Group of Vancouver Island (Surdam, 1973). These occurrences must be explained by other factors, to be discussed in separate sections of this chapter, e.g., differences in porosity and permeability (pp. 417-418), and varying ionic activity ratios in stratal waters (pp. 424-432). Persistence of laumontite into the prehnite-pumpellyite zone. In deeper burial zones, the occurrence of laumontite and pumpellyite appears to be governed, at least in part, by similar mineralogic-compositional factors. In the Broughton Sandstone of the Kiama district, New South Wales, assemblages of pumpellyite-epidote-sphene are mainly restricted to latite fragments in a breccia, whereas laumontite is an almost ubiquitous alteration product of calcic plagioclase and cement in the volcanic conglomerates, lithic-feldpathic sandstones, and siltstones (Raam, 1968, p. 329). One of the most impressive, though comparatively straightforward examples of compositional control of the zonal distribution of authigenic Ca-Alsilicates is the vertical overlap of the occurrence of laumontite, prehnite, and pumpellyite in the different volcaniclastic lithologies in a 12- to 15-km thick composite sequence of Late Triassic and Early Jurassic age in the Aldrich Mountains, Oregon (Brown, 1961; Brown and Thayer, 1963). Pumpellyite occurs in amygdules and as veins in basaltic lavas and coarse lithic tuffs of the lower units, and as interstitial patches in tuffaceous graywacke of the upper unit V, whereas laumontite occurs in the same stratigraphic units as a replacement of vitric debris in silicic vitric tuffs and as interstitial material in a dacitic breccia. Prehnite occurs in the intermediate units I1 to IV. The relationship between lithology and au thigenic Ca-Al-silicates is so marked that these minerals do not appear to be associated in any rock type. As a result, there is no good correlation between authigenic mineralogy and depth of burial, although there is a distinct increase in the extent of albitization of

40 1 detrital plagioclase with depth, and the Jurassic tuffs overlying the upper unit V contain both laumontite and heulandite. Summary of preferential replacement. The alkali-rich, siliceous zeolites clinoptilolite and mordenite tend to occur almost exclusively as diagenetic products of rhyolitic and dacitic pyroclastic rocks, in which they preferentially replace volcanic glass (in which most of these acid pyroclastics abound), and occur as a cement (cf. Iijima and Utada, 1966, pp. 337-338, 342; 1972, p. 70). Such selective replacement of volcanic glass reflects both the higher alkalic contents and the higher silica activity in the acid volcanic glass than in the accompanying feldspars. In contrast, laurnontite mainly occurs as a cement of sands and as a replacement of intermediate plagioclase in non-tuffaceous sandstones and conglomerates, and at the expense of earlier-formed heulandite. Laumontitization of glass is rather exceptional. The replacement of rhyolitic glass by laumontite (e.g., Brown and Thayer, 1963) may be restricted to more deeply buried series in which the less silicic volcanics are already being replaced by prehnite and/or pumpellyite. Extent of reconstitution in the zeolite facies Different zeolites in chemically similar rocks of contrasting mineralogic composition. In the examples given in the previous section, the major compositional control of the variety in burial-metamorphic mineralogy at given burial conditions appeared to be the composition and reactivity of the primary phases being replaced. The differences in chemical composition between the rocks showing contrasting zeolite mineralogy may be comparatively minor: the only significant difference between the chemical compositions of the mordenite-bearing and the laumontite-bearing rhyolitic tuffs of the Izumi mountain range, southwestern Honshu, appears to be the higher Na,O/CaO and (Na,O/CaO K,O)/CaO ratios in the mordenite-bearing rocks (Nakajima and Tanaka, 1967, Table4; cf. Seki et al., 1972, pp. 153-154). The overlap of different zeolites may in fact be due to different extent of zeolitization of interbedded volcanic rocks of similar chemical bulk composition, but different mineralogic composition. In the eastern Tanzawa Mountains, Seki et al. (1972) have observed interbedding of thin-bedded, fine-grained mordeni te-bearing pumice tuff within coarse-grained laumontitized andesitic lapilli tuff and tuff breccia of the same bulk chemical composition. Through subtraction of the chemical composition, of about 20% augite and calcic plagioclase of the pumice tuff, which are not zeolitized, Seki et al. calculated that the composition of the

+

402 remaining fine-grained pumice fragments (about 80%) is more sodic and silicic than that of the bulk material of the lapilli tuffs and the tuff breccias. They suggested that sea water trapped in the relatively impermeable finegrained pumice tuffs reacted with this alkali-rich siliceous glass to devitrify it in to mordenite-bearing assemblages. On the other hand, in the permeable, relatively coarse-grained crystalline tuff breccias the sea water reacted with the calcic plagioclase and augi te to form laumonite- or prehnite-bearing mineral associations. Incomplete reconstitution and mineral facies. The major role of reaction kinetics in zeolite-facies reactions is reflected in the confinement of zeolitization to either the more reactive clastic phases and their immediate surroundings, or to the cement of the clastic framework, and in the failure of the less reactive unstable phases to participate in the reactions. The latter consequently persists as metastable relics until attainment of a more advanced stage of burial-metamorphic alteration. Inasmuch as only some of the mineral phases participate in *thereactions and, thus, only part of the rock is involved in the establishment of chemical equilibrium, a direct relationship between its bulk chemical composition and the authigenic mineral assemblage - as demanded by the concept of metamorphic facies - does not necessarily exist. In incompletely reconstituted rocks such a relationship is to be expected only between the chemical composition of that portion of the rock participating in the establishment of equilibrium and the zeolitic mineral assemblages. Coombs et al. ( 1959, p. 54, footnote), therefore, considered zeolite-facies mineral assemblages in such incompletely reconstituted rocks "in terms of the mineral facies concept of Eskola (1920, pp. 145-146), rather than in terms of metamorphic facies". Progressive reconstitution and albitization in the laumontite zone. The common localization of the Ca-zeolites formed at the expense of calcic plagioclase along the fringes and the cleavages of the albitized plagioclase clasts, rather than within the fragments, indicates that in addition to hydration, the zeolitization and albitization of plagioclase involves movement of ions over about a millimeter, probably in solution. The local character of this migration is reflected by its small scale, which is insufficient to produce stable, compatible mineral assemblages involving the whole rock. As the intensity of alteration increases with depth of burial, the zeolitization tends to become less selective. Coombs (1954, pp. 71-72) and Coombs et al. (1959, pp. 59-60) (cf. Fig. 5-22) have commented on the markedly stronger albitization and zeolitization of detrital plagioclase in the laumontite zone than in the higher heulandite-analcime zone of the Tarin-

403 gatura section of Southland (New Zealand) (cf. Boles and Coombs, 1975). Similarly, in the basic andesites and andesitic pyroclastics of the earliermentioned Karasasawa-zawa tuff member of the "Green Tuff" Formation of the Tanzawa Mountains, central Honshu, heulandite replaces glass of the matrix and of pumaceous fragments (but rarely the calcic plagioclase), whereas the laumontite of the lower horizons of the tuff member generally replaces the medium plagioclase as well as the glass of the matrix (Yoshitani, 1965, pp. 209-210). Albitization of calcic plagioclase is generally well advanced in the laumontite zone of most areas. For example, more than 70% of plagioclase grains are albitized in the non-calcareous laumontite sandstones of the Great Valley sequence, California (Dickinson et al., 1969). Albite is virtually the only plagioclase in laumontite-bearing tuffs, volcanic arenites or siltstones of the Taveyanne Graywacke of Haute-Savoie (Martini, 1968, pp. 592-593); the Upper-Cretaceous and Tertiary graywackes of the western Olympia Peninsula, Washington, and of Vancouver Island and the adjoining Gulf Islands, British Columbia (Stewart, 1970, p. 49; 1974: Stewart and Page, 1974); and the volcaniclastic Broughton Sandstone of coastal New South Wales (Raam, 1968, pp. 323-324) and Robles Formation of south-central Puerto Rico (Jolly, 1970). In the Taringatura section and in the Parry Group of western New England the albitization of calcic plagioclase is completed (apart from local relics) upon the appearance of pumpellyite and authigenic epidote (Coombs, 1954, p. 71; Coombs et al., 1959, fig. 2; Packham and Crook, 1960) (see Fig. 5-23). Albitization may be entirely completed only in the lowest-grade part of the prehnite-pumpellyite-metagraywacke facies, e.g., in northern Otago (Bishop, 1972b), and relics of oligoclase in prehnite-pumpellyite zones have been reported by Hawkins (1967, p. 813) and Seki et al. (1969a, p. 18). Generally, however, albitization of plagioclase in this facies is virtually total (e.g. Martini, 1968, p.633; Hashimoto, 1968, pp. 124-125; Coombs et al., 1970), even in calcareous sandstones (Dickinson et al., 1969). Incomplete albitization in such advanced burial-metamorphic facies is usually associated with other evidence of incomplete reconstitution, such as the persistence of heulandite in the pumpellyite-bearing andesitic clastics of central Oregon (Dickinson, 1962b) (see pp. 404-405). The contrast between the solution of volcanic glass associated with fresh, unalbitized plagioclase at shallow levels, and the replacement of both glass and clastic plagioclase- though often by different minerals-at deeper levels (e.g., Surdam, 1973), constitutes an indication of the progressive reconstitution with depth.

404

Limits of burial-metamorphic mineral zones The cases presented above illustrate the overlap between the zonal ranges of diagnostic zeolites in sequences of pyroclastic rocks with widely different chemical compositions and primary mineralogies. Sharp boundaries between burial-metamorphic mineral zones are to be expected only if consideration is restricted to mineral assemblages in rocks similar not only in chemical bulk composition-as required by the concept of metamorphic facies-but in primary clastic or volcanic mineralogic composition. Most of the discrepancies reviewed appeared to involve the replacement of rhyolitic vitric material by clinoptilolite and mordenite. in otherwise heulandite-, analcime-, or laumonite-bearing sequences, and in at least one example by laumontite in an otherwise prehnite- and pumpellyite-bearing sequence. This effect could mainly be controlled by the oversaturation in silica around silicic volcanic glass. In establishing mineral-zone limits, it may be advisable to consider mainly the authigenic mineralogy of quartz-bearing but crystalline pyroclastics (including rhyolitic compositions) and lithic-feldspathic sandstones and siltstones, and to ignore the anomalous replacements of silicic glass. Such qualifications could appreciably reduce the overlap between zeolite zones due to differences in rock composition. Selective replacement: persistence of earlier-formed mineral assemblages during subsequent alteration In addition to the various cases of telescoping controlled by primary mineralogy, telescoping of mineral facies of burial metamorphism may also be due to the preservation of zeolites formed during an earlier stage of burial upon alteration of the previously-unaltered constituents of the same rocks during a subsequent deeper-burial episode. Such a case is exemplified by a Jurassic sequence of andesitic clastics in central Oregon, in which Dickinson ( 1962b) has distinguished three successive phases of burial diagenetic alteration: ( 1) Devitrification of andesitic glass fragments to heulandite in association with celadonite and chlorite; the medium plagioclase remains unaltered. (2) Local replacement of both plagioclase and the earlier-formed heulandite by laumontite, often with and, sometimes, without accompanying albitization. (3) Widespread, but commonly incomplete, decomposition of the medium plagioclase to albite and pumpellyite with or wit,hout prehnite, while the zeolites and ferromagnesian phyllosilicates- formed earlier from the unstable vitric materials-were preserved. Pyrogenetic augite and hornblende remained unaltered during all stages. The spatial pattern of phases (2) and (3) bear no consistent relationship to stratigraphic or structural position. The dominant mineral assemblage of the

405 andesitic tuffs is the product of two successive phases of alteration, which impose their effects upon two different types of original constituents. I t consists of three principal coexisting sub-assemblages: (a) relic augite and hornblende, (b) albite-pumpellyite-prehnite derived from plagioclase, and (c) heulandite-celadonite-chlorite derived from the unstable glassy constituents. No chemical equilibrium encompassing the bulk rock was attained. The association of these incompatible Ca-Al-silicate phases and the incompleteness of stage (3), as indicated by the persistence of both the earlierformed zeolites and part of the calcic plagioclase, is thought to be controlled by unequal local distribution of interstitial water. Discontinuous burial-metamorphic mineral zoning: successive metamorphic episodes or post-metamorphic thrusting? Discontinuous downward burial-metamorphic zoning has been described by Levi (1970) from the Andean geosynclinal deposits of the Andean Range and the Coastal Range of central Chile, with a cumulative thickness of 15-28 km. Each one of four successive stratigraphic-structural units show a downward increase in grade covering different intervals between zeolite, prehnite-pumpellyite, and greenschist facies. The major unconformities separating these units constitute breaks in the downward mineral zoning, higher-grade assemblages generally overlying lower-grade assemblages, i.e., greenschist facies occur above prehnite-pumpellyite facies, or prehnitepumpellyite facies above “laumontite-bearing prehnite-pumpellyite facies” (Levi, 1970, p. 1008). Levi (1970, p. 1010) explained this repeated burialmetamorphic zoning by assuming at least four successive burial-metamorphic episodes, each one taking place prior to a folding period and leaving unaffected the underlying stratigraphic-structural units, which were supposedly “sealed” by the alteration minerals of the preceding metamorphic episode. Zen (1974b), however, has pointed out that if the whole cumulative thickness of 18 km above the deepest laumontite occurrence was complete at any one time, the laumontite should have decomposed to form lawsonite at the prevailing pressure of 5 kbar (Liou, 1971a). The difficulty would be resolved by verification of Thompson’s suggestion (quoted by Zen, 1974b, p. 449, footnote) that the “unconformities” may actually be large low-angle thrusts, which would invalidate Levi’s interpretation of episodic burial metamorphism. Thrusting of higher-grade over lowerrgrade burial-metamorphic rocks has been recorded from the Glarus Alps, eastern Switzerland, where greenschistfacies Permian volcanics (Verrucano) of the Helvetic Axen nappe overlie autochthonous prehnite-pumpellyite facies of the Taveyanne Graywacke

406 (Martini and Vuagnat, 1970, p. 59). It seems likely that several more cases of tectonic inversion of burial-metamorphic zoning will be shown to exist in various nappe areas. Progressive mineral reactions in the zeolite facies Clastic sediments contain appreciable amounts of water, both as constituents of minerals and as interstitial fluid, which are largely lost by the beginning of the greenschist facies (cf. Fyfe, 1974). Progressive dehydration and the general tendency towards higher density of the prograde phases, which are characteristic of progressive regional metamorphism of sediments in general, are, therefore, particularly noticeable in burial-diagenetic and lowest-grade metamorphic reactions. For instance, the Ca-zeolites stable at higher temperatures, such as laumontite, scolecite, and wairakite, are denser (sp. gr.> 2.2) and less hydrous than the lower-temperature types, such as chabazite, stilbite, and heulandite. Table 5-VI lists the chemical compositions and specifi'c gravities of some of the zeolites and other Ca-Al-silicate minerals discussed in this chapter. In view of the high entropy of water, the AS of dehydration reactions is almost always positive, i.e., the dehydration reactions proceed with increasing temperature. The AV of such dehydration reactions will generally also be positive, owing to the large molecular volume of water at low temperatures

TABLE 5-VI Chemical compositions and specific gravities of zeolites and some other hydrous Ca-Al-silicate minerals mentioned in this chapter Mineral name and formula

Specific gravity

Analcirne, Na[A1Si,06]~H,0-Na[A1Si,0,]~ I f H,O Mordenite, (Na,, Ca)[AI,Si 100,,]. 6 H,O 4-34 H 2O Laumontite, Ca[Al ,Si ,O Heulandite, (Ca, Na,, K2)[AI,_,:Si,:_,0,,].6 H 2 0 .6 H,O Clinoptilolite, (Na,, K ,, Ca)[AI I ;Si 7;01x] Stilbite, (=desmine), (Ca, Na,)[A1,Si,01,]~7 H,O 2 H ,O Scolecite, Ca[Al ,Si ,O Wairakite, Ca[AI ,Si,O,,]. 2 H,O

2.24-2.29 2.12-2.15 2.26-2.29 2.18 2.14 2.14-2.17 2.25-2.29 2.27-2.29

Prehnite. Ca2AI[A1Si,O,, /(OH),] Pumpellyite, Ca,(Mg, Fe, AI)AI,[SiO,/Si,O,/(OH),/H,O, OH] Zoisite/clinozoisite, Ca,(AI, Fe)AI ,[SiO,/Si ,O,/O/OH] Epidote s.s., Ca2(Fe3+,AI)A1,[SiO,/Si2O7/0/OH] Lawsonite, CaAI,[Si ,07/(0H),]. H,O

2.90-2.95 3.18-3.23 3.12-3.38 3.38-3.49 3.05-3.12

I

407 and pressures. Inasmuch as A S and A V are both positive, it follows from the Clausius-Clapeyron equation d P / d T = AS/AV, that the slope d P / d T of the univariant dehydration curves will also be positive, i.e., that the equilibrium temperature T, will rise with increase in pressure. Inasmuch as A V of dehydration reactions is generally positive and large at pressures below 1 kbar, d P / d T will be very small, whereas at high pressures the volume of the liberated water and, hence, A V is markedly reduced. The slope d P/dT, increases and may actually become negative at very high pressures due to the general decrease in total volume of the solids in prograde metamorphic reactions (i.e., AColidsis negative). The equations for a number of important zeolite-facies reactions, indicating the change in molar volumes at 1 bar and 25"C, either calculated from the molar volumes in Robie et al. (1967) or taken from Coombs et al. (1959) and Campbell and Fyfe (1965) are presented below: NaAlSi,O,. H,O analcimr

+ SiO, eNaAlSi,O, + H,O quartz

albite

(AVO

CaAl,Si,O,, . 5.3 H,O eCaAl,Si,O,, . 4 H,O heulandite

- 1.92 cm3),

water

laumontite

+ 2 S O , + 1.3 H,O quartz

water

(AVO = -6 cm3), CaAl,Si,O,, . 4 H,O =CaAl,Si,O,, laumontite

*

wairakitr

2 H,O

+ 2 H,O

(AVO = +22 cm3).

water

The negative AVO and the small A S values in some of these reactions are due to the very large contribution of very loosely held water to both the entropy So and the molar volume V" of hydrated phase such as zeolites. The large entropy of analcime, for instance, causes the A S of the dehydration reaction of analcime in the presence of quartz to be close to zero (Coombs et al., 1959, p. 81; Campbell and Fyfe, 1965, pp. 809-810). Due to the large apparent molecular volume of water in analcime, the AVO will be unusually small [ - 1.92 cm3 at 25°C according to Campbell and Fyfe (1965, p. 813)] and the volume of the solid phases ( ATolld5) will decrease appreciably. Consequently, the d P/dT, slope of the univariant equilibrium curve for the reaction at relatively low pressures will not only be negative, but surprisingly flat (Campbell and Fyfe, 1965; cf. Turner, 1968, p. l58), i.e., increasing fluid pressure should cause dehydration. As a result, restricted low-pressure fields for such minerals as wairakite are to be expected. The equilibrium conditions of these zeolitic reactions are functions of temperature and pressure only as far as water pressure equals total pressure. Conditions of several diagnostic phase equilibria have been determined experimentally for p H l O= Ptntal and are discussed below.

408 Experimental study of lowest-grade metamorphic equilibria as a function of temperature and water pressure In the initial stages of the study of burial metamorphism, direct experimental calibration of equilibrium conditions was hampered by the impossibility of reversing the reactions, a requirement for experimental establishment of univariant reaction curves. Considerable amounts of experimental work over the last fifteen years, however, and the introduction of new approaches, such as using a reaction-rate method involving the weight change of quartz or albite crystals, have provided a reasonably satisfactory insight into the physical conditions of diagnostic low-grade reactions at conditions of p H Z = O Ptota,. The current state of knowledge is briefly summarized here, concentrating on the more recent results. Diagnostic laumontite-zone equilibria (experimental euidence) The beginning of the laumontite zone is defined by the appearance of the Ca-zeolite laumontite, commonly at the expense of heulandite. The analcime quartz assemblage commonly disappears approximately at the beginning of the laumontite zone. The critical reactions are as follows:

+

+ 2 or 3 quartz + H,O, and analcime + quartz = albite + H,O.

heulandite = laumontite

The latter reaction has been experimentally studied by Coombs et al. (1959), Campbell and Fyfe (1965), Liou (1971b) and Thompson (197 la). The latter two authors both found equilibrium curves with negative d p H Z o/dT slopes at higher pressures. Equilibrium was found to be about 200°C at 2 kbar p H I O , 196" I+ 5" at 3 kbar pHZO, about 190°C at 4kbar pcoZ,and 183" 5°C at 5 kbar pHIOby Liou (1971b), or 190" 10°C at 2 kbar pHZO, 170" =k 10°C at 4kbarpFZq,and 150°C at about 4.75 kbarpHZoby Thompson ( 1971b). These equilibrium temperatures are lower than those of Coombs et al. (1959), but higher-at least at H,O pressures above 2 kbar-than those of Campbell and Fyfe (1965), whose equilibrium curve has a more pronounced negative d p H I 0 / d T slope (see Fig. 5-32). At a geothermal gradient of 25"C/km, the analcime quartz assemblage would thus persist to a maximum depth of 7 km, which is in good agreement with the disappearance of analcime in the Taringatura and the adjoining western Hokonui Hills sequences below the present depth of approximately 5 km, or. a maximum depth of burial of 8-83 km before erosion (Coombs, 1954, 1961; Boles and Coombs, 1975, p. 170). It has not been possible experimentally to reverse the reaction heulandite = laumontite 2 or 3 quartz H,O, and the equilibrium relations of heulandite and laumontite are thus not quite clear. Coombs et al. (1959) could only show that below 2 kbar pHIOheulandite was unstable at temperatures above 280°C. Subsequent work by Nitsch

*

+

+

+

409

Temp. ( . C )

Fig. 5-32. P-T diagram showing low-grade Na-Al-silicate and Ca-Al-silicate univariant equilibria after Thompson (1971a). A - B is the equilibrium curve for analcimei quartz= albite+H,O. A - C is Campbell and Fyfe's (1965) curve for the same reaction. E - F is the estimated equilibrium curve for the reaction heulanditea laumontite + quartz + H 20.The shaded region represents pressures and temperatures inferred for the analcime-albite reaction from Coombs (1954) at Taringatura. The curves relating laumontite, lawsonite, and anorthite are from Thompson (1970b).

(quoted in Winkler, 1967, p. 158, and 1968, p. 1013) suggested that heulandite was stable up to 235" 10°C at 10 kbar and about 230" at 7 kbar The equilibrium temperatures were thought to be only slightly lower at decreased pressures.

*

410 Observations in the Wairakei geothermal area, New Zealand, however, suggest that laumontite may form at the expense of zeolite at about 200°C at relatively low pressures; the reaction temperature is lowered with increasing pH,O(Coombs, 1971). In the Wairakei drillhole 225, laumontite is very locally present at a temperature of 140°C at a very low water pressure (Steiner, 1968, fig. 3). Tentative results on the heulandite-laumontite reaction by Thompson (1971a). which suggest that equilibrium temperature is in the range of 12O0-14O0C for 2 kbar p H ? O tend , to confirm that the equilibrium temperatures are significantly lower than believed previously. According to Thompson's ( 1971a) estimate, the d p H , o / d T slope for the reaction curve is strongly negative, and this reaction curve should intersect the curve for the analcime-albite reaction (see p. 408) in the temperature region of 175"210°C and p H , Opressure of 1.5-2 kbar (see Fig. 5-32). If correct, this pressure for the crossover point is only slightly higher than that presently prevailing at a depth of 5 km in the Taringatura section, below which analcime is absent and heulandite is being replaced by laumontite (Coombs, 1954). This suggested to Thompson (1971a, pp. 89-90) that only very minor erosion could have affected the Taringatura section since the metamorphism. Greater depression of the temperature of formation of laumontite (at the expense of heulandite) by higher pressure than that of albite (at the expense of analcime + quartz) could explain the persistence of analcime throughout the laumontite zone in the absence of heulandite in terranes such as Puerto Rico (Otalora, 1964; Jolly, 1970), the Tanzawa Mountains (Seki et al., 1969a, table I; Shimazu et al., 1971, table 2) and Vancouver Island (Surdam, 1973). I t should be noted, however, that in the latter area, analcime occurs mainly in basic rocks. In the absence of quartz the reaction: analcime

+ quartz -, albite + water

is inhibited and analcime persists to higher grades of 6urial metamorphism. At pressures above 2 i to 3 kbar, however, lawsonite rather than laumontite is the stable Ca-Al-hydrosilicate phase in the temperature range, as indicated by studies of the reaction: laumontite = lawsonite + quartz

+ H,O

(Nitsch, 1968; Liou, 1969; Thompson, 1970b). The equilibrium pressures are relatively little affected by temperature (see Fig. 5-32). An unreactive rock undergoing metamorphism along a low geothermal gradient of 10"C/km may thus bypass the zeolite field entirely, lawsonite being the first Ca-Alhydrosilicate formed rather than a Ca-zeolite.

41 1 Physical conditions of the high-grade limit of the zeolite facies (experimental evidence) The high-grade boundary of the zeolite facies is defined by the disappearance of laumontite. Thompson's (1970b) study of the reaction: laumontite = anorthite quartz H,O

+

+

suggested an equilibrium of 310" f. 10°C at a pressures of 1 kbar, 317" 2 10°C at 2 kbar, 338" k 10°C at 4kbar, and 347" k 10°C at 6 kbar ( p H , O= P!ota,).The steep dpHz,/dT slope of this reaction curve indicates that the disappearance of laumontite in progressive regional metamorphism or deep burial should be relatively pressure-independent below about 3 kbar. at geothermal gradients of more than 25-3O0C/km. At 3 kbar pressure the reaction curve is intersected at about 320°C by that for the laumontitelawsonite equilibrium curve mentioned above; this represents the highpressure limit of the laumontite zone (see Thompson, 1970b. fig. 3: Fig. 5-32 in this chapter). The occurrence of wairakite the Ca-analogue of analcime at a highertemperature range than laumontite in geothermal areas (Coombs et al., 1959: Steiner, 1968; Seki et al., 1969b) and its appearance in the higher-grade part of the laumontite zone of the regional-metamorphic terrane of the Tanzawa Mountains (Seki et al., 1969a), where it even persists into the prehnitepumpellyite zone (Shimazu et al., 1971), suggests that at low pressures a wairakite field may be located between the field of laumontite and anorthite quartz. Experimental work by Coombs et al. (1959) and Liou ( 1970. 1971a) indicated that the low-pressure dehydration of laumontite does not follow the laumontite-to-anorthite reaction, presented above, but instead follows a two-stage dehydration process via wairakite. Liou ( 1971a) found the equilibrium dehydration curve:

+

laumontite = wairakite

+ H,O *

to pass through about 230°C at 0.5 kbar, 255" 5 5°C at 1 kbar, 282" 5°C at 2kbar, and 297"*5"C at 3kbar (see Fig. 5-32). According to Liou ( 1970), the equilibrium dehydration: wairakite = anorthite

+ quartz + H,O *

occurs at 330" +5"C at 0.5 kbar, 348" t 5°C at 1 kbar, 372" 5°C at 2 kbar, and 385" 5°C at 3 kbar, so that wairakite has a thermal stability range of about 90°C for p H z O= Ptota,up to about 3 kbar (see Fig. 5-32). The laumontite-to-anorthite reaction as determined by Thompson ( 1970b), falls between Liou's dehydration curves presented above and, thus, may be metastable at low pressures. The .three curves meet in an invariant point at about 3.5-4 kbar, which is already within the lawsonite stability field. That

*

412 wairakite is nevertheless absent from many burial-metamorphic terranes is related by Thompson (1970b, p. 272) to its preferential formation in silicasaturated environments, such as may be expected in geothermal rather than in regional metamorphic areas, in which silica in solution is normally in equilibrium with quartz. In reality, the low-temperature part of the anorthite field is represented by Mg and/or Fe-bearing Ca-Al-hydrosilicates such as epidote, prehnite, and pumpellyite, coexisting with albite (cf. Thompson, 1970b, fig. 3, caption). The prehnite-pumpellyite-metagraywacke and the pumpellyite-actinolite -schist facies are characterized by the occurrence of pumpellyite, with or without prehnite and by the absence of laumontite. Prehnite, which is often associated with laumontite, and, therefore, is not by itself characteristic for the higher-grade facies, may form at the expense of laumontite below its upper thermal stability limit through a reaction such as the one suggested by Coombs et al. (1959): laumontite = prehnite

+ “kaolinite”+quartz + H,O

Although direct experimental study of this reaction is complicated by the concurrent breakdown of kaolinite to montmorillonite, Thompson ( 1971b, pp. 152- 153) established indirectly that at 2 kbars p H 2 0the equilibrium temperature for the reaction: prehnite

+ “montmorillonite”+ quartz = laumontite + H,O

was 270” t 15°C (see Fig. 5-39). Other prehnite-forming reactions at the expense of laumontite, such as the one proposed by Coombs (1961; see also Winkler, 1967): laumontite

+ calcite = prehnite + quartz + H,O + CO,

are strongly dependent on the p c o 2 / p H , 0 ratio, in addition to temperature and pressure. The effect of the pco,/pH,o ratio is discussed below (p. 432ff.). Pumpellyite has also been noted to appear already in the high-grade part of the zeolite facies, for instance by a reaction such as the one suggested by Coombs (1961, p. 208): laumontite

+ prehnite + chlorite = pumpellyite + quartz + H,O.

I t may be concluded from these data that the upper stability limit of laumontite, which constitutes the high-grade boundary of the zeolite facies, is comparatively pressure-independent (at conditions which preclude stability of lawsonite). The extent of the zeolite facies may be camouflaged, however, if laumontite is widely replaced by wairakite, prehnite, or pumpellyite before this stability limit is reached.

413 Constraints on the application of experimentul equilibrium dutu In evaluating experimental mineral-equilibrium data and applying them to naturally occurring mineral assemblages, it should be kept in mind that these equilibrium relationships have been studied under conditions that constitute simplifications with regards to most natural occurrences, i.e., in idealized chemical systems with few components and under conditions of p H , O= P,,,,l. In considering the relevance of the equilibrium conditions thus obtained, account should be taken of the effects on equilibrium conditions of the following factors: ( 1) The reduction of dehydration temperatures by lowered chemical activity of water, either by various solutes in the fluid phase ( p H , O< Pr,llic,) or by high porosity and permeability ( p H L O=ePnuid < PI,,,,). (2) The expansion of the stability range of siliceous minerals by oversaturation with silica. ( 3 ) Extension of mineral stability as determined on pure end-member compositions in natural phases, which depart considerably from such compositions due to non-stoichiometry and ionic substitutions in the zeolites. Effects of reduced water pressure Inasmuch as most prograde lowest-grade metamorphic reactions involve dehydration, the water pressure will play a particularly important role in determining the equilibrium conditions. In the preceding considerations. p H , Ohas been taken to be equal to the lithostatic pressure. Pr<,ta,.resulting from the overlying rock column. In buried sedimentary rocks. however. p tl,O may also be lower than Plots,. Inasmuch as the dehydration reactions generally proceed with increasing temperature due to their positive AS, lowering of the partial pressure or the chemical activity of water will lower the equilibrium temperatures of such reactions. Moreover, inasmuch as AYolidsof dehydration reactions is invariably strongly negative, the d P / d q slopes of univariant dehydration equilibria are likely to be more steeply positive under conditions of pH,()< than when p H , O= Prota,.In reactions that liberate large amounts of water, the slopes may well become negative. The dependence of equilibrium conditions on the pH,O/Plo,a, ratio is shown for the reaction studied by Liou (1971a):

laumontite

+

wairakite

+ water (see Fig. 5-33).

Decrease in pH,o has the effect of displacing facies boundaries involving dehydration reactions to lower temperatures. Consequently. dehydration sequences of zeolites and other hydrous Ca-Al-silicate minerals with depth could be favored not only by increase in temperature, but also by decrease in the pHzo/Pl,talratio with depth. If water is being expelled by dehydration reactions and porosity is low,

414

Temp. PC)

Fig. 5-33. P-T diagram for the reaction laumontite=wairakite+ fluid, under various Pnu,,, /PtOta, ratios, and for various conditions of constant pH,O(from Liou, 1971a, fig. 4).

which is likely to be the common case at greater depths, Pflujd is likely to be equal to Plead (e.g., see Fyfe, 1974). At shallow depths, however, Pfluid may be lower than Plead. Whenever intergranular pore spaces are connected for an appreciable vertical distance and in the presence of open vertical fissures, the pore-fluid pressure may be much lower than Plead. In fissures open to the surface, pore pressure theoretically may be as low as one third of Plead, assuming that sedimentary rocks have an average specific gravity of around 2.8. Coombs et al. (1959, p. 83) have pointed out that analcime might grow in well-sealed rocks in which Pnuid= Plead, while at the same time albite might grow in an adjoining fissure. Such reduced Pfluid with respect to lithostatic pressure could also explain the local occurrence of more dehydrated minerals, for instance, the occurrence of laumontite in joints and fissures in clinoptilolite- and mordenite-bearing tuffs (Coombs, 1958; Nakajima and Tanaka, 1967), and replacement of laumontite by less hydrous Ca-Al-silicate phases (prehnite, pumpellyite, and epidote) along joints and minor faults in the Taveyanne graywackes of the Helvetic zone, Swiss Alps (Martini and Vuagnat, 1965; see also Coombs, 1971, p. 323).

415

Reduced fluid pressure-osmotic conditions. Coombs et al. ( 1959, pp. 8 1-83) have drawn attention to the effects of Pnuidand, thus, of pHlo, which of necessity is =Z Pnuid,being smaller than Plead on the solid phases in contact with it. Such conditions, whereby the pressure is not uniform in all parts of a system, are said to be osmotic. Whenever intergranular spaces containing fluid in a very permeable and porous rock are intercommunicating for an appreciable vertical distance, Pnuid will be lower than Plead, the differential pressure being taken up by grain-grain contacts (Plead = Pnuid Pgrain.to.grain). In the extreme case of a connection to a fissure open to the surface, the Pnuid will be the product of the vertical fluid column in the open fissure and the specific weight of fluid, and thus will be equal to about 1/3 of Plead, assuming 2.8 for the specific gravity of the sedimentary rock column. Fyfe et al. (1958, p. 125; see also Turner and Verhoogen, 1960. p. 22) have derived an expression for the vapor pressure of water in equilibrium with a solid on which the pressure can be varied independently. If the pressure on a solid is increased at constant ‘temperature, its free energy G, which for a pure substance is equal to the chemical potential p , changes by the following amount:

+

dG = d p = &,li&d P . Inasmuch as GP,

=

Colidsundergoes only very small changes with pressure:

+ Kolids( 1‘

- ‘0)

*

If the vapor is to remain in equilibrium with the solid, its chemical potentia1 must change by a similar amount. Inasmuch as for an ideal gas at constant temperature and vapor pressure d p = Kolidsd P = R T d In P, the relationship between P on the solid and its vapor pressurep is as follows: d l n p =-% A i d s d p RT Neglecting the compressibility of the solid phase (i.e. assuming that is constant) and assuming that the vapor behaves as an ideal gas, if the vapor pressure is po when Plead = Po, the vapor pressure p , when the solid phase is under pressure P, is given by the following equation:

v,<,l,d

The corresponding expression for the pressure of water in a dehydration equilibrium is as follows:

416 where fo and f,are the fugacity coefficients of water at water-vapor pressures p o and p I , respectively, and AFolids is the average difference of the volumes of the anhydrous minus the hydrous phase in the pressure range P,,to P,. If A
), might, If the conditions in the rock are osmotic (i.e., Plead B P " ~ , ~ albite therefore, grow in the rock, whereas analcime could grow in an adjoining open fissue at the same low Pnuid, i.e., the relations are opposite to those when the rocks are well sealed, and Pfluidin the rock is much greater than that in the open fissure. In highly permeable rock columns of active geothermal areas, p H Z Omay be approximately equal to 1/3 Plead (Coombs et al., 1959, p. 81). This may account for the low temperatures of the occurrence of wairakite in such areas-140"-25O0C at depths of 180-1 100 m in Wairakei, New Zealand (Coombs et al., 1959; Steiner, 1968), and 75-190°C at depths of 130-600m in Onikobe, northern Honshii (Seki, 1966; Seki et al., 1969b; see Fig. 5-24, this chapter), compared to the experimentally determined thermal-stability ranges of 230"-330°C at 500 bars p H Z Oand 150"-240°C at 1 atm p H , O (Liou, 1970). Equilibrium curves calculated for the laumontite-wairakite reaction by Liou ( 1971a, pp. 392-292) at Paqueous nu,d being equal to 0.8, 0.5, and 0.3 of Ptota,, have progressively steeper positive d P / d T slopes and lie at progressively lower temperatures. On assuming that p H l O= 1/3 Ptotal, the calculated crystallization temperatures of wairakite are about 170"-270°C at Ptotal = 500 bars (Liou, 1970, p. 279; 1971a; see Fig. 5-33, this chapter). In these geothermal areas laumontite appears in the temperature range of 70"-170°C at depths of 60-250m (Seki et al., 1969b; Steiner, 1968),

417 compared to the higher equilibrium temperatures at low pressures for the reaction: heulandite -,laumontite

+ quartz + H,O

(above 250°C for p H z O <1 kbar according to Thompson, when p H z O= Ptotal 1971a, fig. 3; see Fig. 5-32 in this chapter). High content of dissolved components such as NaCl or sulfur, as found in many hydrothermal fluids, may be expected to lower the temperature stability ranges of these zeolites even further through reduction in the chemical activity of water in the fluid phase.

Overlap of zeolite zones related to differences in porosity and permeability. Intensive zeolitization during burial is enhanced by high permeability. This relationship is reflected in the suppression or reduced intensity of zeolitization in fine-grained tuffs, graywackes, and siltstones compared to that in coarser pyroclastics or volcanic sediments (e.g., Otalora, 1964; Hoare et al., 1964; Jolly, 1970; Stewart, 1970, 1974; Stewart and Page, 1974). Many instances of telescoping of authigenic zeolite zones in rocks of similar chemical and mineral composition, however, can also be ascribed to differences in porosity and permeability. The local presence of different, less hydrous zeolite and Ca-Al-silicate minerals in more permeable and porous rocks than in adjoining, less permeable rocks, indicates relative increase in their P and T stability ranges (i.e., decreased equilibrium temperatures of dehydration reactions due to reduced water pressures in the interstices of permeable rocks). Such effects are unlikely to be due merely to differences in flow rate of fluids. They must be ascribed either to different composition of such fluids (to be discussed below), or to incidence of osmotic conditions in the permeable rocks, even considering that the high permeabilities that may give rise to osmotic conditions are likely to be restricted to shallow depths. Drilling data indicate the general prevalence of the P n u , d = P,oadcondition, which is anticipated during progressive dehydration of rocks with low permeability at depths of a few kilometers (cf. Fyfe, 1974). Coombs et al. (1959; see also Hay, 1966, pp. 77-78, 90) have pointed out that heulandite or analcime could be preserved or even form in relatively impermeable, “sealed” beds where pHzO is equal to P,,,,, , whereas in permeable beds under osmotic conditions ( p H z O< Ptota,) the less hydrous laumontite or albite could already form at equivalent levels of burial. This could account for the restriction of the heulandite, which persists into the laumontite zone at Taringatura, Southland (N.Z.), to tightly cemented, impermeable sandstones (Coombs et al., 1959, p.83; Coombs, 1971, p.318), in which a close approach to p H Z O= Plead may be expected. Conversely, alteration of

418 detrital plagioclase to laumontite plus albite occurs locally in the coarsergrained sandstones of the heulandite zone at least 7 km higher in the section than the lowest heulandite (see Fig. 5-22). A similar effect probably accounts for the local interbedding of laumontite-bearing coarse-grained silicic tuffs or tuff-breccias with the mordeniteand clinoptilolite-bearing fine-grained, compact vitric tuffs or pumice tuffs of similar chemical composition in the Izumi Mountain Range, southwestern Japan (Nakajima and Tanaka, 1972) and in the East Tanzawa Mountains, Honshti (Seki et al., 1972). This view is supported by the earlier-mentioned occurrence of laumontite in joints and fissures in clinoptilolite- and mordenite-bearing tuffs, in these and other areas (e.g., Coombs, 1958). Differences in permeability could also contribute towards telescoping of mineral zones through their influence on the flow rate of fluids, which remove CO, evolved in decarbonization reactions. These effects, however, would be opposite to those of osmotic conditions: the decrease in pHz0due to the retention of such CO, would be greater in the more impermeable rocks. As a result of the concomitant decrease in the temperature of dehydration reactions, more dehydrated Al-silicate assemblages might be favored in such impermeable beds, e.g., pumpellyite in laumontite-bearing sequences. At even higher local molar fraction xco, values in the fluids in impermeable rocks, formation of hydrous Ca-Al-silicates may be suppressed entirely, with calcite and aluminous clay mineral appearing instead. Low permeabilities may also cause the local maintenance of high salinities, with their effects on authigenic minerals. Control of zeolite zonation by composition of the fluid phase Reduced Pnuid,discussed above, is not the only factor affecting the chemical activity of water in zeolite reactions. Reduced chemical activity of water in the interstitial fluid ( p H l O< Pfl,jd) due to dissolved components, may have similar effects on zeolite equilibria, even where Pnuid is equal to Ptota,.The effect of reduced pHz0 may be complicated considerably if the dissolved components (which may be either mineral substances such as alkalis or silica, or fluids such as CO,) themselves participate in the equilibrium reactions. Effects of salinity and alkalinity in interstitial solutions on zeolite equilibria Solution of soluble mineral substances reduces the chemical activity of water in the interstitial liquid and, thus, lowers the temperature of dehydration reactions with respect to those in pure water at the same fluid pressure. Campbell and Fyfe (1965, p. 814) have calculated that the decrease in the chemical potential of water upon saturation with NaCl is sufficient to lower

419

+

+

the temperature of the equilibrium analcime quartz e albite water at low pressures from 190°C in pure water to less than 100°C. Other saturated solutions could even achieve stabilization of albite at 25°C. Moreover, increase in the alkalinity of the medium is considered by Senderov (1968) to lower the temperature boundary between analcime and albite by extending the time of recrystallization and, thus, decreasing the tendency of analcime to crystallize as a metastable phase within the albite field. Liou (1971a, pp. 390-391) has shown experimentally that the upper thermal stability limit of laumontite (i.e., the equilibrium. temperature of the reaction laumontite * wairakite water) is reduced by 10"- 15°C at 3000 bars in equilibrium with a 0.25-N NaCl solution compared to excess pure water.

+

Imposition of zeolite zonation by salinity and alkalinity gradients in trapped lake waters Progressive downward increase in the chemical potential of water, leading to a dehydration sequence of zeolites, may be incident upon increasing salinity of interstitial aqueous solutions in rocks that have never been deeply buried. Several examples of zeolite zonation at shallow depth, reflecting gradients of salinity in interstitial waters, are known from lacustrine acid-tuff deposits of Cenozoic age in the western United States. These include (1) the Eocene Green River Formation of Wyoming (Hay, 1966, pp. 44-52, 90-93; Goodwin and Surdam, 1967; Roehler, 1972; Surdam and Parker, 1972), (2) the Miocene Barstow Formation (Sheppard and Gude, 1969), and (3) the deposits of the Pleistocene Lake Tecopa, eastern California (Sheppard and Gude, 1968). These deposits show a succession from alkali-zeolite tuffs (mainly clinoptilolite and mordenite, but locally also phillipsite, erionite, etc.), which are characterized by a vitroclastic structure and relic glass, to analcime and/or alkali-feldspar zones, usually towards the central parts of the basin. Locally (e.g., in the Lake Tecopa deposit of eastern California), a fresh-glass zone is present near the lake margins and inlets. The vitroclastic structure and relic glass are rarely preserved in association with the silica-rich analcime and the K-feldspar. The latter are considered to have formed from diagenetic alkalizeolite precursors, rather than directly from the rhyolitic glass, under the influence of hypersaline and alkaline waters (see also Iijima and Hay, 1968). The preservation of the original textures and sedimentary structures in the altered tuffs proves them to be lateral equivalent of the fresh-glass tuffs (where present). This also constitutes evidence that the present differences in composition and mineralogy within these tuffs do not reflect primary differences, but are due to postdepositional processes.

420 These processes took place under the influence of the chemical zonation of the trapped lake waters, from the relatively fresh near the margins and inlets of the lakes to highly saline, with a p H of 9 or more, basinwards. The chemical zonation of the trapped water reflects paleosalinity and paleoalkalinity gradients inherited from the lake environment. Temperature and pressure must have been uniformly low throughout, at least in the Lake Tecopa deposit, the base of which was never buried deeper than about 100 m. A vertical mineral zonation from more hydrous to less hydrous alkalialuminosilicates, similar to mineral zonation in these lacustrine deposits that have never been deeply buried, would be favored by the common increase in the salinity of interstratal waters with depth in sedimentary basins (Hay, 1966, p. 89; Coombs, 1971, pp. 322-323). The possibility of partial control or enhancement of vertical compositional gradients in the interstitial waters in volcanic-clastic deposits by a downward increase in the extent of alteration and solution of unstable clastic relics. i.e., by kinetic factors, is considered after discussing the effects of oversaturation of fluids in silica on zeolite equilibria. Effects of oversaturation of fluids in silica with respect to quartz In the low-temperature environments prevailing in burial metamorphism, equilibrium with silica normally implies equilibrium with quartz. The siliceous alkali-zeolite clinoptilolite in acidic tuffs and bentonites, however, is persistently associated with tridymite, cristobalite or opal (e.g., Iijima, 1961; Reynolds and Anderson, 1967; Reynolds, 1970; Cameron and Sabine, 1969; Brown et al., 1969; Hallam and Snellwood, 1970). In many of these rocks, clinoptilolite (and/or mordenite) may have formed very early during diagenesis, often at the expense of silicic volcanic glass by hydrolysis and solution. The association of these silicic alkali-zeolites with an unstable form of silica, i.e., in an environment oversaturated in silica with respect to quartz, may represent metastable equilibrium. In any reaction 01the type: A,B,C, e A,B,C,-,

+ WC

the phase A,B,C, will form over a wider range of conditions if the state of C is not the stable form at the given pressure and temperature. Consequently, the effect of an unstable form of silica, such as glass or cristobalite, on a low-temperature reaction of the type: A(SiO,),

+ 2 SiOz e A(SiO,),

will be the expansion of the P-T stability field of the more silica-rich phase A(SiO,), (Coombs et al., 1959, pp. 77-78). I n the earlier-mentioned zonal distribution of zeolites in the acid and

42 I

intermediate Neogene pyroclastics of the Shinjo Basin, northern Honshu (see Fig. 5-31), Utada (1965) and Iijima and Utada (1966) noted that the upper zone I is characterized by fresh glass; the clinoptilolitic-mordenite zone 11, by opaline silica (low-cristobalite) and the analcime-heulandite zone 111 and the laumontite zone IV, by chalcedonic quartz. Mizutani (1967, 1970) has related this zeolite zoning to the diagenetic conversion of amorphous silica during progressive burial through low-cristobalite to low-quartz (see Fig. 5-34).In dehydration reactions, the least hydrated phase will generally be the high-temperature phase, in view of the positive A S of such reactions. Oversaturation of fluids in silica with respect to quart; will thus tend to increase the equilibrium temperature of dehydration reactions which release silica, such as: mordenite -, analcime SiO, water,

+

+

but will favor lower equilibrium temperatures of dehydration reactions requiring silica, such as: analcime

+ SiO,

-, albite

+ water.

In the latter case, oversaturation will favor lower dehydration temperatures (e.g., Campbell and Fyfe, 1965, p. 914; Hay, 1966, p. 94),especially because the A S of this reaction is very small compared to that of other zeolitic dehydration reactions. Thick- Age of ness deposition m m.y. BP

Mineralogical composition of silica (SiO2 1 -t

Zonal distribution of silica minerals in Mogami district a f t e r UTADA (1965)

0

.-

.-16

Zonen clinoptiloiitemordenite zone

CI 0

n

20 hculandite zone

Fig. 5-34. Mineralogical composition of silica in the stratigraphic column of Neogene pyroclastic rocks of the Mogami district, Shinjo Basin, northern Honshii as computed from the kinetic parameters and the thermal history of the succession (to the left) compared with the zonal distribution of silica minerals after Utada (1965; see Fig. 5-31 in this chapter). (After Mizutani, 1970, fig. 6.)

422 The effect of oversaturation in silica thus restricts the occurrence of analcime. Inasmuch as the composite dehydration reaction: mordenite (or clinoptilolite)

+ealbite

+ SiO, + water

releases silica, oversaturation in silica would relatively increase the stability of the alkali zeolites. The same would be true for heulandite in the following important dehydration reaction: heulandite * laumontite

+ SiOz + water.

Local oversaturation in silica, caused by the presence of volcanic glass, might sufficiently expand the P-T stability field of (1) the alkali-zeolites to account for their local occurrence in pumice tuffs within laumontite-bearing crystalline tuff-breccias (as reported from the eastern Tanzawa Mountains of central Honshu by Seki et al. (1972), and (2) of heulandite to explain its persistence in glass-bearing rocks into the laumontite zone (e.g., Read and Eisbacher, 1974). Imposition of zeolite zoning by ionic-activity gradients controlled by hydrolysis and solution of acid volcanic glass upon burial In the preceding sections, some sequences were reported in which the zonal distribution of zeolites appeared to be due to gradients in the chemical activities of appropriate components, such as alkalis and silica, rather than to the increase in temperature and pressure with burial depth. The salinity and alkalinity gradients, with the concomitant variations in the chemical activity of water, were externally imposed upon the solid phases by preexisting, fossil compositional gradients in trapped interstitial fluids (surface waters). The material necessary for the formation of the zeolites and K-feldspars in the lacustrine acid-tuff deposits, however, is largely provided by acid volcanic glass. In other cases it is also provided by clastic feldspars. Such zeolitization involves hydrolysis and solution followed by precipitation of the zeolite from solution, rather than mere devitrification or hydration. This is demonstrated by the common occurrence of zeolites around and along the cleavages, rather than within the clastic grains. Solution of acid volcanic glass. The sequence of formation of authigenic minerals within the zeolite facies of the Lake Tecopa deposits is as follows: montmorillonite, phillipsite, and then clinoptilolite or erionite. This sequence was attributed by Sheppard and Gude (1968, pp. 34-35) to the progressive solution of rhyolitic glass by alkaline and saline pore waters. After early formation of montmorillonite, increase in the (Nat K + ) / H activity ratio through solution and hydrolysis of the glass created an environment more

+

+

423 suitable for the formation of phillipsite than of additional montmorillonite; subsequent increases in the SiO, activity and the relative (Na Ca Mg)/K ratio in the pore waters led to formation of the K-poorer and Si0,-richer erionite and clinoptilolite. This suggests the possibility that vertical salinity and alkalinity gradients. which control depth zoning in zeolite sequences, could be controlled or enhanced from within the system by a progressive downward increase in the extent of hydrolysis and solution of the unstable solid phases, i.e., by kinetic factors. Such compositional gradients could parallel the chemical-activity gradients of SiO,. The latter are considered to reflect oversaturation at shallow burial depth due to the presence of unstable forms of silica (glass, cristobalite, or opal) and their gradual diagenetic conversion to the stable form. quartz, upon progressive burial. Such vertical alkalinity, salinity, and .silica activity gradients of ground waters are reflected by the stratigraphically-zoned succession of abundant zeolites and K-feldspar in the volcanic sediments and tyffs of the MiocenePliocene Esmeralda Formation of southwestern Nevada (Robinson. 1966; Robinson et al., 1968; Moiola, 1970). In the 2800-m thick sequence of the type locality, phillipsite is found only in basaltic tuff-breccias in a thin zone (< 150m) at the top. It is followed downwards by a 2300-m thick zone in which clinoptilolite- usually accompanied by opal showing cristobalite X-ray diffraction peaks-is particularly abundant as a replacement of glass in tuffs and tuffaceous sandstones, and in the upper part of which fresh glass is still present. Analcime and K-feldspar are confined to the lower 150 m near the base of the sequence. This zoning is ascribed by Moiola (1970) primarily to an increase in the (Na+ K + ) / H activity ratio in the subsurface waters with increasing depth, as a result of the proceeding hydrolysis and solution of vitric material after burial. The precipitation of opal with clinoptilolite reflects supersaturation of the solutions with silica. These and similar successions of clinoptilolite-mordenite, analcime, and K-feldspar zones reflect gradients in the chemical activity of silica and alkali ions (Na+ and K + ) imposed in this case not by the environment during sedimentation, but by progressive solution of reactive clastic constituents upon shallow burial. Though such zeolite zoning is not primarily the result of successive stability at increasing temperature with depth, it is conceivable that this factor indirectly contributes to the zeolite zoning by enhancing the solution reactions. Nevertheless, the above mineral zones cannot be regarded as facies of burial metamorphism: their occurrence is controlled by non-equilibrium solution reactions.

+ +

+

+

424 Extension of zonal distribution ranges of diagnostic zeolites and hydrous Ca-A1-silicates due to ionic equilibria and coupled solid-solution reactions In the above considerations, the various minerals participating in the low-grade reactions have been regarded as having constant chemical composition. Many of the zeolites and other hydrous Ca-Al-silicates involved in these reactions, however, show marked departures from the stoichiometric compositions used due to ionic substitutions and solid solution. Whenever one or more of the phases participating in an equilibrium has a variable chemical composition in the system under consideration, the equilibrium becomes bivariant, i.e., the reaction takes place over a range of temperatures at fixed pressure. In the case of incomplete solid solution, the two minerals may coexist over a range of P-T conditions. Variation in composition of analcime. Analcimes found in sedimentary and burial-metamorphic rocks are commonly richer in silicon and poorer in aluminium and sodium, than the stoichiometric composition,NaAlSi ,O, H,O. Compositions of analcimes from quartz-bearing burial-metamorphic rocks cluster around NaAlSi,;O, - laH,O ( e g , Coombs et al., 1959; Wilkinson and Whetten, 1964; Nakajima and Koizumi, 1966; Coombs and Whetten, 1967). Inasmuch as the analcimes form a complete solid solution series, the following hypothetical dehydration-solid-solution equilibrium becomes bivariant: +

NaAlSi,O,. 1 5 H,O #NaAlSi,O, Si-rich analcime

. H,O

+ SiO, + H,O.

Si-poor “ideal” analcime

The Si-content of analcime coexisting with free quartz should vary as a function not only of T and p H I O but , also as a function of the chemical activities of sodium, alumina, and silica. Both the above reaction and the coupled reaction: Si-poor “ideal” analcime

+ quartz e albite + water

have small positive entropies. Consequently, the equilibrium temperature of the following sum reaction must be even lower:

+

Si-rich analcime e albite water. The Si content of analcimes coexisting with quartz should thus decrease with increasing temperature at constant pressure (Coombs and Whetten, 1967, p. 277). Conversely, the maximum temperature at which Si-rich analcime is present in the presence of quartz should be lower than that determined experimentally by Campbell and Fyfe (1965) and Liou (1971b) for the “ideal” composition NaAlSi,O, . H,O.

425 Nakajima and Koizumi (1966) have suggested that the %/A1 ratio of analcimes coexisting with quartz varies with depth of burial and can be used as a geological thermometer. Coombs and Whetten (1967, p. 278), however, have disputed this view, showing that five analcimes occurring through a 4600-m thick burial-metamorphic sequence in the Taringatura district, southern New Zealand, have very similar compositions. The large differences in composition of analcimes from approximately the same stratigraphic horizon of the earlier-mentioned Green River Formation of Wyoming reflect variations in the parent materials. The Si/Al ratio varies inversely both with the salinity of interstitial solutions and with the amount of alkali feldspar present, due to the progressive desilication of the less stable high-silica analcime in alkali-feldspar-forming reactions (Iijima and Hay, 1968; see also Coombs and Whetten, 1967; and Liou, 1971b). It may be concluded that, although comparatively high temperature would favour low %/A1 ratios, analcime composition is probably of little value as a geothermometer in burial-metamorphic rocks. Influence of sodium substition on Cu-zeolite equilibria in the laumontite zone. Many zeolites show various extents of coupled substitution of Na+ Si4+ for Ca2+A13+,interstitial substitution of 2 Na+ for C a 2 + , or of NafAl'+ for S i 4 + . The influence of variation in composition on the temperatures of zeolite equilibria will be particularly marked when major substitution is possible in one of the reactants but not in the other. For instance, virtually all of the above ionic substitutions are present in the clinoptilolite-heulandite solid-solution series, whereas Na-substitution in laumontite is negligible. The following hydration and solid-solution reactions can be regarded as coupled reactions, with the equilibrium temperature TI > T,: (1) laumontite

+ quartz + water -, heulandite (large negative A S , ),

(2) heulandite

+ clinoptilolite

-+

heulandite-clinoptilolites, .

(small positive A S 2 )

The coupled hydration-solid-solution reaction: (3) laumontite clinoptilolite quartz water -,heulandite-clinoptilolite,,

+

+

+

+

will then have a negative AS, ( = AS, AS,) and an equilibrium temperature T, > TI (cf. case 4 in Turner, 1968, pp. 74-75 and fig. 2-3b). Consequently, heulandite-clinoptilolite,, may be stable above the upper temperature limit of heulandite (TI) in a temperature range TI-T,, in which it can coexist with laumontite. Thus, increased Na' activity - and the resulting substitution of Na+Si4+ for Ca2+A13+ in heulandite - will result in an

426 overlap between the stability fields of heulandite-clinoptilolite,, and that of laumontite. An analogous increase in Ca content appears to characterize the mordenites from acidic tuffs, which persist into the laumontite zone, compared to those in lower-grade zones (Nakajima and Tanaka, 1967; Seki, 1973b). Seki (1973b) considers that in the presence of high silica activities and high partial pressures of water, mordenite with intermediate Ca/Na ratios could persist in the laumontite zone beyond the thermal stability of clinoptilolite, stilbite, and heulandite. A somewhat similar effect has been reported for clinoptilolites from nothern Akita prefecture, northern Honshii, by Iijima (1974): Na-K type in the clinoptilolite-mordenite zone 11, and Ca-type into the analcime-calcite zone 111. On the other hand, Boles and Coombs (1975) have shown that the %/A1 ratios of the heulandites (and clinoptilolites) in the now classical 4.8-8.5 km thick section of heulandite-, analcime-, and laumontite-altered andesitic and rhyolitic tuffs in the Hokonui and Taringatura Hills, southern New Zealand, were largely controlled by the Si/Al ratios of their glass precuxsors. They show no correlation with depth of burial, which is reflected by (1) the absence of analcime in the lower part of the section, and (2) the abundance of laumontite to the exclusion of both heulandite and analcime in the 2-km thick volcanogenic sandstone underlying the tuffs. Some bulk-composition control is suggested by the restriction of the formation of laumontite and prehnite in the tuffs-apparently at the expense of heulandite-to the andesitic intervals. On the high-temperature side of the laumontite stability range, a similar relationship exists with wairakite, the Ca-analogue of analcime. Wairakite shows a major substitution of 2 Na+ for Ca2+, whereas the Na-substitution in laumontite remains negligible (e.g., Surdam, 1973). The following dehydration and solid solution reactions both have a positive AS, (with AS, > AS,), with the equilibrium temperature T , > T2:

( I ) laumontite + wairakite (2) wairakite

+ analcime

+

+ water wairakite-analcime,, .

Thus, the coupled dehydration-solid solution reaction: (3) laumontite

+ analcime -,wairakite-analcimes, + water +

must also have a positive AS, ( = A S l AS,), with the equilibrium temperature T3< T I (case 2 of Turner, 1968, p. 72 and fig. 2-2b). High activity of Na+ should thus decrease the crystallization temperature of Na-bearing wairakite and widen its p H l o - T stability field to overlap that of laumontite. Whereas it would appear to be impossible for Na-free wairakite to coexist stably with laumontite over a range of burial temperatures and pressures,

427 stable association of Na-bearing wairakite and laumontite may be thus possible because of the addition of another component to the system (Liou, 1970, p. 278). This may explain the overlap and local coexistence of the two minerals in the laumontite zones of (1) the Karmutsen Group of Vancouver Island, British Columbia (Surdam, 1973); (2) the Miocene volcanics of the Tanzawa Mountains and the Fujikawa district, central Honshu (Seki et al., 1969a; Shimazu et al., 1971; Seki, 1971); and (3) over a depth range of up to 100m in the hydrothermal wells of Wairakei, New Zealand (Coombs et al., 1959; Steiner, 1968), and Onikobe, northern Honshii (Seki et al., 1969b). The net effect of high Nai activity, therefore, will consist in the widening of the occurrence range of heulandite-clinoptilolite,, in the low-temperature, and of wairakite-analcime,, in the high-temperature part of the laumontite zone. Ionic activities in solution and the stabilities of laumontite, wairakite, prehnite, pumpellyite, and epidote. In many low-grade metamorphic sequences, the gross mineral zoning of the above diagnostic Ca-Al-silicate minerals reflects depth of burial and thermal stability of the minerals, whereas in detail their occurrences may overlap over considerable intervals. Such sequences include (1) the classical Taringatura section of southern New Zealand, and (2) the 5500-m thick sequence of mainly tholeiitic tuffs and volcanogenic rocks of the Karmutsen Group of Vancouver Island. Surdam (1973) found that in the latter sequence laumontite and prehnite overlap over more than 3000 stratigraphic meters; in the lower part of this interval pumpellyite and epidote also occur (see Fig. 5-35, after Surdam, 1973, fig. 3). Surdam (1973) explained the overlap in the occurrence of these minerals in terms of ionic equilibria in solution. Using the stoichiometric compositions, he showed that within an appropriate range of temperature and pressure, and in the presence of water, any of five Ca-Al-silicates could be stable, depending on the activities of C a 2 + , SiO,, and H + . The phase relationships of the stability fields of (1) laumontite or wairakite, (2) prehnite, and (3) pumpellyite or epidote are depicted on a schematic aca2+/a(,+p vs. as,O, activity diagram (Fig. 5-36, after Surdam, 1973, fig. 6). The overlap between the stability ranges of wairakite and laumontite, which are identical stoichiometrically except for water content, may be accounted for by small differences in the activities of Na' and Ca'* due to substitution of Na+ for Ca2+ in wairakite but not in laumontite, as discussed above. A similar relationship may exist between Na-bearing wairakite and prehnite. It is noteworthy that both in the north Tanzawa Mountains and the Fujikawa district of the south Fossa Magna area the wairakite solid-solutions associated with prehnite are markedly richer in the analcime component than those associated with epidote (Seki, 1971).

428 Southland

New Zealand (Coombs, 1954 1 Albitized Plag.

I

Pumpelly ite Epidote Prehnite Laumontite Heulondite Anoldme

Butte Lake Albitized Plag. Pumpellyite Epidote Prehnite Loumontite Analcime Albite

-

0

5000 I t

Fig. 5-35. Authigenic mineralogy of tuffaceous rocks from the Karmutsen Group, Butte Lake area, Vancouver Island (British Columbia) compared with that of the Triassic of the Taringatura Hills, Southland, New Zealand (Coombs, 1954). The reference for the comparison is the analcime-albite transition. There is 5,ooO-15,000 ft of overburden on top of the Karmutsen Group. (After Surdam, 1973, fig. 3.)

The overlap of the ranges of piimpellyite and epidote (the iron-free compositions of which are stoichiometrically identical, except for water content), could partly be due to differences of aMgZ+and aFcZ4and the oxidation-reduction potential in the aqueous phase. Substitution of Fe in Ca-Al-silicates. Of the major diagnostic burialmetamorphic minerals, laumontite is one of the very few to have a more or less constant chemical composition. Its range of occurrence, therefore, cannot be extended by high activities of cations such as Na’ , Fe’+, F e 3 + ,etc. These may, however, result in its displacement by wairakite, prehnite, pumpellyite, or even epidote. This justifies Coombs’ (1961) assignment of

429

log %io2

-

Fig. 5-36. Chemical activity diagram depicting phase relations of hydrous Ca-Al-silicate minerals at an unspecified temperature and pressure in the presenFe of water after Surdam (1973, fig. 6). Assuming that the phases are stoichiometric, wairakite and laumontite are the same, except for water content, as are epidote and pumpellyite.

those zones of burial metamorphism in which both laumontite and prehnite and/or pumpellyite occur to the laumontite zone of the zeolite facies, rather than to the prehnite-pumpellyite-metagraywacke facies. The Ca-Al-silicates diagnostic of the more advanced stages of burial metamorphism, including epidote, prehnite, and pumpellyi te, show appreciably substitution of A1 by F e 3 + , and of Mg by FeZf in pumpellyite. Moreover, one or both of these substitutions are prevalent in other common minerals in pumpellyite-bearing facies, such as chlorite and actinolite. Introduction of the component FeO in the experimentally studied system Ca0-Mg0-A1,0,-Si02-H,0 will thus usually lead to Mg-Fe2+ substitution in one of these phases rather than to the appearance of a new phase. Consequently, the univariant equilibrium curves as determined in the FeOfree system will be controlled by the Fe/Mg ratio in addition to pressure and temperature, and thus become bivariant (cf. Nitsch, 1971, p. 253). Thompson (1970b) has pointed out that presence of Fe and Mn might greatly reduce the pressure-stability of pumpellyite (i.e., allow its formation at much lower pressures) with respect to that found for Fe-free pumpellyite by Hinrichsen and Schiirmann ( 1969). Hashimoto (1968, p. 132) found markedly higher Fe contents in the pumpellyites of the prehnite-pumpellyite-metagraywacke facies of the Katsuyama district of southwestern Honshu than in those of the pumpel-

430 lyite-actinolite schists of the higher-grade glaucophane-schist zone (cf. Coombs et al., 1970). The epidotes in prehnite-pumpellyite zones have higher Fe-Mg ratios than pumpellyites from the same intervals, for instance in the Karmutsen Group (Surdam, 1973, p. 1918), which suggests that high Fe/Mg activity ratios may be a factor in their stabilization versus pumpellyite. Epidotes from prehnite-pumpellyite zones also appear to be markedly richer in Fe than epidotes from higher-grade zones (e.g., Hashimoto, 1968, p. 133; Bishop, 1972b, p. 3184; Coombs et al., 1976). Bishop (1972b, p. 3191) has thus suggested that with progressive metamorphism pumpellyite should also disappear at lower temperatures in Fe-rich systems than in the Fe-free systems studied by Nitsch (1971). This may be due to a reduced equilibrium temperature of epidote-forming reactions such as : pumpellyite chlorite quartz e epidote actinolite water.

+

+

+

+

Although Fe-rich epidote is common in the pumpellyite-actinolite-schist facies, it may already appear-although generally in very minorlquantitiesin the prehnite-pumpellyite-metagraywacke facies sensu strict0 (cf. Packham and Crook, 1961; Martini and Vuagnat, 1965, p.285; Matsuda and Kuriyagawa, 1965; Smith, 1968; Hsshimoto, 1968; Seki et al., 1969a; Seki et al., 1971; Surdam, 1973; and some other references in Seki, 1972). Fe-rich epidote may locally even appear in the higher-grade part of the laumontite zone, e g , in the Taringatura section and in the Tanzawa Mountains (Seki et al., 1969a). Moreover, inasmuch as the iron of low-grade epidotes is predominantly in the trivalent state (Coombs, 1961; Strens, 1965), it is to be expected that high Fe3+/FeZf activity ratios (i.e., high oxidation state) will extend the P-T stability range of epidote at the expense of pumpellyite (cf. Thompson, 1970b, p.274; Seki, 1972). This could explain the sporadic occurrence of pumpellyite, compared with the common presence of the epidote-actinolite assemblage, in hematitic metavolcanic rocks of the pumpellyite-actinoliteschist facies in northern Otago, New Zealand (Bishop, 1972b, p. 3191). Iron-poor zoisite or clinozoisite appears only in the greenschist facies or “low-stage metamorphism” of Winkler ( 1970).

Solution of basaltic glass and calcic plagioclase. Enrichment of interstitial waters in alkali ions (Na’ and K t )and silica as a result of hydrolysis and solution of vitric material are also in accordance with the observations on the chemical changes upon alteration of basaltic glass. Relative losses of Si and alkali, and relative gains in Fe, Mg, and, in‘some cases, A1 were found upon the palagonitization of Quaternary basaltic glass (sideromelane). This occurred through reaction with groundwaters in Hawaii (Hay and Iijima, 1968)

43 1 and through reaction with sea water near Surtsey, Iceland (Alexandersson, 1970), the released components precipitating as zeolites, opal, and other minerals. The changes in composition of the sideromelane are not necessarily identical in all stages of the alteration process. For instance, the cores of the 14 A phyllosilicate fragments, which are the alteration products of the basaltic glass shards and fragments in the tuffs of the Triassic Karmutsen Group of Vancouver Island, show significant relative losses of Si and Na, and slight gains in Fe, Al, and Mg in comparison with the chemical composition of the average feeder dike (Surdam, 1973). The more altered rims of the fragments, in addition show relative loss of Ca, complete loss of K, and relative gains of Fe and Mg (Surdam, 1973). In the upper part of the sequence, this altered glass is surrounded by laumontite, while part of the plagioclase remains fresh and unaltered. In the lowermost part, the glass has been altered directly to prehnite and phyllosilicate, whereas the calcic plagioclase has been largely decomposed to pseudomorphs of albite and hydrous Ca-Al-silicates, generally pumpellyite, but locally prehni te, laumontite, wairakite, or epidote. Surdam (1973) ascribed the differences in activity ratios of ionic species, whch controlled the vertical overlap of these critical Ca-Al-minerals, to hydration and differential solution of volcanic glass and plagioclase after burial. Mobility of stratal waters of varying composition. Movement of stratal waters of different ionic composition through a sedimentary pile may lead to different replacements, with concomitant change in bulk composition of originally similar rocks. Dickinson (1962b) has drawn attention to the role of pore waters in the albitization of plagioclase and the formation of pumpellyite, prehnite, and laumontite (at the expense of diagenetic heulandite) during genesis of a Ca-Al-silicate zoning unrelated to depth of burial. The water acted both as a catalytic agent and as a carrier of critical components on whose mobility the alteration reactions depended. Otalora (1964, pp. 730-73 1) attributes the overlap in the vertical distribution of analcime and laumontite throughout the Barranquitas tuff member in east-central Puerto Rico (where the minerals do not occur together, though about 75% of the samples investigated contains either mineral) to formation of analcime as a result of interaction of connate water with clay minerals in the matrix, whereas laumontite formed by interaction of Ca and A1 released by albitization of calcic plagioclase and clay minerals. In the rhyolitic to andesitic tuffs of the Hokonui-Taringatura Hills area of southern New Zealand, Boles and Coombs (1975) found that heulanditeand laumontite-altered tuffs were Ca-enriched and Na-depleted, whereas

432

analcime-bearing tuffs were Na-enriched relative to the unaltered volcanic rocks. According to these authors the breakdown of heulandite, which takes place locally but over a stratigraphic interval of several kilometers, is favored by a local lowering of a H Z Oor aao,. The latter is brought about by movement of stratal waters of varying composition through the sedimentary pile. The alternative breakdown of the heulandite to Ca-Al-silicates (laumontite or prehnite) or to Na-Al-silicates (analcime or albite) is attributed to the varying aNa+ /a,,> + and a N a+ /aH ratios in these waters. +

Zeolite zoning due to variation of the pcq,/p,,,o ratio in the fluid phase CO, content of intergranular fluids will causepHZoto be lower than Pr,,,,d. Zen ( 1961) has introduced diagrams showing the stability of Ca-zeolites, Al-phyllosilicates, and calcite in the presence of quartz in the pseudo-binary system CaO-Al,O, as a function of the chemical potentials of H,O and CO,. Assuming that all phases have constant composition, the trivariant reactions will be univariant when T and Ptotal are fixed and, appear as lines on such diagrams. For reactions in which both volatiles participate, these univariant lines will have a positive GpH20/SpC02slope if the volatiles are evolved on opposite sides of the reaction equation, and a negative GpHzo/SpcOzslope if evolved on the same side, the slopes flattening as more water is involved compared to CO,. For reactions that involve H,O but no CO, ( GpH~o/Sp,oz = 0), the reaction lines will be parallel to the pco2 coordinate. In the case of such reactions, an increase in the molar fraction xcoz at constant Pnu,d will, thus, merely affect the equilibrium conditions as any inert solute such as nitrogen or sugar would (i.e., through the concomitant reduction of p H z Oand hence of pHz0), thus depressing the temperature of facies reactions that involve dehydration. Such reactions, which thus proceed to the right as pH,o is decreased, include: (a) heulandite -+ laumontite (b) laumontite

(c) prehnite

+

prehnite

+ quartz + H,O,

+ “kaolinite”+quartz + H,O, and

+ kaolinite -, zoisite + quartz + H,O.

Albee and Zen (1969, especially figs. 6 and 7) have amplified Zen’s (1961) isothermal, isobaric p H z 0-pro, diagrams for quartz-bearing assemblages with heulandite, laumontite, kaolinite, pyrophyllite, and calcite, into the condensed pseudo-ternary 6-phase multisystem SO,-A1 ,03-( Fe, Mg, Mn)OCaO-H,O-CO,. The latter includes the additional phases prehnite. pumpellyite, chlorite, and Ca-montmorillonite (Fig, 5-37, after Albee and Zen, 1969, fig. 7). This diagram shows that the following progressive sequence may

433

P

Fig. 5-37. Synoptic pco,-pHz0 diagram showing possible quartz-bearing assemblages in the complex pseudo-ternaj six-phase multisystem SiOl -A1,0, -CaO-(Fe, Mg, Mn)O-H,OCO, at some constant temperature and load pressure; all phases are assumed to have fixed chemical composition. A potassium mica (assumed to be the only K-bearing phase) may be present. The albite-analcime transition (taken to be closely associated with the heulanditelaumontite transition) is schematically indicated by the hachured zones. Coordinates are equally spaced on each axis. The Al,O,-CaO-(Fe, Mg)O projections of the mineral compositions are shown in the inset. (After Albee and Zen, 1969, figs. 3 and 7.)

occur at decreasing pH,o at low valus of pco2 and some arbitrary, constant value of T and Pnuid. ( 1 ) Appearance of prehnite (at the expense of Ca-zeolite calcite), followed closely by the appearance of montmorillonite (the order might be reversed).

+

434

(2) Heulandite -, laumontite reaction. (3) First appearance of pumpellyite (from laumontite + prehnite chlorite). (4) Final disappearance of laumontite (upon formation of prehnite Alrich clay mineral). ( 5 ) Disappearance of pumpellyite followed by disappearance of montmorillonite (this order may be reversed). This sequence corresponds to that observed in various burial sequences and is commonly attributed to increasing temperature. Of the above reactions at low pco,, the only one to involve CO,, and thus to be dependent on pcoz, is the formation of prehnite from laumontite and calcite according to the following reaction:

+

+

laumontite

+ calcite -, prehnite + quartz + 3 H 2 0+ CO,

(with a negative GpH,,/Spco, ratio of - 5). Such a sequence could thus also be due to a decrease in pHz0because of increasing salinity’at low values of El. co; In considering the effect of variations in the molar fraction xcoz at constant Pfluid and T, it should be remembered that the chemical potentials p H Z 0and pco, are logarithmic functions of the fugacities (or activities) of H,O and CO,. Thus, for the very low xco, range of the Ca-Al-silicate equilibria (cf. next section), one can take the line representingpHzo+pcoz = Pflu,d for binary H20-CO, fluids (the “binary vapor line” of Kerrick, 1974, pp. 739-740) to be subparallel to the pHz0coordinate. Within the T, Pfluid., and pHZ0ranges of the high-grade part of the laumontite zone in which prehnite and pumpellyite may occur, increase in xco, will result successively in the following carbonate-forming reactions at the expense of Ca-Al-silicate phases (see Fig. 5-37): (1) Incompatibility of prehnite and chlorite23 p r + 2 c h + 4 H 2 O + 6 C O , + 10 p u + 6 c a + 15qz (prehnite without chlorite; pumpellyite; and laumontite are stable). (2) Final disappearance of prehnitepr qz 3 H,O CO, lu ca (pumpellyite and laumontite are stable). (3) Final disappearance of pumpellyite (laumontite stable; no prehnite). (4) Final disappearance of laumontitelu+CO, + k a + c a + 2 q z (no prehnite or pumpellyite). In a very narrow range of slightly lower pHlo values, laumontite could disappear before pumpellyite with increasing xco,. High pco, values, therefore, will restrict the appearance of prehnite in the laumontite zone, but

+ +

+

+

+

435 extend the range of compositions containing pumpellyite, at least between reactions (1) and (3). The assemblage prehnite-pumpellyite-chlorite without laumontite may occur in the low-pco2 range in some bulk compositions, but laumontite will occur in others. However, at the same T and Pnuid,and similarly low pco, values, but in the 1ower-pH2, range (i.e., due to high salinity) laumontite cannot form, .or will disappear through the following reaction: 2 lu -+ pr ka 3 qz 5 H,O.

+ +

+

Under such conditions, the diagnostic prehnite-pumpellyite-chlorite assemblage, though not forming in a wider range of bulk compositions, will define the facies as prehnite-pumpellyite, in view of the absence of associated laumontite-bearing assemblages. At somewhat higher grades of burial (or much lower pH20values) epidote appears to be compatible with higher xco2 values than pumpellyite (cf. Fig. 5-38). Surdam (1973, p. 1918) has found (1) a close association between calcite and epidote, and (2) correlation between the presence of calcite and the absence of1 pumpellyite in the alteration products of tuffaceous rocks of the Karmutsen Group, Vancouver Island. The pH20-pco, diagrams indicate that the various diagnostic mineral assemblages of the zeolite facies and the prehnite-pumpellyite facies may in principle be obtained isothermally and isobarically by progressively decreasing pH20with respect to pco2 (cf. Fig. 5-38). There is little doubt, however, that the regional transition from zeolite facies through prehnite-pumpellyite facies and pumpellyite-actinolite facies to the greenschist facies in New Zealand and Japan reflects increasing temperature. In a series of analogous p H 2 0 -p co2diagrams representing increasing temperatures, the field would be progressively restricted towards the facies assemblages in the lower left of the diagrams (see also Kerrick, 1974, p.739). At a pressure of 2kbar, for example, the heulandite-, laumontite- and prehnite-bearing assemblages would be eliminated above about 2OO0C, at 3OO0C, and at 400°C, respectively (Coombs et al., 1970, p. 152). Suppression of hydrous Ca-AI-silicates at high pco, The pH,0-pco2 diagrams indicate that at the T and Ptota,at which variations of pH20 could give rise to the diagnostic Ca-Al-silicate assemblages of the various lowest-grade facies, the aluminous clay minerals-chlorite-calcite-quartz assemblages remain stable over a similar range of pH,o, but at higher values of pco2. In the presence of high values of xco2 in the fluid phase, the hydrous Ca-Al-silicates will be displaced entirely by Ca-free aluminous layer-silicates calcite, and the zeolite and prehnite-pumpellyite -metagraywacke facies will not appear at all.

+

-

PY

ZPGC

I

PCO2

-

Fig. 5-38. Schematic pHzo-pcO, diagram for part of the system A120,-Ca0 in the presence of excess quartz, at some arbihary value of T and total P. The precise position of the kaolinite-pyrophyllite reaction, relative to the other reactions shown, is uncertain. C= calcite: G = grossular; H = heulandite; K = kaolinite; L = laumontite; P =prehnite; P Y = pyrophyllite; Z=zoisite. In any divariant assemblage, two of these phases can co-exist as indicated by adjacent symbols in the phase compatibility diagrams. For example, in the uppermost field, i.e.. above the univariant lines 2, I , and 8, the phases kaolinite, heulandite, and calcite are stable where the ratio of Al,O,/CaO is 1 :0, I : 1, and 0 : I , respectively. Kaolinite and heulandite co-exist for finite ratios AI,O,/CaO> 1 : 1. (From Coombs et al., 1970, fig. 2, as reproduced in Coombs, 197I , fig. 2.)

In zeolite-bearing sequences, zeolitization of calcic plagioclase is commonly absent in the calcite-rich rocks of otherwise suitable non-volatile composition, e.g., in the calcareous unit IV in the, Aldrich Mountains, Oregon (Brown and Thayer, 1963, pp. 415-416). If present, such zeolitization is much less advanced than in the associated non-calcareous laumontitized sandstones, e.g., in the Cretaceous of western Alaska (Hoare et al., 1964, p. C77) and in the Great Valley sequence in California (Dickinson et al., 1969). In the western Olympic Peninsula, Washington (Stewart, 1970, pp. 47-49; 1974), and the adjoining Vancouver Island and Gulf Islands, British Col-

437 umbia (Stewart and Page, 1974), the calcareous graywackes are unzeolitized. Moreover, the calcareous concretions within the zeolitized sandstone beds contain the original detrital oligoclase-andesine and volcanic-lithic detritus, whereas the plagioclase in the enclosing beds is almost exclusively albite, with blocky laumontite replacing the grains and only minor presence of oligoclase-andesine relics. This contrast is likely to be due to the low permeability of these concretions. Similarly, Kubler et al. (1974, p. 463; see also Martini, 1968, pp. 632-633) noted that in the Taveyannaz Graywacke of Haute Savoie, eastern France, the basaltic andesite fragments with fresh plagioclase that has escaped the widespread albitization, are invariably accompanied by a strongly calcitic cement. Calculations for determining the low-grade mineral equilibria for the Ca0-A1,0,-Si0,-C02-H,0 system in T - xH,o-co, space at a constant fluid pressure of 2000 bars have been made by Thompson (197 1b; see Fig. 5-39). Using the position of the “invariant” point A [prehnite-anorthitewollastonite-calcite-quartz-( CO, -H ,0)], obtained by extrapolation into space of experimental T-pco, and T-pHi0 data on two wollastonite-forming reactions, Thompson ( 1971b) calculated 6T/6xCo2values for the following “univariant” reaction emanating from point A: calcite quartz anorthite H,O = prehnite CO,. On extrapolating into TxH,o-co, space of Thompson’s (1970) experimental T-pHzOdata for the (laumontite = anorthite 2 quartz 4 H,O) reaction, Thompson (1971b, p. 151) located a second ‘‘invariant” point B [laumontite-prehnite-anorthite -calcite-quartz-(CO,-H,O)] at Pflu,d = Ptota, = 2000 bars at 317°C and xcoz = 0.0072 (i.e., pco, = 14.4 bars) (see Fig. 5-39, after Thompson, 1971b, fig.3). It should be noted that the association of six phases (five solid and one fluid) in the 5-component system is univariant in P-T-x space, and is only invariant in T-x space when Ptota, = constant. A further univariant curve emanating from the “invariant” point B (laumontite calcite = prehnite quartz CO, 3 H,O), intersects the univariant (laumontite CO, = calcite kaolinite 2 quartz 2 H 2 0 ) curve in a third “invariant” point C [laumontite-kaolinitequartz-calciteprehnite-(CO, -H,O)], tentatively located near 270°C and xco, = 0.0070 (at Pnuid = 2000 bars), both of these univariant curves having extremely steep 6T/6xc,, slopes. This suggests that at Ptotal = P n u i d = 2000 bars, Ca-zeolites are stable with respect to calcite, clay minerals and quartz only if xco, is less than about 0.0070. At lower Pnuid, which is likely to prevail at depths of less than 7 km, and even on assuming that Pnuid 1Plead, the T and xco, of the “invariant” points B and C and, thus, the equilibrium xco, of laumontitekaolinite-calcite-quartz equilibria will be even lower. Thompson ( 197 1b, pp. 153-154) noted that the very high values of xHl0.required for the completion of the formation of Ca-zeolite from a Ca-zeolite-calcite-kaolinite-quartz

+

+

+

+

+

+

+

+

+

+ +

+

+

+

438 I

I

I

I Ptda,

I

I

I

= 2000 bars

:PllUid

0

200

V

t mole lroction X

Fig. 5-39. T - x ~ ~ ,section - ~ , for ~ calculated equilibria in the system CaO-AI,O, -!GO, C0,-H,O, for P,0ta,--PH20+PCOz2000 bars. The insets show the schematic arrangements of univariant cumes around the invariant points B, C. It is probable that C is a metastable invariant point. The indicated portions of ( I , ) , ( 2 ) and (3) may be metastable. (From Thompson, 1971b, fig. 3.)

assemblage can be produced only if the pco, is maintained below the equilibrium pco, for the reaction. This will b e the case only if the CO, evolved is continuously flushed from the system, which necessitates flow

439 rates of the fluids that must exceed the reaction rate of the carbonate decomposition reaction in question, i.e., high permeability. It is obvious that volcanic rocks or derived sedimentary rocks with initially small carbonate contents could recrystallize to stable zeolitecarbonate assemblages without necessarily evolving CO,. If the x C ~in , the associated liquids is low, the pco, will be buffered by the mineral assemblages. AIbee and Zen (1969) summarize the above considerations as follows: “a volcanic graywacke composed of plagioclase, chlorite, quartz, glass, etc., can, with the addition of H,O, form minerals that typify the zeolite facies without thereby evolving CO, , whereas another rocks of the same nonvolatile bulk composition, composed of the assemblage montmorillonite-calcitechlorite-quartz can only react to form zeolitic assemblages if the CO, evolved from the decomposition of calcite is removed from the system. Hence the original mineral composition and the permeability of the rock affects the reaction path.” Incompatibility of dolomite and Ca-Al-silicate minerals: Whereas Ca-zeolite often coexists with calcite, Ca-zeolite-dolomite and prehnite-pumpellyitedolomite assemblages are rare or lacking. The Ca-Al-hydrosilicates are entirely lacking in the upper part of the Salton Sea geothermal field ( 17O0-2OO0C), which is characterized by a high concentration of CO, due to the decomposition of dolomite and ankerite. Epidote and, subsequently, actinolite appear only after the disappearance of dolomite (Muffler and White, 1969). McNamara ( 1965) had ascribed the absence of epidote, prehnite, pumpellyite, and actinolite from the lowest-grade part of the chlorite zone of a Dalradian terrane in Argyllshire, Scotland, to a high value of pcoz, as witnessed by the abundance of calcite and dolomite. Epidote (but not actinolite) appears only in the dolomite-free higher-grade part of the chlorite zone. Coombs et al. (1970, p. 152) suggested that decarbonation reactions of dolomite may well buffer pco, at levels inhibiting formation of Ca-zeolites, prehnite, or pumpellyite. These levels must be rather high to be compatible with the findings of Glassley (1974). With the aid of T-xco, phase diagrams for the system CaO-MgO-A1,0, -SiO, -CO, -H,O, which he constructed using the Schreinemakers method, Glassley has shown that the prehnite-pumpellyite mineral assemblage will exist in equilibrium with a vapor phase in which xco, is equal to up to 0.2 at 1 kbar, and up to 0.15 at 2 kbar. Calcite and Ca-zeolites in hydrothermal systems. In hydrothermal systems the coexistence of calcite with Ca-zeolite is surprising in view of reactions such as: (1) laumontite CO, = calcite kaolinite 2 quartz 2 H,O and (2)

+

+

+

+

440

+

+

+

wairakite CO, = calcite kaolinite quartz. The latter reaction will occur at slightly higher xco, values than the corresponding laumontite reaction (Thompson, 1971b, p. 154). In a H,O-CO, fluid phase bearing the components for Ca-zeolite, calcite, etc., either calcite or Ca-zeolite should precipitate. This corresponds to the contrast between ( 1) the clay-mineral-carbonate trend represented by the Salton Sea geothermal system (Muffler and White, 1969), and (2) the zeolitic trend, with successive appearance of mordenite (and locally some heulandite), laumontite, and wairakite, at Wairakei, New Zealand (Coombs et al., 1959; Steiner, 1968), and several geothermal fields in Japan (Seki, 1966; Seki et al., 1969b), summarized by Zen and Thompson (1974; see also Zen, 1974a). The sequence at the Ohaki-Broadlands geothermal area, New Zealand, with rarity of both wairakite (at 230”-280°C only) and epidote (some above 260”C), the absence of other Ca-zeolites, and abundance of calcite, is intermediate between these two trends (Zen and Thompson, 1974; Zen, 1974a). Browne and Ellis (1970, pp. 122-128) have attributed the contrast between the mineralogy of hydrothermal areas at Ohaki-Broadlands and at Wairakei to the higher initial CO, content of the deep waters, and to the effects on mineral stability of subsequent changes in water composition caused by near-surface boiling-off of steam and CO,. They made use of the log( aCa2+/aH2+) value above which calcite will precipitate at a particular CO, concentration- the calcite solubility line. This calcite “blind” is lowered down the log( aCa2+/aH2+) scale as xco2 is increased (and pH is consequently decreased) at constant temperature. At 260”C, the K-Ca-H water composition at Wairakei is close to the wairakite-epidote-K-feldspar coexistence point, but is well below the calcite solubility line which lies at log(a,az+/aH2+)= 9.1 for the 0.01 pc0, concentration. On the other hand, the deep waters at Ohaki-Broadlands, with a K-Ca-H composition close to the K-mica-K-feldspar-wairakite coexistence point, but having ten times higher initial CO, concentrations ( pco, of about 0.15) (see also Mahon and Finlayson, 1972), are nearly saturated with calcite due to the lowering of the calcite “blind” to a log(a,,,+/aH2+) value of 7.8. Upon boiling-off of steam and CO, from the Ohaki-Broadlands waters at 26OoC, the pH rises, and the K-Ca-H and Na-Ca-H water compositions move to equilibrium first with wairakite and feldspar, then with epidote and feldspar, and finally (after about 3-4% steam loss and at pco2 of about 0.03) with feldspar without Ca-Al-minerals. However, calcite would precipitate throughout, because the calcite “blind” would admittedly lift in proportion to the pH rise, whereas the log(a,,~t/a,2t) value of the solutions rises in proportion to twice the pH change. Wairakite is unlikely to nucleate during formation of calcite (Browne and Ellis, 1970, pp. 124-128).

44 1

The common coexistence of calcite with Ca-zeolites in hydrothermal alteration is ascribed by Thompson ( 1971b, p. 155) to non-equilibrium situations, e.g., zeolite formation after precipitating calcite had lowered pco, to equilibrium value. Gross regularities in burial-metamorphic mineral zonation In summarizing the effect of (1) reduction of the chemical potential of water, and (2) increase in the chemical activity of CO,.in fluid phases on lowest-grade mineral assemblages, Coombs (1971, pp. 322-323) noted that: “...differences in pressure, salinity, CO, content of waters within different beds of the one sedimentary sequence, and osmotic effects can account for the seemingly bewildering overlap of zeolite subfacies”. Understanding the influence of these differences and diagnosing their symptoms and mode of occurrence, can help to establish the sequence of burial-metamorphic assemblages formed under a P / T gradient at otherwise comparable chemical conditions. In establishing the zonation, it is necessary, for instance, to exclude from consideration anomalous mineral assemblages from veins, highly impermeable (“scaled”) lithologies, saline environments, carbonate concretions, etc. More complicated (i.e., harder to detect) are discordant assemblages formed as a result of ionic equilibration with interstitial fluids of modified ionic composition. In some cases, the anomalous nature of such assemblages may be suspected by the presence of incompletely devitrified volcanic glass and un-albitized clastic plagioclase within sequences in which the “normal” burial-metamorphic assemblages are associated with albite and lack volcanic glass. Such criteria, however, may be hard to apply to sequences affected by migration of such interstitial waters. Still, despite the overlaps and complications discussed above, there are striking gross regularities in the vertical or lateral zonation of zeolites and Ca-Al-hydrosilicates, which are diagnostic of successive stages of lowest-grade metamorphism. Most zeolite-facies areas show a distinct laumontite zone intervening between an upper zone characterized by zeolites (such as heulandites, clinoptilolites, mordenite. etc.) on one hand, and a higher-grade facies characterized by pumpellyite without zeolites on the other. Facies series of lowest-grade metamorphism and pressure-dependence of .sonic lowest-grade metamorphic hydrous Ca-AI-silicate minerals Miyashiro (1961) first defined metamorphic facies series in orogenic belts. Such facies series are characterized by different metamorphic assemblages and represent different P / T gradients of metamorphism.

442

In the description of the mineral zonation in the zeolite facies and pumpellyitic facies, the writer noted the existence of differences in the sequences of diagnostic minerals and mineral assemblages in these lowestgrade metamorphic successions. There are prima facie indications that at least some of such differences reflect differences in P / T gradients. Some of these sequences constitute the lower-grade part of metamorphic successions, which show the high-pressure mineral assemblages of the glaucophane-lawsonite-schist facies in their higher-grade zones, indicating low geothermal gradients of metamorphism. In contrast, the low-grade sequences in active and fossil geothermal areas have obviously been formed under high geothermal gradients: in some sedimentary basins in which zeolites are currently forming, the geothermal gradients are known from direct measuremen t s. Moreover, in the preceding section, dealing with the experimentallydetermined mineral-stability fields, it was noted that some minerals or mineral assemblages are restricted to certain ranges of pressure, or that their temperature-stability range is pressure dependent. Unequivocal examples are the restriction of the wairakite stability field to relatively low-pressure gradients, and of the lawsonite stability field to relatively high-pressure gradients. In view of these differences, it is of interest to assess to what extent the differences in the sequences of diagnostic minerals and mineral assemblages in lowest-grade metamorphism can be interpreted in terms of metamorphic facies series. I t has been claimed that one difference between areas with low and high geothermal gradients is the greater diversity of zeolite mineralogy in the latter. In areas with high geothermal gradients (e.g., the Akaishi Mountains, the Tanzawa Mountains, and active geothermal areas) there is a widespread development of zeolites such as mordenite, stilbite, yugawaralite, and wairakite in addition to heulandite, laumontite, and analcime. Under lower geothermal gradients (e.g., in southern New Zealand and the Sanbagawa metamorphic belt) only the last three zeolites are produced. It seems to this writer, however, that this difference may be more apparent than real, because the lowest-grade zones in most intermediate- and high-pressure terranes appear to be lacking, probably due to more advanced erosion. Dlfferences in metamorphic zoning of pumpellyite-hearing sequences Differences in the mineral sequences in pumpellyite-bearing terranes have been noted before. Successive occurrence of prehnite-pumpellyite and pumpellyite-actinolite zones, in which prehnite disappears at or before the actinolite isograd, whereas pumpellyite persists, characterizes ( 1) the southern part of the Wakatipu metamorphic belt of the New Zealand geosyncline

443 (Landis and Coombs, 1967; Landis and Bishop, 1972; Bishop, 1972b), (2) the central Kii Peninsula in the Sanbagawa metamorphic belt (Seki et al., 1971), and (3) the western Appalachian belt in northern Maine (Coombs et al., 1970; Zen, 1974a). Hashimoto (1966, 1972), Seki et al. (1969), and Seki (1969, 1972) have shown that such a succession tends to be restricted to the high-intermediate-pressure and the intermediate-pressure types of metamorphism (e.g., in the above-mentioned terranes). I t was noted earlier that in several terranes, some containing wairakite at lower grades, the pumpellyite-actinolite facies does not intervene between the prehnite-pumpellyite-metagraywacke facies and the greenschist facies proper. On the other hand, it has been pointed out that prehnite occurs rarely, if at all, in the lower grades of glaucophane-bearing terranes, whereas pumpellyite occurs both in the glaucophane-lawsonite-schist facies and as a low-grade “precursor”. In these terranes, the pumpellyite or pumpellyiteactinolite zones are not preceded by a lower-grade prehnite-pumpellyite zone. These field relations constitute a prima-facie indication that pumpellyite, in contrast to prehnite, tends to be stable during high-pressure, lowest-grade metamorphism. Seki (1969; see also 1972) has proposed five lowest-grade metamorphic facies series with different zonal relationships between prehnite, pumpellyite, and actinolite: ( 1) highest-pressure type, (2) high-pressure type, (3) intermediate-pressure type, (4) low-pressure type, and ( 5 ) lowest-pressure type. The type of sequences of critical mineral assemblages in pumpellyitebearing terranes for Seki’s five facies series are schematically depicted in Table 5-VII. This series is amplified by the addition of types (2a), (2b), and (2c), which this author, for various reasons, considers to be intermediate between Seki’s high-pressure (2) and intermediate-pressure type (3) facies series. The zonal sequences for these intermediate to high-pressure types (given in the order of increasing grade of metamorphism) are given below. (2a) Intermediate to high-pressure type I: (i) Prehnite-pumpellyite zone with subordinate actinolite, rarely associated with pumpellyite (no glaucophane or stilpnomelane). (ii) Pumpellyite-actinolite zone with glaucophane (rarely in association with pumpellyite) and stilpnomelane (no prehnite). (iii) Greenschist facies (no pumpellyite). Example: Katsuyama area, Sangun metamorphic belt of western Honshn (Hashimoto, 1968). Because of the prehnite-pumpellyite zone, this terrane is assigned by Seki (1969, 1972) to the intermediate-pressure type (3), from which it differs, however, by the presence of glaucophane in the pumpellyite-actinolite zone.

TABLE 5-VII

Types of progressive successions of low-grade metamorphic prehnite- and pumpellyite-bearing assemblages (facies series) Pumpellyite or prehnite; no zeolites Laumontite or neither laumontite nor pumpellyite (or prehnite) (1) Highest-pressure type: Franciscan metamorphic belt, Diablo Range, central California

(2) High-pressure [vpe: Kanto Mountains, Sanbagawa metamorphic belt central Honshii (Japan) Chichibu metamorphic belt, western Shikoku (southern margin of Sanbagawa belt) (2a) High-intermediate pressure type I ; Katsuyama area, Sangun metamorphic belt, western Honshii (2b) High-intermediate pressure type 11: Tamba area, Honshii

Greenschist facies, no pumpellyite or prehnite

lau

Pum.; laws.; stilp.; Laws.; cross.; stilp.; in some areas act. pum. and minor (no glauc. or pr.) act. in some areas (no pr.. glauc., or jad.)

Laws.; jadeitic px.. glauc.; arag. (no pr., pum., or act.)

chl., stilp. (no pum.)

pum.-chl.; stilp.; very minor pr. (no act. or glauc.)

pum.-epi.-act.; stilp.; epi.-act.-chl.; stilp.; very rare glauc. (rarely with glauc. (greenschist or epi. pum.-epi.-glauc. amphibolite facies) "jad.

pum.-act.; stilp., glauc. (rarely with pum.)

?

pum.-chl.; stilp.; pum.-epi.-act., glauc.. stilp. minor epi.; act. only replacing px.

9

pr.-pum.; minor act. and epi. (no stilp. or glauc.)

9

pum.-chl., extremely pum.-chl., epi., stilp., act., rare pr. and epi. only replacing px. (no glauc.) (no act., stilp. or glauc.)

pum.-act.; epi., glauc. (rarely with purn.), stilp. (no pr.)

Epi.-amphibolite facies

?

epi.-act. (no glauc.)

epi.-act.

(2c) High-iniermediate pressure type I I I : Upper Wakatipu and western Southland districts, southern New Zealand geosyncline

lau. with pr. and pum.

pr.-pum., epi., stilp. (no act. or laws.)

laws.-pum.; epi.; stilp.

pum.-act.; epi.: epi.-act. stilp. some alkaliamph.. not with pum.

( 3 ) Intermediate-pressure type : Dansey Pass area, northern Otago, Southern New Zealand Geosyncline Central Kii Peninsula, Sanbagawa metamorphic belt, southern Honshii Greenlaw-Mooseleuk lake area, northern Maine (western Appalachian belt)

lau., also pr. in highgrade part

pr.-pum.; epi.; local stilp. (no act.)

pum.-act.-chl.; epi.; stilp. ( n o pr. or glauc.)

epi.-act. (greenschist)

?

Iberian pyrite belt ( Hercynian)

(4) Low-pressure t-ppe: Tanzawa Mountains, southern Fossa Magna, central Honshii

?

lau., also wai., some pr. and pum. in high-grade part

Eastern Akaishi Mountains, zeolites (unspec.): also pr. in high-grade part central Honshu

Modificutions; Northeastern Tanzawa lau., also pr., pum. Mountains. southern Fossa and wai. in highMagna. central Honshu grade part Tamuarth Trough. New South Wales

pum.-act.-chl.; local pr.

lau.; some pr.

pr.-pum.: epi. (no act. or stilp.)

pr.-pum.; wai. (no act. or rpi.)

pr.-pum.-act.. wai.: epi.

act.; epi.: minor pr. ( n o puni.)

epi.-act. (greenschist)

greenschist P

pr.-pum.: epi. ( n o act.)

epi.: pr. as relics ( n o pum.)

;ict.;

act.-epi. (greenschist)

P m

TABLE 5-VII (continued) Types of progressive successions of low-grade metamorphic prehnite and pumpellyite-bearing assemblages (facies series) Laumontite or neither Pumpellyite or prehnite; no zeolites laumontite nor pumpellyite (or prehnite) Lyndhurst area, Molong Geanticline, New South Wales

(no zeolites, pr., or pum.)

Wairakei geothermal area ’ North Island, New Zealand lau., wai. in deeper part Onikobe geothermal area, northern Honshu

i?

&

Greenschist facies, no pumpellyite or prehnite

pr. (no pum. or act.)

wai. (no pr. or pum.)

9

?

447 (2b) Intermediate to high-pressure type 11: (i) Pumpellyite zone (no actinolite; extremely rare prehnite). (ii) Transitional zone with pumpellyite and stilpnomelane (actinolite only at the expense of relic clinopyroxene; not associated with pumpellyite). (iii) Epidote-actinolite zone (greenschist facies; no pumpellyite). Example: Tamba area, Kyoto Prefecture, Honshii. This terrane, though considered by Hashimoto ( 1972) to represent a lower pressure terrane than the Chichibu belt (absence of glaucophane), differs from the intermediatepressure type (3) by the absence of a real pumpellyite-actinolite zone and the extreme rarity of prehnite. (2c) Intermediate to high-pressure type 111: (i) Prehnite-pumpellyite zone (without actinolite). (ii) Lawsonite-albite zone with pumpellyite (without prehnite or actinolite). (iii) Pumpellyite- actinolite zone with local alkali-amphibole (without lawsonite). (iv) Greenschist facies (no pumpellyite). Examples: Upper Wakatipu and western Southland districts, southern New Zealand Geosyncline (Landis, 1974; Kawachi, 1974, 1975). In addition, some terranes show modifications of the low-pressure type (4). The succession of the northeastern part of the Tanzawa Mountains (Shimazu et al., 1971) is exceptional in the persistence of prehnite, and to a lesser extent of pumpellyite, beyond the actinolite isograd, although the occurrence of wairakite in the prehnite-pumpellyite zone attests to the relatively low pressures. The persistence of prehnite (without pumpellyite) into the low-grade part of the actinolite zone in the Tamworth Trough and the Molong Geanticline of New South Wales (Packham and Crook, 1960; Smith, 1969) similarly indicate low-pressure environments. The estimated temperature-pressure gradients of Seki’s lowest-grade metamorphic facies series are shown in Fig. 5-40 after Coombs (1971). Nature of the actinolite isograd. In accounting for these different zonal sequences, Hashimoto ( 1972) has stressed the three different possible relationships of the actinolite isograd to the decomposition of pumpellyite and prehnite: ( 1 ) In the pumpellyite-actinolite zone of glaucophane-bearing terranes such as the Chichibu metamorphic belt (and apparently also in the KantB Mountains and the Katsuyama area), the actinolite commonly forms in association with pumpellyite or epidote (i.e., before decomposition of pumpellyite) and not by pumpellyite-decomposition reactions. The formation of actinolite in the lower-grade pumpellyite zone as a replacement of

448

4r

LAWSONITE - A L B I T E CHLORITE

h

IJl

L

0

n

2-

Y

I

100

200

300

400

Temp. (OC1

Fig. 5-40. Possible P - T fields for low-grade mineral facies and subfacies, calibrated for the laumontite-wairakite (Liou) and analcime-t quartz-albite (Campbell and Fyfe, 1965) equilibria at p H , O= p,,,,,. Arrows represent facies series slightly modified from Seki (1969), as follows: ( 1 - 2 ) highest-pressure to high-pressure type: ( 3 ) intermediate-pressure type: ( 4 ) low-pressure type; ( 5 ) lowest-pressure type. (After Coombs. 1971, fig. 3.)

relic clinopyroxene exclusively, does not constitute a stable progressive equilibrium reaction and. thus, does not define a true actinolite isograd. (2) In the lower-pressure Tamba area, the actinolite isograd appears to be defined by the decomposition of pumpellyite into actinolite and epidote at the beginning of the epidote-actinolite (greenschist facies) zone. The actinolite in the transitional zone from the lower-grade pumpellyite zone to this epidote-actinolite zone occurs as a replacement of clinopyroxene only, and is not associated with either pumpellyite or epidote. (3)In the Greenlaw-Mooseleuk Lake area (Coombs et al., 1970), the actinolite isograd is defined by the decomposition of prehnire to the actinolite-epidote assemblage, whereas pumpellyite persists and stably coexists with actinolite. This relationship presumably also exists in northern Otago, New Zealand, and in the central Kii Peninsula, southern Honshu. Inasmuch as none of these successions contain high-pressure assemblages with glaucophane. they are considered by Hashimoto ( 1972) to represent lower-pressure conditions than the Chichibu belt and the Tamba district.

449

At low pressures, pumpellyite disappears with progressive metamorphism upon or before the appearance of actinolite. The actinolite isograd is defined by the decomposition of both prehnite and pumpellyite in the Tanzawa Mountains (metamorphism under the influence of a large quartz-diorite intrusion), (2) in the Lyndhurst district of New South Wales (Smith. 1969). and probably (3) in the Tamworth Trough of New South Wales (Packham and Crook, 1960), and (4)in the eastern Akaishi Mountains (Matsuda and Kuriyagawa, 1965), although actinolite may already form at a lower grade at the expense of clinopyroxene in these terranes as well. High pressure thus appears to favor the formation of pumpellyite in basic rocks. According to Hashimoto ( 1972), the temperature of pumpellyite deconiposition increases with increasing pressure. At high pressures, pumpellyi te occurs at a higher temperature than that necessary for the following actinoli te-producing reaction : chlore (clinochlore) -+

actinolite

+ quartz + calcite

+ epidote (or pumpellyite) + water + C 0 2 .

This is in agreement with the experimental results. quoted below, that delimit the actinolite-chlorite-pumpellyite-quartz (without prehni te) paragenesis to geothermal gradients lower than about 35"C/km. Prehnite, on the other hand, is indicative of pressures lower than those that prevailed in both the Chichibu belt and the Tamba district. Physical conditions and pressure dependence of diagnostic equilibriu Hiitliin the prehnite-pumpellyite and pumpelIvite-uctinolite facies: experimentul ei$ dence In an earlier section, the writer discussed the equilibria between launiontite, prehnite, and pumpellyite in the deeper parts of the laumontite zone of the zeolite facies. The disappearance of laumontite marks the transition to the higher-grade metamorphic facies characterized by the presence of pumpellyite and the absence of zeolites (except wairakite) in the low-grade part. The equilibrium relationships which exist between pumpellyite, prehnite, and actinolite within the pumpellyite-bearing zones and at their high-grade limit are discussed here with particular reference to the pressure dependence of the mineral assemblages. The following two reactions which are respectively characteristic of ( 1 ) the high-grade boundary of the prehnite-zone within the prehnite-pumpellyite metagraywacke facies, and (2) the low-grade boundary of the greenschist facies, have been experimentally studied in the Fe-free CaO-MgO- Al 20,-

450 SiO, -H,O system by Hinrichsen and Schurmann ( 1969, 1972):

+ chlorite (amesite molecule) + H,O = pumpellyite + quartz (Seki, 1961),

(1) prehnite

and (2) pumpellyite = zoisite

+ grossular + chlorite (arnesite molecule)

+ quartz + H,O. In the 54 -8 kbar pH,, pressure range, the first equilibrium lies between 245" and 265°C (negative dpH2,/dTslope) and the second one lies between 310" and 345°C (positive dpH2,/dT slope). The temperature stability range of pumpellyite thus widens with increasing pressure. The characteristic pumpellyite-chlorite-actinolite assemblage of the pumpellyite-actinolite-schist facies is related to the prehnite-bearing assemblages of the lower-grade facies by the following reaction: prehnite

+ chlorite k H,O = pumpellyite + actinolite 5 quartz.

This reaction has the advantage over the earlier pumpellyite-forming reaction in that it requires a less aluminium-rich chlorite composition, which may vary from clinochlore (e.g., Seki et al., 1969a, p. 67) to talc-chlorite (e.g., Hashimoto, 1966; Seki, 1969, p. 259). The appearance of minor amounts of water and quartz on either side of the equation depends on the precise composition of the chlorite. In any case, the AV of the above reaction should be markedly negative (Coombs et al., 1970, p. 149; Nitsch, 1971, p. 242, footnote). Nitsch's (1971) provisional (as yet un-reversed) experimental results on this reaction in the system CaO-MgO-Al,O, -SO, -H,O show close agreement with those of Hinrichsen and Schurmann (1969, 1972). They indicate that the prehnitepumpellyite-chlorite-quartz assemblage is stable up to 260" 20°C at 7 kbar, and up to higher temperatures at lower pressures (negative d P/dT slope). The following pumpellyite-breakdown reaction (with variable chlorite composition, cf. Banno, 1964; Seki, 1969; Seki et al., 1969a, p. 18) was shown in reversed experiments by Nitsch (1971) to be in equilibrium at a temperature of approximately 370°C at a pressure of 7kbar, and at a temperature of approximately 34Oo-350"C at 4 kbar:

*

pumpellyite

+ chlorite + quartz = clinozoisite + actinolite + H,O.

I t should be noted, however, that the substitution of Fe and Mn for Mg in pumpellyite, the effect of which has not been evaluated by Hinrichsen and Schurmann (1969) and by Nitsch (1971), could considerably lower both its temperature and pressure stability (Thompson, 1970, p. 274; Hinrichsen and

45 1 Schurmann, 1972). Moreover, the pumpellyite-consuming reaction should proceed at lower temperatures at high oxidation states, because ferrous purnpellyite is related to epidote by a simple oxidation-dehydration relationship: Ca,Al,Fe”[Si,O,,/O/(OH),] Fe-pumpellyite = Ca,Al,Fe”’[Si,O,,/O,/

.2H,O (OH),]

+4 0

+ 2:

H,O.

epidote (pistacite) Purnpellyite may thus form already at lower grades of metamorphism than are represented by the above pumpellyite-forming reactions. The temperature limits found by Hinrichsen and Schurmann (1969) and by Nitsch (1971) nevertheless show remarkable agreement with the equilibrium conditions of 2OO0-325”C and at least 6-7 kbar pressure found by Taylor and Coleman ( 1968, p. 1951) by ’xO/160determinations on pumpellyite-bearing metabasalts (low-temperature blueschists), and a temperature bf 350°C found by Devereux (1968; see also Bishop, 1972b, p. 3193) on calcite from a segregation lens in a metavolcanic rock very close to the pumpellyite-clinozoisite isograd. The temperatures of 50Oo-55O0C, found earlier by Landis and Rogers (1968) for the irreversible decomposition of pumpellyite to anorthite clinopyroxene in the presence of quartz in the pressure range of 500-3000 bars, are improbably high in view of the natural occurrence of pumpellyite and, probably, represent a metastable reaction. The steep slopes in p H I 0 - T space (i.e., the relatively small dependence on pressure variations) of the pumpellyite-breakdown reactions studied by Hinrichsen and Schurmann (1969, 1972) and Nitsch (1971) make the disappearance of this mineral an approximate geothermometer, and justify its use by Winkler (1970) as an isograd diagnostic for the high-grade limit of “very-low-grade’’ metamorphism. Accepting this criterion, the writer includes all regionally pumpellyite-bearing metasediments or metavolcanics in the realm of incipient metamorphism, which as shown later is approximately equivalent to Kubler’s anchimetamorphic zone in clay-rich sedimentary rocks.

+

Pressure-temperature gradient and succession of metamorphic ussemhlugrs The stability of the paragenesis actinolite-chlorite-purnpellyite-quartz (without prehnite) is thus delimited by two univariant equilibrium curves intersecting in an invariant point. The latter was located at 345 20°C and 2.5 2 1 kbar (see Fig. 5-41) by Nitsch (1971), using the extrapolated lowpressure extensions of the univariant curves. Consequently, the stability of

45 2

I

Km

30

20

10

Fig. 5-41. P-T diagram showing the equilibrium curves for the reactions: prehnite+chlorite=pumpellyite+ actinolitei quartz ( - 7 ) pumpellyite+chlorite+quartz=epidote+actinolite ( 3 ) pumpellyite+quartz= prehnite+epidote+ actinolite ( 4 ) prehnite + chlorite=pumpellyite+epidote+ actinolite ( 5 ) prehnite-t chlorite+quartz=epidote+ actinolite ( 6 ) pumpellyite+ quartz=prehnite+epidote+ actinolite in the system CaO-MgO-Al,O,Drawn curves: experimentally reversed; broken curve: experimentally approxiSiO, -H mated; stippled curve: schematic position. The quartz-bearing assemblages expected to occur commonly in each of the sectors I , 11, I V , and V I , are shown framed. A=actinolite; C =chlorite; E=epidote; P=prehnite; Pu=pumpellyite; Q=quartz. The bands represent the hivariant range of the curves when FeO is incorporated in the system, and J.P. indicates the range of the invariant point for different Mg/Fe ratios. The equilibrium curves for the reactions calcite=aragonite and jadeite+quartz==albite are indicated (Ab=albite; Ar= aragonite; Cc=calcite; Jd=jadeite), as well as four geothermal gradients. (After Nitsch, 1971, fig. 6.) (/ )

453

the paragenesis, which is diagnostic for the pumpellyite-actinolite-schist facies, is not only restricted to higher pressures (geothermal gradients lower than 35"C/km), but its temperature range is strongly pressure dependent. I t widens from only approximately 300" to 355°C at a geothermal gradient of 20°C/km, but from approximately 255" to more than 380" at 10"C/km. I t should be kept in mind, however, that in natural systems the position of the univariant equilibrium curves, and thus also that of the invariant point, is likely to be affected by factors such as the Fe content of the minerals and the oxidation state, and thus to be slightly different (Thompson, 1970, p.274; Bishop, 1972b, p. 3 191; see pp. 429-430 in this chapter). The various sequences of pumpellyite-bearing assemblages or burial-metamorphic facies series (Seki, 1969, 1972) may now be interpreted using these experimentally determined equilibrium relationships. The successive occurrence of prehnite-pumpellyite zone and a prehni tefree pumpellyite-actinolite zone, characterizing the intermediate type (3) sequences (Table 5-VII), is to be expected at geothermal gradients of less than 35"C/km. With the increasing pressures of lower geothermal gradients, the stability field of the pumpellyite-actinolite-chlorite assemblage progressively widens to lower temperatures at the expense of that of prehnite-pumpellyite-chlorite assemblage. According to a recent evaluation of experimental data by Seki (1973c), the invariant point defined by the intersection of the univariant lawsonitelaumontite-quartz equilibrium (the pressure of which is only very slightly temperature dependent) and the high-temperature stability limits of lawsonite and laumontite lies at a higher pressure (3.5 kbar) and at a lower temperature than the invariant point under discussion (on a geothermal gradient of approx. 23"C/km). This explains why lawsonite only appears in terranes with sequences showing a wide stability range of the pumpellyite-actinolite assemblage, and little or no prehnite-bearing zone. With increasing pressures, the prehnite-bearing assemblages may not appear at all before the occurrence of pumpellyite-actinolite-chlorite assemblage. At a geothermal gradient of 10"C/km, the appearance of the latter assemblage at 260°C could already be associated with Na-rich amphibole. This is the case in the sequences from the Sanbagawa and Chichibu metamorphic belts representing the high-pressure type (2). At the lower pressures of the invariant point itself, i.e., at the relatively high geothermal gradient of 35"C/km, the conditions of the stability of the pumpellyite-actinolite-chlorite assemblage are not attained at all (see Fig. 5-41). At about 350"C, the greenschist-facies assemblage actinolite-chlorite epidote-quartz will immediately succeed prehnite-pumpellyite-chlori te assemblage, i.e., prehnite and pumpellyite both disappear close to the ap-

454

pearance of actinolite (Coombs et al., 1970, p. 149). Such a sequence is found in the low-pressure type (4) areas in southeastern Honshu, Japan. At pressures below that of the invariant point, i.e., at still higher geothermal gradients, the prehnite-chlorite-quartz assemblage is actually stable to a somewhat higher temperature than pumpellyite-quartz assemblage. The former assemblage persists even through a narrow zone in which actinolite is already forming according to the following reaction (Nitsch, 197 1, p. 256 and fig. 6): prehnite

+ chlorite + quartz = epidote + actinolite + H 2 0 .

This may account for the persistence of prehnite to a somewhat higher degree of metamorphism than pumpellyite in the Tamworth Trough and the Molong Geanticline of New South Wales (Packham and Crook, 1960; Smith, 1969). These sequences, therefore, are considered by Hashimoto ( 1972) to represent very-low-pressure conditions. The experimental evidence on the stability of pumpellyite-, prehnite-. and wairakite-bearing mineral assemblages thus confirms that the facies series of lowest-grade metamorphism reflect differences in the P and T gradients during burial and recrystallization. Tentative P and T gradients for the five main low-grade metamorphic facies series recognized are summarized in Fig. 5-40 (after Seki, 1969 and Coombs, 1971). As more detailed mineral-assemblage stability data become available and as the variety of mineral assemblages that may appear under isometamorphic conditions become better understood, i t will no doubt become possible to resolve apparent inconsistencies and recognize facies series of lowest-grade metamorphism in other areas. It will also be necessary to relate these facies series to the facies series recognized in the higher-grade, regionalmetamorphic terranes. Rank of coal associated with burial-metamorphic mineral fucies The extent of coal-rank ranges associated with the various burialmetamorphic and lowest-grade metamorphic mineral facies in volcanic and lithic-feldspathic clastic sediments has been investigated in a number of papers by this writer (Kisch, 1966b; 1969; 1974; 1980a, b; in preparation), to which the reader is referred for details and for references. For details on coal rank and its parameters see Chapter 5 of Volume1 of Diagenesrs qf Sediments and Sedimentary Rocks. Only the general relationships are outlined below: ( 1) Heulandite-clinoptilolite and analcime have been found in several localities associated with coals ranging in rank from brown coal to approximately the limit between high-volatile A and medium-volatile bituminous coals (31% volatile matter, d.a.f.), e.g., in the Werrie Basin, New

455 South Wales. In the Triassic section of the Taringatura district, heulandite persists into the laumontite zone and thus is associated with high-volatile A and medium-volatile bituminous ranks in the lower units (Kisch, in preparation). (2) Laumontite is associated with coals of high-volatile A and/or medium-volatile bituminous ranks, e.g.. in: (a) the Lower-Cretaceous coal measures of the Lena Coal Basin. northern Yakutia, Siberia (Zaporozhtseva et al., 1961; 1963; cf. Kisch, 1966b. 1969; Kaplan, 1974); (b) the Permian Coal measures of the lower Hunter Valley, New South Wales (Kisch, 1966b, 1969); (c) the Lower Jurassic of Cape Paterson, Victoria (Gill, 1956; Kisch, 1974, table 111); (d) the Lower Tertiary of the Diablerets area and the lower Kien Valley in the Helvetic zone of the Swiss Alps (Kisch, 1980b); and (e) the lower units of the Triassic section of the Taringatura district, Southland, N.Z. (Kisch, in preparation). In some areas, the associated coals seem to be of somewhat higher rank. In some sections in the Permian and Upper Carboniferous of New South Wales, e.g., in the Illawarra area, southern Sydney Basin. and the Werrie Basin, Tamworth Trough, the coal ranks associated with laumontite-bearing sequences as obtained by downward extrapolation (from the ranks of coals overlying the laumontite-bearing sequences) seem to range from mediumvolatile bituminous at the top to at least low-volatile bituminous at the base of the laumontite-bearing sequence in the Illawarra area (Kisch, 1966b, 1969). In the Helvetic zone of the Swiss and French Alps, coals of low-volatile bituminous rank (14-22% volatile matter, d.a.f.) have been found to be closely associated with the higher grade part of the laumontite zone in the Taveyanne Graywacke in the Massif de Plate, Haute Savoie (Martini. 1972, p. 261; Kisch, 1974, table III), and at Kandergrund in the Kander Valley (Kisch, 1980b). The low-volatile bituminous coals at Kandergrund may admittedly be from a higher tectonic unit (Wildhorn nappe) than the laumontite-bearing Taveyanne Graywacke, but the zeolite-facies metamorphism is likely to post-date the emplacement of this higher nappe (e.g., Martini, 1972; Frey et al., 1973). In some younger sequences, on the other hand, for instance the EoceneMiocene of the Tejon area in the San Joaquin Valley (Castaiio and Sparks, 1974) and Lower Cretaceous of the Great Valley sequence of Cache Creek (Dickinson et al., 1969; Bostick, 1974; Castaiio and Sparks, 1974), both in California, but also in the higher units of the Lower and Middle Triassic North Range Group of Southland, N.Z., the top of the laumontite-bearing sequences is associated with lower, high-volatile B and, locally, possibly

45 6 high-volatile C bituminous ranks (0.4-0.62 Rmo,,for most of the Californian localities). (3) Coal- rank data from the prehnite- pumpellyite facies -without laumontite-are available only from the Lower Tertiary of the Helvetic zone of the Swiss Alps. f i s c h (1980b) has found that prehnite- and pumpellyitebearing, laumontite-free Taveyanne Graywacke of (i) La Tieche in the northern Rhone Valley, (ii) Reichenbachfall (central Switzerland), and (iii) two localities in the Schachen and Urnerboden Valleys in eastern Switzerland, are associated with vitrinite reflectances ranging between 3.3 and 4.0% R,,, (i.e., medium-rank anthracite). (4) Coal rank data from the lawsonite-albite facies and lawsonite-bearing pumpellyite facies (the latter as a lower-grade “predecessor” to the glaucophane-lawsonite-schist facies) are restricted to data from the Brianqonnais “zone houillere” of the French Alps, the Franciscan of central California, and some extrapolated data from the Sanbagawa metamorphic belt of HonshU. Proximate coal analyses of Carboniferous coals from the southern part of the “Zone houillere brianqonnaise” (Feys, 1963, pp. 55-73, and Annex II), in the Alpine lawsonite- and pumpellyite-bearing zone of Saliot (1973), show 10-4% volatile matter (d.a.f.), i.e., high-rank semi-anthracite to medium-rank anthracite (coal analyses with more than 25% ash not considered). Reflectances of 5-5.5% R,, oil (high-rank anthracite) have been measured on coals from the southern termination of the “zone houillere” south of Brianqon (Robert, 1971, p. 120), but 3.72% R,, oil (medium-rank anthracite, 4.3% V.M.) was obtained on Dogger coal from the same area by Chateauneuf et al. (1973). These anthracite ranks are quite similar to those found associated with prehnite-pumpellyite facies Taveyanne Graywacke in central and eastern Switzerland (Kisch, 1980b). In the Sanbagawa belt in the Chichibu district of central Honshil, high-rank anthracite with 1.46%H occurs in zone V, the “pumpellyite-epidoteactinolite stage” of Seki ( 1958), which succeeds the glaucophane-lawsoniteschist facies (with lawsonite and jadeite) of zone IV (see Kisch, 1974, p. 103). I t seems likely that the ranks in the glaucophane-lawsonite-schist terrane of the area are anthracitic. Bostick (197 1, 1974) has given vitrinite reflectance data on coaly particles from shales interbedded with lawsonite- and locally aragonite-bearing graywackes of the Franciscan assemblage of the Mt. Hamilton Range (part of the

’

’

Anthracite is here delimited from meta-anthracite using the 2% volatile-matter limit of the ASTM classification (corresponding to about 6% R,,, o,, vitrinite reflectance and about 1.5% H). and by the appearance of three-dimensional ( h k l ) X-ray diffraction peaks in metaanthracite (Mentser et al., 1962, 1963).

457 Diablo Range), central California. The “upper” (flank) unit-which locally contains jadeitic pyroxene at the southwestern margin of the Diablo Range -shows reflectances of 1.1 - 1.3% R,,, (high-volatile A to medium-volatile bituminous ranks), whereas the “lower” (core) unit shows 1.75- 1.9% R,,, (low-volatile bituminous rank). ( 5 )Assuming that coal rank is essentially unaffected by pressure (see Chapter 5 of VolumeI), the similarity of the ranges of rank associated with the lawsonite graywackes in the Diablo Range and with the laumontite zone of the zeolite facies in several areas indicates that these mineral facies were formed at similar burial temperatures (but the former under much higher pressures). The much higher, anthracitic, ranks found in the lawsonite- and pumpellyite-bearing terrane in the Pennine realm of the French Alps, and most likely also associated with the glaucophane-lawsonite-schist facies in the Sanbagawa metamorphic belt are in agreement with the postulated lower temperature (and higher pressure) during Franciscan than during Sanbagawa metamorphism (e.g., Ernst and Seki, 1967; Ernst et a]., 1970); this difference is evident, among other things, from the absence or rarity, of lawsonite and jadeitic pyroxene in metaclastics, and the absence of aragonite in Sanbagawa metamorphism. In the French Alps, the thermal gradients during the earlyAlpine phase of incipient metamorphism are likely to have been more similar to those in the Sanbagawa belt, in view of the absence of lawsonite north of the Vanoise, the rarity of jadeitic pyroxene, and the absence of aragonite (Bocquet, 1971; cf. Kienast and Velde, 1970; Ernst, 1973). Clay minerals and authigenic layer silicates associated with burial-metamorphic mineral facies In order to obtain an integrated picture of the mineral zonation in burial metamorphism, it is necessary to correlate the mineral facies recognized in volcanic clastic rocks and tuffs with the stages of burial diagenesis and initial metamorphism in clay-rich clastic rocks. One way of arriving at such a correlation is to relate the clay mineralogy of rock series showing diagnostic mineral assemblages of the zeolite- and the pumpellyite-bearing facies with the diagnostic clay-mineral alteration processes described earlier. For this purpose, it is not necessary to be restricted to the clay minerals in the-usually volcaniclastic-zeolite-bearing rocks; one should preferably also consider the clay minerals in other “iso-metamorphic” clay-rich clastic rocks from the same sequence. Unfortunately, relatively few studies of zeolite facies and prehnitepumpellyite-bearing assemblages include detailed data on the clay mineralogy of the rocks themselves or of associated rocks. Nevertheless, the scanty data available provide some indication as to the approximate position of these mineral facies with respect to stages of late diagenesis and initial metamorphism in clay-rich sediments and rocks.

45 8 Occurrence of expandable clays and celadonite in burial-metamorphicfacies Montmorillonite. Montmorillonite and illite-montmorillonite mixed-layers are major constituents of the clay-minerals of zeolite-facies sequences. Montmorillonite appears to be ubiquitous in association with the alkalirich zeolites characterizing the uppermost zones of the zeolite facies. I t is associated, often together with celadonite, with analcime, clinoptilolite and/or mordenite in the Cenozoic altered vitric tuffs and lacustrine shales of the Green River Formation of Wyoming (Hay, 1966; Roehler, 1972; Surdam and Parker, 1972) and elsewhere in the western United States (e.g., Sheppard and Gude, 1969; Robinson, 1966; Robinson et al., 1968; Moiola, 1970). Montmorillonite is associated both with clinoptilolite (or heulandite) and with cristobalite or tridymite in tuffs and bentonites of Jurassic to Eocene age in various parts of the world (Iijima, 1968; Reynolds and Anderson, 1967; Brown et al., 1969; Hallam and Snellwood, 1970; Reynolds, 1970). It is also associated with analcime without clinoptilolite or heulandite in lacustrine clastic rocks and tuffs ( e g , Sabine, 1963; High and Picard, 1965; locally also in the Green River Formation). Montmorillonite is also a common clay mineral in the low-grade part of the zoned zeolite-facies sequences mentioned earlier. In several of such sequences, the smectite appears to become less common downwards, but to persist into the laumontite zone, e.g., in the “Green Tuff” of the Shinjo Basin, northern Honshii (Utada, 1965; Iijima and Utada, 1966), and in the burial-metamorphic Miocene sandstones of the central California coast ranges (Murata and Whiteley, 1973). However, here it tends to be accompanied or, locally, even replaced by illite-montmorillonite mixed-layers (e.g., Iijima and Utada, 1971; Shimazu et al., 1971). Illite-montmorillonite mixed-layers. Montmorillonite as a discrete mineral commonly disappears within the laumontite zone by progressive mixedlayering, commonly in the upper part of the zone ( e g , Kossovskaya and Shutov, 1963, p. 1162), well before the appearance of pumpellyite, e.g., in the vulcanites of eastern Angola (Portugal Ferreira et al., 1973) and in the Tanzawa Mountains (Shimazu et al., 1971). In these and many other laumontite zone terranes, di-octahedral montmorillonite is largely restricted to occurrence as a component of expandable mixed-layers (e.g., Kaplan, 1974), which have been noted to persist even in the lower, prehnite- and pumpellyite-bearing part of the laumontite zone, for instance at the base of the Triassic section in the classical Taringatura district of Southland, N.Z. (Kisch, in preparation). In contrast, only traces or absence of expandable mixed-layers have been reported from clayey horizons intercalated with the laumontite-bearing

45 9 Taveyanne Graywacke in the Massif de Plate, Haute-Savoie, French Alps, which is immediately underlain by a coal of low-volatile bituminous rank (Martini, 1972, p. 261 and fig. 2; see also fisch, 1974, p. 99). No discrete mixed-layer diffraction peaks were found in pelitic rocks associated with the higher-grade, prehnite-pumpellyite-bearing part of the laumontite zone in the Taveyanne Graywacke of the Diablerets area, the middle Kander Valley, and lower Kien Valley in the Helvetic zone of the Swiss Alps (Kisch, 1980b), and in the Torlesse terrane of the Benmore dam area in northern Otago, N.Z. (Kisch, in preparation): the expandable layers are apparent only as unresolved low-angle ( = high d-spacing) tails to the I O A illite diffraction peak, the extent of which varies upon cation- and ethylene-glycol-saturation (“open illite” I of Lucas, 1962). The half-height widths of these 10 A diffraction peaks (“illite crystallinity”-see p. 344ff.) indicate that the anchimetamorphic zone of Kubler (1967a) has not yet been reached. The illite accompanying the prehnite-pumpellyite-metagraywacke facies in the Taveyanne Graywacke of the Helvetic zone of eastern Switzerland (Kisch, 1980b), and in the Caples terrane of southern Otago, N.Z. (Kisch, in preparation) carry only very subordinate proportions of smectitic interlayers as apparent from the minor changes in the low-angle “tails” and the 10A peak widths upon cation- and ethylene-glycol-saturation; the 10 A half peak widths are characteristic for the anchimetamorphism of Kubler ( 1967a). Expandable chlorite- or uermiculite-rich mixed-layers. In other laumontite zones, the mixed-layering of montmorillonite as well as vermiculite occurs largely with chlorite. From the Miocene pyroclastics of the Tanzawa Mountains of central Honshfi and the Shinjo Basin of northern HonshU, for instance, Seki et al. (1969a, pp. 50-51) and Shimazu et al. (1971), and Utada (1965, pp. 149-196), respectively, found a marked decrease in the amount of montmorillonite layers in the montmorillonite-vermiculite and vermiculitechlorite mixed-layers in the laumontite-bearing zones. Montmorillonite layers almost completely disappear in the highest-grade part of this zone (see Fig. 5-42). The occurrence of mixed-layer phases evidence the gradual alteration of montmorillonite through chlorite-montmorillonite mixed-layers into chlorite, which has also been studied in laumontite-zone sandstones of the Irkutsk Coal Basin and of Kamchatka, Siberia, by Koporulin (1961, 1968, 1972). Locally, such mixed layering in the laumontite zone takes the form of regular, corrensite-type chlorite-swelling chlorite or chlorite-montmorillonite mixed-layers (Kubler, 1969; Iijima and Utada, 1971; Kubler et al., 1974). Corrensite has also been reported in chlorite-prehnite and chloritelaumontite alteration zones in a gabbro as a deuteric alteration product and

460

~

( 1 ) Neogcne of the southern Tonzowo mountains (Seki et o

"p

Zlinoptilolite jtilbite Heulonditc vlor denite -0urnontite rhomsonitc Nairakit e hnolcime klontmorillonite/vermiculite Celodonite iermicullte I chlorite Dumpellyite 'reh nitc E pi dote 4ctinolite -lorn blende Chlorite Sericite Biotite

1969 c

P

-

-:++T

( 2 ) Neogene of the northeasteh Tan: wa mountoins(Shim0zu

et

.,1971)

I 1 vlordenite Heulondite 5 t ilbit e T homsonite 4nolcime ,oumontite Naira k it e Montmorillonite Soponite Chlorlte / montmorillonite Celodonite Pnhnite Pumpellyite Epidote Actinolite Chlorite Sericite

I

(3)Eastern Akaishi mountoins'(M0ta

Prehnite

Fig. 5-42. Lowest-grade metamorphic zoning of zeolites and other hydrous Ca- A1 silicates, and of phyllosilicates in three low-pressure type areas in central Honshii (broken lines: minerals are not common). References for correlations of zones: ( I ) disappearance of laumontite; (2) disappearance of common pumpellyite and prehnite.

46 1 hydrothermal fracture-filling material (Furbish, 1973, confirming the similarity in conditions of corrensite formation. Corrensite is the main mineral in the fine fraction of the laumontite-bearing Taveyanne Graywackes of the Helvetic zone in western Switzerland and France, but disappears in the higher-grade prehnite-pumpellyite rocks, in which chlorite takes its place (Kubler et al., 1974, p.465; Aprahamian, 1974, p. 10; Kisch, 1980b). Fe-saponite. The trioctahedral smectite, Fe-saponite, appears to persist to a higher grade of burial than di-octahedral montmorillonite. In both the northeastern Tanzawa Mountans and in the eastern Akaishi Mountains (central Honshii), it persists until the high-grade boundary of the zeolite facies, but is absent from the laumontite-free prehnite-pumpellyite zone (Shimazu et al., 1971; Matsuda and Kuriyagawa, 1965; see Fig. 5-42). Read ( 1968) has found that the Fe-saponite previously reported from the Lake Wakatipu area, South Island, is in fact stilpnomelane. In New Zealand, as in Japan, Fe-saponite may be expected to be absent from rocks of higher grade than the zeolite facies. In summary, therefore, it appears that smectite (with the exception of Fe-saponite) rarely appears in the laumontite zone other than as a component of mixed-layers with illite or chlorite. Such smectite-rich mixed-layers may locally persist in the higher-grade, prehnite- and pumpellyite-bearing part of the laumontite zone, but rarely into the prehnite-pumpellyite facies proper. Celadonite. The Mg-Fe-rich di-octahedral 1 M mica celadonite is another common three-layer phyllosilicate in the zeolite facies. In the Taringatura and Hokonui Hills, Southland (N.Z.), i t is much less abundant in laumontitized than in heulanditized tuffs (Coombs, 1954, p. 80; see also Boles and Coombs, 1975). Dickinson (1962a) has reported chloritization of celadonite concurrently with the laumontitization of heulandite. Nevertheless, celadonite is comparatively common in laumontite-bearing assemblages, not only on South Island, New Zealand, but also in the Sustut Group of British Columbia (Read and Eisbacher, 1974, p.535-536) and in the Cretaceous of coastal Angola (Portugal Ferreira et al., 1973). In the latter area, in central Honshii, and elsewhere, celadonite is associated with Fesaponite (e.g., Ross, 1958; see also Wise and Eugster, 1964, p. 1047). In their review of celadonite occurrences in low-grade altered volcanic rocks, Wise and Eugster (1964, p. 1047) listed the metavolcanics of eastern Oregon as the only instance of association of celadonite with pumpellyite or prehnite. In Oregon, moreover, celadonite is also common in green bands in the prehnite-, pumpellyite- and laumontite-bearing tuffs and lavas of the

462 lower units of the earlier-mentioned Triassic sequence of the Aldrich Mountains (Brown, 1961; Brown and Thayer, 1963). Still, celadonite, as Fesaponite, seems to be largely restricted to the zeolite facies. It is absent from the zeolite-free prehnite-pumpellyite zones of New Zealand and central Japan, and sparse in those of northern Maine (Coombs et al., 1970, pp. 146- 147). Occurrence of kaolinite in burial-metamorphic facies Kaolinite is not a predominant clay mineral in the altered intermediate volcanic clastic rocks and tuffs which are the main locus of the zeolite facies. It is likely to be less common than montmorillonite in zeolite-facies mineral assemblages. Nevertheless, kaolinite has been reported not only to occur but even to form in laumontite zones, e.g., as an alteration product of feldspars in New South Wales (Whetten, 1965; Raam, 1968). It also occurs throughout the zeolitic Sustut Group of central British Columbia, being well-ordered towards the base (Read and Eisbacher, 1974, p. 531). Kaplan (1965) reported that in the Triassic sandstones of the southern Maritime Provinces, eastern Siberia, laumontite persists to somewhat greater depth than kaolinite. Laumontitic and kaolinitic alteration. In the Blairmore Group of the Alberta Foothills, laumontite and kaolinite occur respectively in sandstones rich and poor in volcanic detritus (Carrigy and Mellon, 1964; Mellon, 1967). In the polymict sandstones and conglomerates of the coal measures in the southeastern part of the Irkutsk Basin, Siberia, a zone with a complex clay-zeolite cement containing montmorillonite, illite-montmorillonite mixed-layers, chlorite, laumontite, and thomsonite, is underlain by 80-230-m thick zone with monomineralic laumontite cement with minor chlorite. The latter is followed at depths of 250-400 m by a zone with predominantly kaolinitic cement. Difference in the type of epigenetic (burial-diagenetic) alteration is ascribed by Koporulin (1961, 1962) to the nature of the interstitial waters. Saline, low-pH waters favor the development of kaolinite cement (1) directly from detrital biotite, (2) from chlorite or montmorillonite alteration products of biotite, or (3) at the expense of zeolite formed earlier under influence of alkaline interstitial waters, particularly in the upper 20-25 m of the kaolinite cement zone. The less permeable shales and siltstones of the sequence have been much less affected by these epigenetic alterations. Throughout the whole sequence they retain a varied clay composition closely reflecting the depositional environment (Koporulin, 1967). Chloritization followed by laumontitization in one unit and hydromicatization following chloritization in other units of the Cretaceous of the Lena Coal Basin, northern Siberia, have been interpreted by Zaporozhtseva ( 1960)

46 3 to reflect the mineralogy of the sediment and the interstitial water composition. Both of the above-mentioned processes are controlled by the sedimentation conditions (see also Hoare et al., 1964). Whatever the controls, kaolinite occurs in several laumontite-zone terranes. As far as known to the author, however, it has not been reported from higher-grade burial-metamorphic facies. Completion of the decomposition of clastic biotite and K-feldspar The diagnostic Ca-Al-silicates of burial metamorphism, as well as secondary Mg-Fe-silicates such as chlorite, form to a major extent at the expense of clastic minerals such as calcic plagioclase, biotite, pyroxenes, and amphiboles. It is of interest to ascertain if the completion of the decomposition of certain clastic minerals can be correlated with particular stages of burial metamorphism. It should be taken into account, however, that the decomposition of relics of igneous and metamorphic origin (including volcanic glass) constitutes a transition from a metastable high-temperature phase to a stable low-grade assemblage, rather than a stable equilibrium reaction. As a result, the extent of relic decomposition is not subject only to the chemical controls, which were found to affect the appearance of one low-grade assemblage at the expense of another, such as the availability and rate of removal of water (i.e., permeability), composition of the fluid phase, etc. To a much greater extent than these reactions, it is dependent on kinetic controls such as duration and intensity of metamorphic recrystallization. If only for this reasons, no stringent correlation between completion of relic decomposition with progressive lowest-grade mineral zones is to be expected. The albitization of plagioclase has been discussed in earlier sections. I t was found to be well-advanced in laumontite zones and to be virtually completed (with minor exceptions) by the appearance of pumpellyite. High carbonate content appeared to inhibit or retard albitization of plagioclase, at least in the laumontite zone. The presence of detrital biotite and K-feldspar in clastic rocks is reported from relatively few burial-metamorphic sequences. This may be partly because no importance is assigned to their presence and, in part, due to the scarcity of these minerals in the intermediate volcanic clastic rocks and tuffs in which the zeolite facies is commonly developed. More or less altered biotite appears to be commonly present in laumontite zone sandstones (e.g., Martini, 1968, p. 592), often together with K-feldspar (e.g., Raam, 1968, p.325; Read and Eisbacher, 1974, pp. 532-535; Stewart and Page, 1974, p. 281). In the Great Valley sequence of Cache Creek, California, the chloritization of clastic biotite in the sandstones appears to increase with depth. It is

464

essentially completed at the same stage as the albitization of calcic plagioclase, at the base of the laumontite zone and before the appearance of widespread prehnite and pumpellyite (Dickinson et al., 1969). On the other hand, Bishop ( 1972b, p. 3186) reported that the replacement of clastic biotite by chlorite and sphene in the metagraywackes of northern Otago proceeds in the prehnite-pumpellyite zone. Both biotite and Kfeldspar disappear only at the boundary of the textural zones 1 and 2A, which is equivalent to the limit of the Chl. 1 and Chl. 2 zones of Hutton and Turner ( 1936), in the low-grade part of the pumpellyite-actinolite facies. Kawachi (1975, p.433) noted the extreme rarity of relic orthoclase in the laumontite-albite and pumpellyite-actinolite zones of the upper Wakatipu district, south Island. These scanty data merely suggest that both biotite and K-feldspar in clastic rocks tend to be completely altered towards or within the pumpellyitic facies of lowest-grade metamorphism. Even more persistent are clastic clinopyroxene and hornblende, which commonly are completely unaltered in the laumontite zone (Coombs, 1954, p. 82; Raam, 1968, p. 325). They are often intact in prehnite- and pumpellyite-graywackes (e.g., Seki et al., 1969a, p. 18), although in the Taveyanne Graywacke, for example, they show a stronger tendency to chloritization in the prehnite- and pumpellyite-graywackes than in the laumontite-graywackes. in which they have almost completely escaped replacement (Martini, 1968, pp. 592, 596). Whereas clastic clinopyroxene and hornblende may be either unaltered or partially chloritized in the prehnite-pumpellyite zone, they are commonly surrounded by an actinolite rim in the pumpellyite-actinolite zone (e.g., Hashimoto and Kashima, 1970; Hashimoto, 1972; Coombs et al., 1970, p. 146; Kawachi, 1975, p. 432). Nevertheless, despite this partial alteration at lower grades, the clastic clinopyroxene and hornblende may persist into the epidote-actinolite zone of the greenschist-facies in the intermediate and mafic tuffs and arenites of several areas (e.g., Hashimoto, 1972, p. 23; Smith, 1969, pp. 153-156). Relation of the laurnontite zone of the zeolite facies to Kossovskayu and Shutov’s “stage of deep epigenesis” and Kubler ’s illite- “crystallinity” zones Kossovskaya and Shutov ( 196 1) have placed the laumontitized sandstones along the western margin of the Verkhoyansk Range in northern Yakutia (cf. Zaporozhtseva, 1963) in their “zone of unchanged argillaceous and authigenic quartz cement”, which constitutes the upper division of the “stage of deep epigenesis”. The disappearance of laumontite in the sandstones of the underlying “zone of chlorite hydromica cement and quartzite-like structure” (which together with the first-mentioned zone constitutes the “stage of deep epigenesis”) in the area is ascribed to the different nature of the clastic

465 components in the lower zone. At any rate, the laumontite does not extend into the higher-grade “metagenesis” or incipient metamorphism. The consequent incorporation of the laumontite “facies” in the “stage of deep epigenesis” has been maintained in subsequent papers (e.g., Kossovskaya and Shutov, 1963; 1970, tableI). No minerals characteristic for the anchizone (or equivalent ”stage of early metagenesis”), such as pyrophyllite, paragonite, or rectorite-allevardite have been reported from any laumontite-bearing terrane. This absence, and the common persistence of illite-montmorillonite mixed-layers (but rarity of discrete montmorillonite), kaolinite, and of clastic biotite, K-feldspar, and calcic plagioclase, into the deeper part of the laumontite zone (cf. preceding section), suggest that the laumontite zone of the zeolite facies may be entirely within the zone of “diagenesis” (in the sense of Kubler, 1967a) and the “stage of deep epigenesis”. Coal ranks in the ‘Stage of deep epigenesis” and Kubfer’szone of “diagenesis”. Moreover, the coal ranks associated with the “stage of ‘deep epigenesis”, as recognized in various parts of the U.S.S.R.,range up to semi-anthracite (see review in Kisch, 1974, pp. 87-90). In the Triassic and Jurassic of the sub-Alpine zone of the French Alps and the Helvetic zone of the Glarus Alps (eastern Switzerland), and in the Upper Paleozoic of the Rheinische Schiefergebirge (western Germany) the illite 10A peak widths, marking the highest-grade part of Kubler’s “diagenetic” zone, appear to be associated with higher-grade ranks, up to 3-4% R,,, ol, or medium-rank anthracitic (cf. Qsch, 1974, table I1 and pp. 111-112; see also Barlier, 1974; Barlier et al., 1974; Wolf, 1975) ’. Subsequent data on coal ranks associated with illite “crystallinities” and “sharpness ratios” (after Weaver, 1961) characteristic for the more advanced stages of the zone of “diagenesis” (after Kubler) appear to corroborate this general relationship: In the Late Cretaceous Buckinghorse Formation of northeastern British Columbia (Foscolos and Kodama, 1974, table 2; Foscolos and Stott, 1975, pp. 22-24) the associated vitrinite reflectances range up to 2.5% R,,,,,,,, (Foscolos et al., 1976, table 5); in the Cambro-Silurian Jamtland Supergroup, western Sweden, the associated reflectances range up to 2.6% R,,,,,,,,, (Kisch, 1980a), corresponding to highest-rank semi-anthraci te. I In some areas that underwent heating by deep plutons, e g . in the Devonian and Carboniferous of the Lippstadt Upwarp in the Rheinische Schiefergebirge, the reflectances associated with high “diagenetic” illite 10 A peak widths are even higher, ranging up to 6% R,,,,i,, or high-rank anthracite (Weber, 1972b. fig. 3; Wolf, 1972, fig. 9; Wolf, 1975; see also Hoyer et al., 1974).

466 Comparison of this range of coal ranks (up to high-rank semi-anthracite or even anthracite) with those found associated with the laumontite zone (absence of ‘higher than low-volatile bituminous ranks-see p. 455) similarly suggests that the latter extends to a somewhat less advanced degree of burial-diagenesis than do the “stage of deep epigenesis” and the zone of “diagenesis” of Kubler. “Illite crystallinities” associated with the laumontite zone in the Helvetic zone. Kisch (1974, p. 100) has suggested that the limit between the laumontite zone and the prehnite-pumpellyite facies in the Taveyanne Graywacke of the Helvetic and sub-Alpine chains is associated with a somewhat lower, lowvolatile bituminous to semi-anthracitic coal rank than is the onset of the anchimetamorphic zone or the equivalent “stage of early metagenesis” in most areas. Frey and Niggli (1971, p. 232) found that the laumontite-bearing Taveyanne Graywacke of the Wageten Range, northwestern Glarus Alps, contains illites with “crystallinities” characteristic for Kubler’s ( 1967a) “diagenetic” zone; and Kubler et al. (1974, p. 466) similarly found “diagenetic” 10 A illite peak widths from Eocene shales and Senonian limestones underlying the laumontite-, prehnite-, and pumpellyite-bearing Taveyanne Graywacke in the Diablerets area. Kubler et al. (1974, p. 466), therefore, concluded that the upper-i.e., the low-grade-limit of the anchimetamorphic zone lies “somewhere in the pumpellyite-prehnite zone”. Subsequent illite 10 A peak-width data from the flysch directly associated with laumontite-bearing Taveyanne Graywacke in the Diablerets area (Durney, 1974, p.271), the middle Kander Valley, and the lower Kien Valley (Kisch, 1980b), and with the stratigraphically equivalent Champsaur Sandstone of the sedimentary cover of the southern part of the Pelvoux massif (Aprahamian, 1974, p. lo), are also “diagenetic”, but locally close to the low-grade limit of the anchizone. This seems to corroborate the conclusions about the relationship of the laumontite zone and the zone of “diagenesis” arrived at in the previous section; however, this relationship requires further investigation in other areas. Relation of the pumpellvite facies to Kossovskaya and Shutov S “stage of early metagenesis” and Kubler’s anchimetamorphic zone Prehnite-pumpellyite facies and the “stage of early metagenesis”. Kossovskaya and Shutov (1961) have proposed correlation of Coombs’ (1960) prehnitepumpellyite facies in New Zealand with the ‘‘zone of chlorite-muscovite cement and spiniform structures” in the Verkhoyansk Range, on the basis of

467

the occurrence of prehnite in the arenites of this zone (laumontite has disappeared in the overlying textural zone). The zones thus correlated have been assigned to the “stage of early metagenesis” (Kossovskaya and Shutov, 1961, table; 1963; 1970, table I ’), which, on the basis of the appearance of pyrophyllite at the expense of dickite, has been correlated by Kubler (1968, pp. 394-395) with the anchizone (see also Kisch, 1974, pp. 83-84). This correlation, however, is somewhat tenuous because no pumpellyite has been reported from the Verkhoyansk Range. The disappearance of laumontite in the overlying textural zone- the “zone of chlorite-hydromica cement and quartzite-like structures”- ascribed by Kossovskaya and Shutov (1961, p.733) to the change in the nature of the clastic feldspar, may alternatively be interpreted to indicate that the high-grade limit of the zeolite facies had already been exceeded, admittedly without concurrent formation of prehnite and pumpellyite. Acceptance and generalization of the suggested correlation of the low-grade boundary of the prehnite-pumpellyite-metagraywacke facies with the onset of the “stage of late metagenesis” (Kubler’s anchzone), , therefore, requires demonstration that it is associated with the appearance of anchizonal illite crystallinities and of diagnostic minerals such as pyrophyllite, paragonite, and rectorite-allevardite.

Phengite and stilpnomelane in pumpellyitic lowest-grade facies. N o detailed clay-mineral data are available for the clay-rich clastic rocks associated with the prehnite-pumpellyite and the pumpellyite-actinolite facies. The phyllosilicates usually reported are chlorite, di-octahedral micas, hydromicas or illites, and, less commonly, celadonite. References to the occurrence of irregular illite-montmorillonite mixed-layers, celadonite, or kaolinite are notable for their absence beyond the high-grade part of the laumontite zone. Reports on the composition and polytype of phyllosilicates present are included in several studies of the mineralogy and petrology of glaucophanelawsonite-schist terranes (e.g., Ernst et al., 1970; Seki et al., 1969). Ubiquitous in the pelitic schists and metasandstones is a white mica of phengitic composition, i.e., rich in Mg and/or Fe, intermediatebetween the muscovite, ferrimuscovite and celadonite compositions, and predominantly showing the 2M polytype. It is a characteristic product of low-temperature, high-pressure metamorphism (Velde, 1965b, 1967).

’

In the text of Kossovskaya and Shutov’s (1970) paper (p. 16), prehnite and pumpellyite are stated to characterize volcanogenic graywackes and intermediate arkoses in the “stage of early metagenesis”, the characteristic textures of which are those of the “zone of chlorite-muscovite cement and spiniform structures” of earlier papers, and of table1 of the same paper. The heading of the section, “zone of quartzitic structures and hydromica-chlorite cement”, seems inappropriate to the text, and appears to this writer to be misplaced.

468 The white micas tend to become more aluminous (i-e., somewhat more paragonitic and muscovitic, and less phengitic) with the increase in degree of metamorphism from the glaucophane-lawsonite-schist facies proper to the glaucophanitic greenschist facies in the “schistes lustres” nappe of the Piemont zone in the Franco-Italian western Alps (Dunoyer de Segonzac, 1969, pp. 185-192; Dunoyer de Segonzac and Hickel, 1972), and to the greenschist facies in the Sanbagawa metamorphic belt of the Shirataki district, Shikoku (Ernst et al., 1970, pp. 155-160). Some of the authigenic K-rich layer silicates of the glaucophane-lawsonite -schist facies are increasingly shown to be also comparatively common in the other pumpellyite-bearing facies. Phengitic composition of the white micas, for instance, has been reported from the prehnite-pumpellyite facies by Hawkins (1967) and in the pumpellyite-actinolite-schist facies by Bishop (1972b, p. 3185). Stilpnomelane is not restricted to the highest- and high-pressure facies series of glaucophanitic terranes. It is common throughout the prehnitepumpellyite and the pumpellyite-actinolite facies of the intermediatepressure series of the Wakatipu metamorphic belt, New Zealand (Coombs, 1960; Bishop, 1972b; Kawachi, 1974, 1975), the Sanbagawa belt of Honshu, Japan (Seki et al., 1971), and the west side of the Appalachian metamorphic belt, U.S.A. (Zen, 1974a). Stilpnomelane is also common in the prehnitepumpellyite facies of the southeastern margin of the Rheinische Schiefergebirge, Germany (Meisl, 1970, p. 59, fig. 70), and the Helvetic zone of eastern Switzerland (Frey, 1970; Frey et al., 1973: cf. Kisch, 1974, pp. 100-lOl), and in quartzite-like sandstones in the “stage of initial metagenesis” of the Russian platform (Veselovskaya, 1967). The appearance of stilpnomelane in the Helvetic nappes of Valais and Glarus appears to be associated with mid-anchizonal illite crystallinities (Frey et al., 1973, p. 201 ; Durney, 1974, p. 27 1). The development of stilpnomelane in Cretaceous glauconitic sandstones (“Gault”), underlain by some 200 m of Taveyanne Graywacke with pumpellyite, prehnite, and laumontite in the southern part of the Massif de Plate, Haute-Savoie (Martini and Vuagnat, 1970, p. 58; Martini, 1972, p. 262; cf. Kisch, 1974, p. lOl), suggests that this mineral may already appear in the highest-grade, pumpellyite-bearing part of the laumontite zone. On the other hand, the stilpnomelane occurrence in the Simme nappe of the Prealpes near Zweisimmen does not constitute appearance of stilpnomelane at a lower grade than prehnite-pumpellyite facies, as erroneously suggested earlier by this writer (Kisch, 1974). Illite “cristallinity ” and diagnostic anchimetamorphic la-ver silicates associated with the prehnite-pumpellvite facies. The relationship between the onset of the

469 characteristics of Kubler’s ( 1967a) anchimetamorphic zone and the prehnite -pumpellyite (and the higher-grade pumpellyite-actinolite) facies has been studied in the Helvetic zone of the Swiss Alps, and in the transitions of the Caples and Torlesse terranes to the Haast schists ( = Otago schists) in Otago, N.Z. The prehnite- and pumpellyite-bearing Taveyanne Graywacke of the Helvetic zone of the Schachen, upper Linth, and Sernft Valleys in eastern Switzerland (Martini and Vuagnat, 1965, fig. 1) lie in the higher-grade part of the anchizone. This is indicated already by the occurrence of characteristic anchimetamorphic minerals such as pyrophyllite, paragonite, regular paragonite/phengite mixed-layer and rectorite, and of illite lOA peak widths of < 5 mm in the Upper Triassic Quartenschiefer and the Lower Liassic (Frey, 1970, fig. 10 and tableI). Moreover, the glauconite of the Lower-Cretaceous and Eocene glauconite-bearing horizons of the Linthal area and the Urnerboden area have been replaced by stilpnomelane (Frey, 1970, fig. 10; Frey et al., 1973, pp. 199-201). However, several of these minerals are from higher tectonic units than the prehnite-pumpellyite-bearing Taveyanne Graywacke (see also Fig. 5-43.). Illite “crystallinities” from the slates directly associated with some of these pumpellyites were measured by Frey and Niggli (1971, p. 232; see also Kubler et al., 1974, p. 466; Kisch, 1974, p. 96) and Kisch (1980b): the 10A peak widths are characteristic for the middle and high-grade part of the anchimetamorphic zone. The coal ranks associated with four occurrences of prehnite-pumpellyite-bearing Taveyanne Graywackes are medium-rank anthracitic, with approx. 3.3 - 4.0% R,,,,,, (Kisch, 1980b; see also p.456). The pumpellyite-actinolite facies Taveyanne Graywacke near Leuk, Valais (Martini and Vuagnat, 1970, p. 57; Coombs et al., 1976), is already associated with “epimetamorphic” 10 A peak widths (Kubler et al., 1974, p. 466; Qsch, 1980b). In the Caples-Pelorus terrane of south Otago, N.Z., Kisch (in preparation) has found predominantly anchimetamorphic illite 10 A peak widths associated with the prehnite-pumpellyite facies, whereas predominance of “epimetamorphic” peak widths seems to occur somewhat before the appearance of unequivocal pumpellyite-actinolite facies mineral assemblages towards the Haast schist terrane to the northeast. In one area near the Torlesse-Haast schist transition in north Otago, “epimetamorphic” 10 A peak widths seem to appear at slightly higher grade, already within the pumpellyite-actinolite facies. In both the Helvetic zone and around the margins of the Haast Schists terrane the prehnite-pumpellyite facies thus appear to be essentially associated with anchimetamorphism, the onset of the mica 10A peak widths characteristic of “epimetamorphism” lying at somewhat lower, similar or

470 TabuIarJura; Northernmost Northern and Southern Autochtonous Cover of tht )oreholes be-Glarus Alps middleGIorus Glorus cover of the Gotthard Alps eastern Aar Massif ow Molasse Alps

m ass if*

A upper Triassic “Quartenschiefer” and lower Liassic lMd>ZMi Mite (and polytype) --- Phengite Kaolinite Pyrophyllite Irregular illitejrnontmorillonite m-1 ? Regular illite /montmorillonite m-l (allev. ?Regulor paragonite jphengite rn-l Parogonite . _ . Regular chlorite/montmorlllonlte rn - I Chlorite Chloritoid Biotite >?omn 7-13 mrn lllite crystallinity**

_ _ _ _ - - -2 M 1 --

2M 1

, 7

---

__

LindaL Frick

Localities studied in Qtail by Frey ngures showing typical textures (Frey. 1969a)

Ouarten

32 B

32A

4.5-7.5rnm

Dia3enesis

Anchimetarnorphism

“Epimetornorphism”

Diagenesis I

I

I

I

8. Mesozoic and Eocene glauconitic hOriZOnS Glauconite Stilpnornelane Riebeckite Biotite 01.

< 4mm

< 4 5rnrn

Lirnrnernbmen Garvera Wnixerpass. Kistenpass 320

~

~

Metamorphic zones (based on illite crystallinity and presence of pyrophyllite and paragon i t e )

Zones of Frey et

4-53mm

Guggenegg. W. of Kobelruns,Wiss Linthal rnilen, Laui furkel 32C

1

I

(1973)

C. Volanic graywackes (“Gres de Taveyanne”) and associated slates Laumontite PurnDellvlte Prehnite Illite crystallinity’* (after Frey and Niggli, 1971)

, I

I

1

Fig. 5-43.Ranges of occurrences of some minerals in the Helvetic zone of eastern Switzerland, compiled after Frey (1969a, 1969b, 1970), Frey and Niggli (1971) and Frey et al. (1973). * Including the southernmost Glarus Alps. ** Low-grade limit of anchizone: 74 mm; of “epizone”: 4 mrn. *** In the Wageten, west of Lake Walen (Frey and Niggli, 1971, p. 232).

somewhat higher-grade than the boundary between that facies and the higher-grade pumpellyite-actinolite facies. Illite “crystallinity and paragonite associated with lawsonite-bearing pumpelbite facies. Data on these crystallinities can be obtained from a comparison of the lowest-grade Alpine metamorphic assemblages of the “zone houillere brianGonnaise” in the Pennine realm of the French Alps with the illite “crystallinites” in its Mesozoic sedimentary cover. Prehnite and pumpellyite are common in the external centra1 massif of the Pelvoux and its sedimentary cover and in the external, westernmost margin I’

47 1 of the “zone houillere brianqonnaise” (Saliot, 1973), accompanied in the Gres de Champsaur (the equivalent of the Taveyanne Graywacke of HauteSavoie and Switzerland) of the southern cover of the Pelvoux by laumontite (Martini and Vuagnat, 1965; Bocquet, 1971) and corrensite (Kubler, 1970; Aprahamian, 1974, p. 10). The illite lOA peak widths associated with the laumontite-bearing southwestern cover are still “diagenetic” (Aprahamian, 1974, p. 10). However, unequivocally anchimetamorphic (in part high-grade anchimetamorphic) illite 10 A peak widths predominate in the autochthonous and para-autochthonous Lower Tertiary flysch cover to the southeast and the northeast of the Pelvoux, whereas those in the Liassic autochthonous and parautochthonous Lower Tertiary flysch cover to the southlimit, and are accompanied by phengite-paragonite mixed-layer or paragonite (Kubler, 1970; Aprahamian, 1974, pp. 10-1 1). In the more internal, eastern part of the “zone houillere”, lawsonite has been reported from microdiorites west of Modane, with prehnite according to Fabre (1961, p.64; cf. Bocquet, 1971, p.98), but according to Bearth (1962) and Guitard and Saliot (1973, pp. 508-5 10) usually with pumpellyite without prehnite or actinolite. Guitard and Saliot (1971) and Saliot (1973) report several additional occurrences of lawsonite and/or pumpellyite from the “zone houillere brianqonnaise” further to the south. Predominance of low-grade “epimetamorphic” illite 10 A peak widths (2.5-3.4 mm) were reported in the Upper-Cretaceous calc-schists (“marbres en plaquettes”) of the sedimentary cover of the central part of the Brianqonnais zone (from around Brianqon in the north to the Guillestre area in the south) by Dunoyer de Segonzac (1969, pp. 176-180); some of the samples contain paragonite. In the more internal eastern margin of the Brianqonnais zone (Col d’Izoard area and Cristillan Valley), the lOA peak widths (2.2-2.7 mm at the conditions used) are higher-grade “epimetamorphic”, and several of the samples contain paragonite. Abbas ( 1974) has similarly found predominance of low-grade “epimetamorphic” lOA peak widths (means of 3.3 and 3.1 mm, respectively) and occurrence of paragonite in two sections of Rhaetian shales and dolomites at Charvie and Haut-Cristillan in the discontinuous “pre-piemontaise” zone which has a structural position intermediate between the internal Brianqonnais zone and the more internal, easterly Piemont zone (“schistes lustres” nappel. Predominance of low-grade anchimetamorphic 10 A peak widths (mean value = 5.1 mm), on the other hand, has been found by Abbas (1974) in the Rhaetian shales and dolomites of the Peyre-Haute nappe (part of the Brianqonnais zone, and overlying the “zone houillere”) at Lac d’Ascension, some 12 km south of Brianqon and 8 km southwest of the Col d’Izoard. One may note that the anthracitic coal ranks associated with the pumpel-

472 lyite- and lawsonite-bearing “zone houillere” (see p. 456) are quite similar to those found associated with prehnite-pumpellyite facies of Taveyanne Graywacke’in central and eastern Switzerland (Kisch, 1980b; cf. p. 456). These relationships appear to indicate that in the Briangonnais zone the pumpellyite- and lawsonite-bearing assemblages are largely associated with “epimetamorphic” illite 10 A peak widths and anthracite coal ranks, and that such “crystallinities” are thus attained well before the appearance of abundant Alpine glaucophane and jadeite to the east. These minerals appear immediately east of the Briangonnais zone sensu stricto, in the Permo-Trias of the “ecailles intermediaires” of Rio Secco and Chaberton (near the Montgenevre Pass, NE of Briangon), underlying the above-mentioned “pre-piemontaise” zone, and the Aceglio zone to the south (Bocquet, 1971, pp. 94-95, 99; Bocquet, 1974, pp. 165, 168; Saliot, 1973, Fig.5-30). The illite peak widths of associated rocks from the upper Valle Varai ta (northwestern Aceglio zone) reported by Dunoyer de Segonzac and Hickel ( 1972) are unequivocally “epimetamorphic”. Unequivocally “epimetamorphic” mica 10 A peak widths and presence of minor paragonite have also been found by this writer (Kisch, in preparation) to be associated with the lawsonite-pumpellyite (without actinolite or prehnite) zone in the Caples Group of the Humboldt Mountains in the upper Wakatipu district, western Otago, N.Z. (zone I1 of Kawachi, 1974, 1975). These very scanty data seem to suggest that at the metamorphic gradients giving rise to the appearance of lawsonite in pumpellyite-facies metamorphic terranes- presumably at low thermal gradients during metamorphism-“epimetamorphic” illite “crystallinities” are attained at a comparatively low degree of incipient metamorphism. This may be an expression of the existence of a pressure effect in enhancing the development of mica crystallinity at a given metamorphic temperature. Relationship of coal rank and authigenic layer silicates to burial-metamorphic facies in volcanic and lithic-feldspathic sediments in different lowest-grade metamorphic ‘yacies series” In the preceding sections the writer has noted the incorporation of the laumontite zone of the zeolite facies in Kossovskaya and Shutov’s “stage of deep epigenesis”. This incorporation, originally suggested by Kossovskaya and Shutov (1961), has been confirmed and refined by comparisons between the maximum coal ranks associated (either directly or as obtained by downward extrapolation) with laumontite-bearing mineral assemblages in various areas (low-volatile to semi-anthracite), and those associated with the “stage of deep epigenesis” or Kubler’s zone of “diagenesis” (or “non-metamorphic” zone) in other areas (Kisch, 1969, 1974). The writer has also arrived at a tentative correlation of the boundary

473

between the laumontite zone and the (zeolite-free) prehnite-pumpellyite facies with the high-grade part of “deep epigenesis”, or the transition between this stage and the higher-grade “early metagenesis” (this transition is equivalent to the “diagenesis”-anchimetamorphism boundary of Kubler). This correlation is based (1) on the above comparisons of ranks associated with laumontite-bearing assemblages, and (2) on direct correlations of zeolite-facies-prehnite-pumpellyite-facies successions with the associated modifications in layer-silicate mineralogy in very few areas (with Kossovskaya and Shutov’s zones in the Permian and Jurassic of the western Verkhoyanye, Siberia, and with illite crystallinity in the Lower Tertiary of the Helvetic zone of Switzerland and Haute-Savoie, and in the Permian of Otago, N.Z.). The correlation requires further refinement. Even then, it is valid only for terranes like the above, where a laumontite zone is succeeded by a zeolite-free prehnite-pumpellyite zone and then by a pumpellyite-actinolite zone, i.e., the intermediate- and high-intermediatepressure type “facies series” of Table 5-VII. Such a relationship cannot automatically be extrapolated to incipientmetamorphic terranes with different temperature-pressure gradients. I t remains to be seen, for instance, if the onset of pumpellyitic facies is associated with similar coal ranks and stages of clay-mineral alteration in such different temperature-pressure gradients-“facies series”-of lowest-grade metamorphism. For instance, incomplete evidence in terranes in which a lawsonite-bearing pumpellyite facies occurs (as a lower-grade predecessor to blueschist facies) indicates that the onset of such pumpellyitic facies could well be associated with different coal ranks and illite “crystallinities”. In the BrianGonnais zone of the French Alps. for instance, the coal ranks associated with the lawsonite-bearing pumpellyite facies appear to be anthracites, but the illite 10A peak widths are already characteristic for “epimetamorphism”. No illite “crystallinity” data are available for the presumably even higher-pressure (and lower-temperature) lawsonite- and pumpellyite-bearing metagraywackes of the Franciscan in the Diablo Range, California, but the associated coal ranks are appreciably lower: high- to low-volatile bituminous. These differences may appear at first sight disconcerting. However, they probably reflect fundamental differences in the way in which the various parameters of burial metamorphism react to changes in temperature and pressure. It may, for instance, reflect the major influence of temperature and duration of heating on the coalification process, pressure having little or no effect, whereas pressure may provide a major contribution to the improved crystallinity of illite and the disappearance of expandable layers from illite-smectite mixed-layers.

474

The differences between the relationship of the various parameters in various incipient-metamorphic “facies series” may, therefore, point the way to further studies, in which these differences will actually be used and evaluated in evaluating temperature-pressure-time gradients of burial metamorphism.

ACKNOWLEDGEMENT

The author is greatly indebted to his colleague Dr. Reg Shagam (BeerSheva) for improving the style of parts of the manuscript, and for his continuous encouragement. This chapter has benefited substantially from the editing by Dr. G.V. Chilingar, who suggested various improvements. REFERENCES * Abbas, M.. 1974. Mktamorphisme des argiles dans le Rhetien des Alpes sud-occidentales. Etude mineralogique et geochimique. These docteur 3eme cycle, Universite Louis Pasteur, Strasbourg, 73 pp. Albee. A.L. and Zen, E-An, 1969. Dependence of the zeolitic facies on the chemical potentials of CO, and H,O. In: V.C. Zharikov (Editor), Contributions to Physico-chemical Petrology (Korzhinskii Vol.), pp. 249-260. Alexandersson, T., 1970. The sedimentary xenoliths from Surtsey: marine sediments lithified on the sea-floor. Surtsey Res. Progr. Rep., V: 83-89. Althaus, E., 1969. Das System Al ,O, -SO, -H,O. Experimentelle Untersuchungen und Folgerungen fur die Petrogenese der metamorphen Gesteine. Teil I: Apparative und experimentelle Grundlagen; die Stabilitatsbedingungen der hydroxylhaltigen Aluminiumsilikate. Neues Jahrb. Mineral., Abh., 111 (1): 74-1 10. Aprahamian, J., 1974. La cristallinite de I’illite et les mineraux argileux en bordure des massifs cristallins externes de Belledonne et du Pelvoux. Geol. Alpine, 50: 5-15. Aronson, J.L. and Hower, J., 1976. Mechanism of burial metamorphism of argillaceous sediment, 2. Radiogenic argon evidence. Geol. SOC.Am. Bull., 87 (5): 738-744. Artru, Ph. and Gauthier, J., 1968. Evolution geometrique et diagenetique d’une serie miogeosynclinale (Lias inferieur a Berriasien) d’apres I’etude du sondage de Valvigneres (France sud-est). Bull. Centre Rech. Pau-SNPA, 2 ( I ) : 101-1 16. Artru, Ph.. Dunoyer de Segonzac, G., Combaz, A. and Giraud, A., 1969. Variations d’origine sedimentaire et evolution diagenetique des caracteres palynologiques et geochimiques des Terres Noires jurassiques en direction de I’Arc Alpin (France, sud-est). Bull. Centre Rech. Pau-SNPA, 3(2): 357-376. Bailey, S.W. and Brown, B.E., 1962. Chlorite polytypism: I. Regular and semi-random one-layer structures. Am. Mineral., 47: 819-850. Banno. S., 1964. Petrologic studies on Sanbagawa crystalline schists in the Bessi-Ino district, central Sikoku, Japan. Unio. Tokyo J . Fac. Sci., 15: 203-319.

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475 Barlier, J., 1974. Recherches Paleothermometriques dans le Domaine des Terms Noires Suhulpines Meridionafes. These docteur 3eme cycle. Univ. Paris-Sud (Centre d’Orsay). 96 pp. Barlier. J., Ragot. J.-P. and Touray, J.-C.. 1974. L’evolution des Terres Noires subalpines meridionales d’apres I’analyse mineralogique des argiles et la reflectomktrie des particules carbonees. Bull. B.R.G.M. (2eme Serie), 5: 533-548. Beall. A.O. Jr. and Ojakangas, R.W., 1967. Mineralogy of an Upper Cambrian K-bentonite from Missouri. J . Sediment. Petrol.. 37(3): 952-976. Bearth, P., 1962. Versuch einer Gliederung alpin-metamorpher Serien der Westalpen. Sdnt.ei:. Mineral. Petrogr. Mitt., 42: 127- 137. Beuf. S., Biju-Duval, B.. Stevaux, J. and Kulbicki. G.. 1966. Ampleur des glaciations “siluriennes” au Sahara: leurs influences et leurs consequences s u r la sedimentation. Rev. Inst. Fr. Petrol., 21: 363-381. Beugnies, A.. Godfriaux, L. and Robaszynski. F.. 1969. Contribution a l’etude des phengites. Bull. SOC.Beige GeoI. Paleontol. Hvdrol.. 77: 95- 146. Biljon, W.J. van and Bensch, J.J.. 1970. The “crystallinity” of illite as a measure of contact metamorphism in mudstone of the Karroo System, South Africa. In: Second Goriht.urur Symposium, Proceedings and Papers. C.S.I.R.. Pretoria. South Africa. pp. 45 1-453. Bishop, D.G., 1972a. Authigenic pumpellyite and other metamorphic effects in the Kveburn Formation, central Otago. N . Z . J . Geol. Geoph.vs.. 15(2): 243-250. Bishop, D.G.. 1972b. Progressive metamorphism from prehnite-pumpellyite to greenschist facies in the Dansey Pass area, Otago. New Zealand. Geol. Soc. Am. Bull.. 83: 3177-3198. Bocquet. J., 1971. Cartes de repartition de quelques mineraux du metamorphisnie alpin dans les Alpes franco-italiennes. Eclogue Geol. Helu.. 64( I ) : 71- 103. Bocquet. J.. 1974. Etudes Mineralogiques cr Petrologiques sur les MPtamorphisnies cl‘A^ge Alpiri dans les Alpes Franqaises. These. Univ. Grenoble. 489 pp. Boles, J.R. and Coombs. D.S., 1975. Mineral reactions in zeolitic Triassic tuff. Hokonui Hills. New Zealand. Geol. Soc. Am. Bull.. 86(2): 163-173. Bostick. N.H., 1971. Thermal alteration of clastic organic particles as an indicator of contact and burial metamorphism in sedimentary rocks. Geosci. Man. 3: 83-92. Bostick. N.H.. 1974. Phytoclasts as indicators of thermal metamorphism. Franciscan assemblage and Great Valley sequence (Upper Mesozoic). California. In: R.R. Dutcher. P.A. Hacquebard. J.M. Schopf and J.A. Simon (Editors). Carbonaceous Materials as Indi(.ator.T of Metamorphism. Geol. Soc. A m . Spec. Pap.. 163: 1-17. Boudier, F. and Nicolas. A,. 1968. Decouverte de chloritoide dans les schistes ardoisiers d’Angers. Bull. Soc. Fr. Mineral. Cristallogr.. 91 : 92-94. Bradley, W.F., 1950. The alternating layer sequence of rectorite. Am. Mineral.. 35: 590-595. Brindley, G.W.. 1956. Allevardite, a swelling double-layer mica mineral. Am. Minerd.. 41: 91-103. Brindley, G.W. and Sandalaki. Z.. 1963. Structure, composition and genesis of some longspacing, mica-like minerals. A m . Mineral., 48: 138- 149. Brown, C.E., 1961. Prehnite-pumpellyite metagraywacke facies of Upper Triassic rocks. Aldrich Mountains. Oregon. U.S. Geol. Surv., Prof: Pap.. 424-C: 146- 147. Brown, C.E. and Thayer, T.P.. 1963. Low-grade mineral facies in Upper Triassic and Lower Jurassic rocks of the Aldrich Mountains, Oregon. J . Sediment. Perrol., 33(2): 41 1-425. Brown, E.H.. 1967. The greenschist facies in part of Eastern Otago, New Zealand. Conrrih. Mineral. Petrol., 14: 259-292. Brown, G. and Weir. A.H.. 1963. The identity of rectorite and allevardite. In!. C1u.v. Cant 1963, Stockholm Proc.. 1. 27-35. Brown, G., Catt, J.A. and Weir, A.H.. 1969. Zeolites of the clinoptilolite-heulandite type in sediments of south-east England. Mineral. Mag., 37(288): 480-488. I

476 Browne, P.R.L. and Ellis, A.J.. 1970. The Ohaki-Broadlands hydrothermal area, New Zealand: mineralogy and related geochemistry. Am. J . Sci., 269: 97-131. Burst, J.F., Jr., 1959. Postdiagenetic clay mineral environmental relationships in the Gulf Coast Eocene. Clays Clqv Miner., Proc. 6th Nail. Confi, pp. 327-341. Burst, J.F., 1969. Diagenesis of Gulf Coast clayey sediments and its possible relation to petroleum migration. Am. Assoc. Pet. Geol. Bull., 53( I ) : 73-93. Buryanova, Ye.Z. and Bogdanov. V.V., 1967. Distribution of the authigenic zeolites laumontite and heulandite in the sedimentary rocks of the Tarbagatai coal deposits. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 1967 (2): 195-202. Bystrom, A.M., 1956. Mineralogy of the Ordovician bentonite beds at Kinnekulle, Sweden. Suer. Geol. Unders. (Ser. C), 540: 62 pp. Caillere, S., Mathieu-Sicaud, A. and Henin, S., 1950. Nouvel essai d'identification du mineral de La Table pres d'Allevard, I'allevardite. Bull. Soc. Fr. MinPral. Cristallogr., 73: 193-201. Cameron, L.B. and Sabine, P.A., 1969. The Tertiary welded-tuff vent agglomerate and associated rocks at Sandy Braes, Co. Antrim. Inst. Geol. Sci., Rep. No. 69/6: 15 pp. Campbell, A.S. and Fyfe, W.S., 1965. Analcime-albite equilibria. Am. J . Sci., 263: 807-816. Carman, M.F., 1965. Nature of chlorite in some low-grade metamorphic rock in South Island, New Zealand. Geol. Soc. Am., Program 1965 Ann. Meet., Kansas City, p. 27. Carrigy, M.A. and Mellon, G.B., 1964. Authigenic clay mineral cements iQ Cretaceous and Tertiary sandstones of Alberta. J . Sediment. Petrol., 34(3): 461 -472. Cassan. J.P. and Lucas, J., 1966. La diagenese des gres argileux d'Hassi-Messaoud (Sahara): silicification et dickitisation. Bull. Seru. Carte GPO/.Alsace Lorraine, 19: 241-253. Castaiio, J.R. and Sparks, D.M., 1974. Interpretation of vitrine reflectance measurements in sedimentary rocks and determination of burial history, using vitrinite reflectance and authigenic minerals. In: R.R. Dutcher, P.A. Hacquebard, J.M. Schopf and J.A. Simon (Editors), Carbonaceous Materials as Indicators of Metamorphism. Geol. Soc. Am. Spec. PUP., 153: 31-52. Chamley, H., 1967. Possibilite d'utilisation de la cristallinite d'un mineral argileux (illite) comme temoin climatique dans les sediments. C.R. Acad. Sci., Ser. D, 265: 184-187. Chamley, H., 1968. Sur le r d e de la fraction sedimentaire issue du continent comme indicateur climatique durant le Quaternaire. C.R. Acad. Sci., Ser. D,267: 1262- 1265. Chateauneuf. J.-J., Debelmas, J., Feys, R., Lemoine, M. and Ragot, J.-P., 1973. Premiers resultats d'une etude des charbons jurassiques de la zone brianqonnaise. C.R. Acad. Sci. Poris, Ser. D,276: 1649-1652. Chatterjee, N.D.. 1973. Low-temperature compatibility relations of the assembalge quartzparagonite and the thermodynamic status of the phase rectorite. Contrib. Mineral. Petrol., 42: 259-271. Chennaux. G. and Dunoyer de Segonzac, G., 1967. Etude petrographique de la pyrophyllite du Silurien et du Devonien au Sahara. Repartition et origine. Bull. Serv. Carre GkoI. Alsace Lorraine, 20(4): 195-210. Chennaux, G., Dunoyer de Segonzac, G. and Petracco, F., 1970. Genese de la pyrophyllite dans le Paleozoique du Sahara occidental. C.R. Acad. Sci., Ser. D, 270: 2405-2408. Clauer, N. and Lucas, J., 1970. Mineralogie de la fraction fine des schistes de Steige-Vosges septentrionales. Bull. Groupe Fr. Argiles, 22: 223-235. Coombs. D.S.. 1954. The nature and alteration of some Triassic sediments from Southland, New Zealand. R. SOC.N.Z. Trans.,82(1):65-109. Coombs, D.S., 1958. Zeolitized tuffs from the Kuttung Glacial Beds near Seaham, New South Wales. Aust. J . Sci., 21: 18-19. Coonibs. D.S.. 1960. Lower grade mineral facies in New Zealand. Rep. In/. Geol. Congr., 21st Sess.. Norden, 13: 339-351.

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495 Appendix A DIAGENESIS OF IRON-RICH ROCKS (Illustrated by the Role of Diagenesis in Oolitic Iron Ores) L. BUBENICEK

INTRODUCTION- DEFINITIONS

Because of its abundance in the earth’ crust, iron is one of the major constituents of rocks and is always present, in varying quantities, in sedimentary formations. The diagenetic evolution, however, may be related to iron-phase changes only in iron-rich rocks. This is particularly true in sedimentary iron ores. It should be kept in mind, however, that the term “ore” is restricted here to iron formations that are economically minable, which implies meeting requirements such as location, thickness, and grade. Iron-rich rocks have more widespread distribution through geological scale and on the earth’ surface than ores. A major difficulty arises when an attempt is made to assign a definite sedimentary origin to ores. It is not sufficient for an iron formation to be interbedded in sedimentary rocks in order to assign a sedimentary origin to iron, because it may be introduced either by solution through replacement of host rocks (substitution type of iron deposits) or through magmatic injection (Kiruna type). The difficulty in assigning an unquestionable sedimentary origin in such cases is the fact that mineralogical changes that occur in iron-rich rocks are so destructive of pre-existent textures and structures that the various stages of evolution cannot be recognized. Iron ores of sedimentary origin have many different physical and chemical properties and have been found in all types of rocks and from all stages of sedimentary evolution. Classically, four fundamental categories, which are presented in order of importance below, may easily be distinguished: (1) The Lake Superior ore type, including low-grade banded ores, B.H.Q, taconites; etc., i.e., different iron-bearing formations (James, 1955); (2) the oolitic ores, with very varied parageneses; (3) the glauconitic ores; (4)the sphaerosiderite ores. The stages of evolution undergone by the sedimentary iron ores before reaching the present state are numerous and varied. Attempts have always been made to understand the mode of formation by also trying to explain the apparent anomaly of the iron mineralization. The diversity of the proposed classifications reflects the complexity of the phenomena. If one attempts to distinguish the role of the main sedimentary processes and in

496 particular the role of diagenesis, however, it is necessary to study all the stages of transformation. Many authors working mainly on considerably modified pre-Hercynian ores have often denied all possible effects of diagenesis, even of epigenesis. They explained the mineral associations solely by the processes of syngenetic precipitation during deposition. Taking into consideration the present state of knowledge, however, it is impossible to ignore diagenesis. I t is an extremely important process, particulary in the case of iron, an element the chemical behavior of which varies greatly according to its ionic state. (See discussions in Grubb, 1971; Dimroth and Chauvel, 1973; Tsu-Ming Han, 1968.) At present, after tremendous amount of studies, the ores of Lake Superior type, which are generally called “iron-formations”, have still not been clearly related to the original state of deposition that remains conjectural. (See discussions and references in James, 1955; Gross, 1965.) The main contribution is in relative mineralogical changes that may be observed through morphological transformations and justified by mineral equilibria. To date, however, no complete balance from depositional to post-diagenetic stage has been determined for any one of those ores. The importance of sphaerosiderites is mainly historical, because these ores were the mainstay of the British iron industry during the 19th century; but no new exploitation seems likely at present. Diagenetic segregation of the siderite from diffused material in clays appears to be the most common present interpretation (Williams et al., 1954; Kazakov, 1957). The glauconite iron ore deposits that could possibly be exploited are very rare; however, the frequent occurrence of these ores has led to some research. The problems in this case are similar to those of the oolitic ores as they have many common features. These are mainly: (1) texture, (2) nature of associated minerals, and (3) depositional environment. The great peculiarity of these ores is that glauconite is of primary origin, and several hypotheses have been proposed to explain this. The oldest involves the activity of Foraminifera. Also the frequent association of glauconite with biotite led to the hypothesis that glauconite is a weathering product of biotite. Some Soviet authors assigned a diagenetic origin to this mineral in a slightly oxidizing environment (Krotov, 1952; Kazakov, 1957; also see Hower, 1961). New categories probably will be added to this list. For example, personal observations convinced the author (Bubenicek, 1968) that in Kiruna ores many residual features may be explained by sedimentary processes, and that the actual state reached by this iron ore may be explained through metamorphic phenomena applied to a sedimentary formation. In the same way, the substitution type of iron ore, which is well illustrated by the Ouenza deposit in Algeria, leads one to believe in replacement of calcareous formation by

497 ferruginous brines at early stages of diagenesis. The brief general survey presented here shows that the research is not sufficiently advanced to allow a satisfactory definition of the diagenetic history of the iron ore deposits, except for the oolitic ores. All possible evolutionary stages may be noted in the latter ores, which occur in all formations from the Precambrian up to the Lower Quaternary and through various stages of evolution between syngenesis and metamorphism. These stages have been studied in detail by Bubenicek ( I 968).

HISTORICAL REVIEW OF LITERATURE ON DIAGENESIS IN OOLITIC IRON ORES

In earlier publications on oolitic iron deposits, emphasis was placed on the process of diagenesis in the formation of the deposits, i.e., the concentration of the iron. Thus, numerous hypotheses have been proposed, involving mineralizing solutions and the replacement of pre-existing rocks. As a result of progressive accumulation of new data, however, it wag concluded that the concentration of iron in a deposit is a phenomenon related to differentiations antedating the deposits. Many different genetical schemes have, therefore, appeared to explain the existing textures. The role assigned to different processes by various authors varies a great deal and depends largely on which observational scheme is regarded as decisive. The different theories may be classified as follows: (I) Iron has been concentrated and deposited during diagenesis. The iron (and silica) has been brought in by mineralizing solutions of various origins (marine, submarine, thermal springs, compactional fluids, etc.), with replacement of oolitic and calcareous detritus by iron solutions containing SiO, and Fe (Castano and Garrels, 1950). ( 1) Without reworking phenomena: the different minerals were directly formed during the replacement processes (Arend, 1933). (2) With reworking phenomena: the different mineralogical types were formed by oxidation during the reworking process (Cayeux, 1909, 1922; Deverin, 1945). (11) Iron ore deposits resulted from a sedimentary differentiation process predating the deposit, i.e., predating sedimentation. (1) With direct precipitation of iron along with other components at the bottom of the sedimentary basin. (a) Direct formation of different iron ore minerals during precipitation. Formation of oolites in situ as a result of diagenesis. The mineralogical distribution depended on the distance from the shore line (Poustovalov, 1940; Caillere and Kraut, 1954, 1956; Braun, 1963; Petranek, 1964). (b) Formation of oolites in situ in minerals having low iron content

498 (e.g., chlorite). Formation of limonite due to oxidation during a reworking stage, with detrital deposition of oxidized oolites. The siderite is considered by some authors to be of diagenetic origin (Berg, 1924, 1944; Hallimond, 1925; Bichelonne and Angot, 1939; Taylor, 1949; Krotov, 1952; Bushinsky, 1956. Without details: Braconnier, 1833). (2) Formation of oolites in dynamic conditions before deposition. (a) The oolites were formed directly from various mineralogical types of iron minerals prior to deposition: (i) as part of primary cement (Borchert, 1952); (ii) alone: cement or a few minerals only being of diagenetic origin (Popov, 1955; Tochilin, 1956; Courty, 1959, 1961 ; Formosova, 1959; Dunham, 1960; Teodorovich, 1961). (b)The oolites are of an oxidized nature; different reduced iron minerals present are the result of diagenetic processes (Brown, 1943; Harder, 1951, 1957; Correns, 1952; Kolbe, 1958; Bubenicek, 1961, 1963). These various hypotheses reflect the treading of a new path of research, as new data were obtained. Every genetic theory, however, should not only give an explanation of the facts, but also allow practical conclusiohs and inferences to be drawn. A brief review of the present knowledge on this subject is presented below.

PRESENT KNOWLEDGE OF THE DIAGENESIS OF OOLITIC FORMS OF ORESCONDITIONS OF DEPOSITION

The data available at present leads one to believe that the sedimentary differentiation of iron is due to two main processes. The first one, of a pedological nature, is associated with the evolution of the continent. The second process takes place in the basin itself, where iron is precipitated and oolites are formed. It appears that there is no direct relation between the influx of iron and the evolution of the basin. Because of the oolitic form, iron is deposited together with various detrital grains, mainly quartz grains or fragments of various shells. At this stage the last possible concentration o f t h e iron takes place by variation of the relative proportions of detrital particles. Consequently, there are deposits exhibiting primary structures of currentbedding and cross-bedding, which reflect current effects. In clay muds, penecontemporaneous deformation features, such as slumping and animal burrowing, are common. The conditions of deposition are such that the iron-bearing oolites contain the most oxidized form of iron (Fe”) in association with oxides of aluminium, phosphorus, and manganese. The constant composition of the material precipitated in the oolitic envelopes (limonite) and the fact that generally in almost all known oolitic

499 iron deposits subsequent changes did not lead to the subtraction or addition of major components, allow a preliminary interpretation of the chemical composition of these ores. Some conclusions can be drawn on the basis of the almost constant amounts of A1,0, and P, along with the dominant iron, especially in the original limonite: (1) In the case of non-clay ores, A1,0, and phosphorus contents’ are proportional to the amount of iron (the more common Al,O,/Fe ratio is 0.10-0.12 in Recent deposits and 0.05 in Paleozoic deposits). This is valid for all the iron contents, which depend on the degree of original dilution with quartz grains or calcite from shell debris; (2) An excessive amount of A1*03 indicates clay ores.

DIAGENETIC EVOLUTION

Limits of diagnesis

When discussing diagenesis, it is always necessary to define the limits of this process, because of lack of agreement among various authors. The present author assigns to diagenesis all the processes which act on the sediment after the end of mechanical movements of the particles 2 , whether during deposition or in seeking a mechanical stability through slumping, and in a milieu having lost all direct relation to the medium of precipitation. Thus, on the one hand diagenesis can start in fine clay sediments that are only a few millimeters thick, whereas in very permeable sands circulation of water coming from the water-sediment interface can preserve syngenetic conditions over a long period of time. Generally, cessation of the mechanical movement of the particles involves isolation of interstitial solutions. The communication between the depositional environment and these solutions exists only through diffusion or by very slow circulation which depends on the permeability of the sediment and on possible ionic composition and charge differences. Circulation occurs on the beaches where the water of waves returns to the sea through the deposited sand3. When the exchange of matter with exterior is low, the

’ ’

For the whole ore body, taking into consideration the whole thickness of a layer, and not only sections where segregation phenomena could be involved. It should be kept in mind, however, that the “compactional diagenesis” inuolutng movement of particles is very important. (Editorial comment.) Cementation and/or solution occurring as a result of this process is considered as part of diagenesis by many authors. (Editorial comment.)

500 “closed milieu” conditions of evolution are approached. The end of diagenesis is more difficult to establish. The present writer places it at the beginning of epigenesis, when hydrated minerals are transformed into less hydrated forms (e.g., limonite to hematite and silicates to micas) with compaction of the rock and loss of its permeability. It is between these limits, in fact, that the oolitic iron ores acquire the essential characteristics of their mineralogical facies. Other transformations, however, may occur beyond this point as has been shown for metamorphosed ores, e g , for hematite and magnetite type ores (see Bubenicek, 1965). Role of diagenesis in the paragenesis of iron-bearing minerals Nature of the diagenetic environment. The diagenetic environment mainly consists of two phases: (1) solid phase: the detrital sedimentary phase where the dominant iron-bearing components are present in the mdst oxidized forms; (2) liquid phase: the pore solutions filling the pore spaces between the grains or the micropores and microfractures of these particles. These solutions are at first identical to those of the sedimentation environment (e.g., sea water). They are, therefore, rich in salts, organic matter and organisms of all kinds. As a result of lack of oxygen, the environment becomes progressively more reducing. This is usually associated with a notable change of pH through temporary acidification (see Bubenicek, 1964). The diagenetic environment has a certain oxidation-reduction potential and capacity, which is a function of the quantity of the confined organic matter or its possible renewal through slow movements of solutions. At a given pH, there are organisms which are able to live under reducing conditions, in particular the sulfate-reducing bacteria. Furthermore, interstitial solutions contain an important reserve of a variety of different ions in variable concentrations. A new diagenetic environment is established more or less rapidly at various depths in the deposit, depending on the type of material and other conditions. In order to define these ideas more accurately, one has to consider the level at which Eh = 0 with respect to the depositional interface: (1)If the level at which Eh=O is within the sediment, its depth will depend on the conditions given above. In this case the changes in the sediment will be of a diagenetic nature. (2) If the level at which Eh = 0 is above the top of the water-sediment interface, deposition of products by direct precipitation from solution could occur. In t h s case one may only apply the term syngenesis, as defined by sedimen tologists.

50 1 Nature of the transformations. The passage from one environment to another, having different physicochemical characteristics, results in a reorganization of the chemical elements in order to establish a new and more stable equilibrium. The previously-formed minerals become unstable and new minerals are precipitated in a stable form, accompanied by exchange of chemical elements (addition or subtraction) with the interstitial solutions. The formation possibilities of the principal iron minerals were discussed by Garrels (1961), who defined the theoretical conditions of their stability (Fig. A-1). Although, theoretical, these diagrams show that all pH-Eh variations involve crossing of stability boundaries which occurs during diagenesis. The following should be noted, however: (1) These diagrams should be expanded by adding areas where the mineral species can exist in a metastable state. (2) The stability diagrams only give an imperfect picture of the reactions, because on assuming the existence of a closed environment the concentrations would vary in accordance with the displacement ,of the reactions. Consequently, threshold to other reactions would appear (Fig. A-2). Main transformations. The main reactions which can occur, as already suspected on studying many deposits, have been confirmed by studies of the Lorraine iron ore deposits. The first fundamental reaction for the non-clay ores is:

limonite

+ quartz

-

siderite

+ chlorite.

The determination of the relative proportions of siderite and chlorite is made possible by the invariance of A1,0, and iron contents and by the knowledge of the Al,O,/Fe ratio in the original limonite. This ratio is equal to 7/52 in the case of the Lorraine limonite. It seems that this ratio is applicable to all Lorraine-type oolitic iron ores. Taking the average chemical composition of the phases given in Table A - I the reaction becomes: 6.05 limonite

+ 1 quartz

+

4.72 siderite

+ 4.26 chlorite.

These values only represent the first attempt to roughly estimate the proportions of the constituents. In this reaction, the retained constituents are A1,0,, Fe, SiO,, Mn, and P; MgO and S are added, whereas H,O is subtracted. The origin of the CO, is open to discussion; however, pyrite forms in the siliceous sediments where calcite is not abundant. This would suggest that at least this CO, is derived from the calcite, and that there would be a corresponding loss of CaO with the gradual destruction of the calcium carbonate. Depending on the initial quartz/limonite ratio, the first threshold would

0 10

N

3

Fig. A-1. Stability diagrams of the main iron minerals under various experimental conditions; T = 25°C; p = 1 atm. (After GARRELS, 1961.) A. Absence of silica and LO,, ZS = 10 '. B. Absence of silica, ZCO, = ZS = C. Presence of silica, CCO, = zs =

503 100

A 50 c .-

c

c 0

0 3

0

A

100

0

Fig. A-2. Diagrams of the evolution of limonite-quartz-calcite ores during diagenetic reduction. These ores have not undergone metamorphism. Lorraine, France; central Great Britain; Jurassic deposits of Germany, etc. (After Bubenicek, 1963.) A. Evolution when quartz is in excess. B. Evolution when limonite is in excess.

TABLE A-I Chemical composition of the iron-bearing phases of the minette of Lorraine

Limonite Chlorite Siderite

52.0 31.9 31.1

4.0 29.1 -

0.5 1.5 5.3

1.2 6.5 5.6

6.0 9.9

0.7 0.2

-

-

11.5 10.9 39.2

504

appear when one of these components had been completely consumed in the reaction forming chlorite and siderite. New transformations then intervene: ( 1 ) In the case of excess of the limonite, there is formation of hematite, siderite and then magnetite. (2) When quartz is present in excess, there is destruction of the first generation of chlorite and formation of siderite, chlorite, and secondary quartz. Equilibrium is reached when the limonite/quartz ratio becomes equal to 6.05. The various parageneses and their evolutionary relationships are presented diagrammatically in Fig. A-2. The limonite and the quartz transformations as a result of fundamental reduction reaction can cease in the case of exhaustion of reducing agent or of one or the other of the original components. If in the initial stage these two components are consumed, an excess of quartz or limonite would be the result. The boundaries beyond which new parageneses occur may be determined through calculations.

Fig. A-3. Calcite concretions. There is a continuity of the laminae between the concretion and the inter-concretionary ore. The inflection of laminae along contact with the two margins indicates a settling of the inter-concretionary ore material and increase in volume of the concretions during the process of CaCO, displacement; X 0.25.

505 Role of diagenesis in modifying structures As noted earlier, the primary and fundamental structures within the oolitic iron ores are current-bedding (most common) and contorted structures. Diagenesis can modify these structures in two different ways: (1) By the general volume shrinkage due to the formation of more compact new minerals. This is followed quite frequently by the development of upright cracks, which can be filled mechanically. (2) By change in the distribution of components (Fig. A-3). usually by segregation of some primary component (e.g., calcite) or of components produced during diagenesis (e.g., siderite and pyrite). These segregation phenomena correspond to the reorganization of the constituents, in order to reach a greater degree of stability. As Ramberg (1952) pointed out, due to surface-energy differences, the free energy of concretion material is lower when the material is concentrated than when it is dispersed. As the stability of a mineral increases with a decrease of free energy, this phenomenon is frequent for minor components (Seibold, 1955; Pettijohn, 11956). Differences in structure or texture can control the arrangement of the concretions, which may be scattered in beds rich in concretions or may be perfectly localized and rounded. The presence of center of preferential attraction and the role of the permeability in the emplacement of concretions and aggregates should also be considered. It must be noted that migration of components may appear to modify the above rules at the level of individual samples.

ROLE OF DIAGENESIS IN THE FORMATION OF TEXTURES

The final texture of the rock reflects all of the above described changes and also those which arise as a result of metamorphism and the effects of meteoric waters, which largely tend to destroy earlier textures. Nonetheless textural studies alone often enable one to determine the history of the physical and chemical reorganization of the components. Three groups of textures due to three fundamental processes may be distinguished: ( 1) Filling textures (cementation textures) due to displacement of material at the time of segregation. At the actual time of crystallization an increase in volume frequently occurs, On the other hand, in impoverished zones physical reorganization is brought about by compaction, quite often with breaking of oolites or plastic deformation. There is frequently secondary growth in the zones of enrichment and also sometimes of saturation, especially in the case of quartz. (2) Solution textures are frequently associated with corrosion textures, and

506 the disappearing minerals are clearly different from the new ones. The disappearance can take place in a progressive or alternating way, as in the case of the transformation of oolitic limonite into chlorite. (3) Growth textures are the fundamental textures which distinguish the growth (by precipitation of whatever origin) of a mineral at a given point and include: (a) Incrustation: cementation on all minerals; the pores may remain open or may be filled completely (the most common case for chlorite; Fig. A-4). (b) Corrosion: with the new mineral filling up gaps left as a result of the disappearance of destabilized minerals (frequent for siderite corroding quartz grains; Fig. A-5). (c) Authigenic (with euhedral shapes): the mineral crystallizes into euhedral shapes and forces away the surrounding material. Siderite crystals could develop crystalline faces against clay minerals, but would only adjust their shape to that of an oolite or its envelope. At the interface between the authigenic crystal and the surrounding material, the effect is comparable to that of corrosion. Siderite shows this very often against clay minerals and

Fig. A-4. Development of chlorite from the limonite of oolites. Chlorite appears as a pellicular cement around the limonitic oolites. The advanced transformation of limonite into chlorite appears clearly on some oolites. Lorraine ore; grey bed=chlorite paragenesis; X 150, natural transmitted light.

507

Fig. A-5. Corrosion of a quartz grain by siderite. The two residual areas of quartz show intended rounded outlines, and have the same optical orientation. The outlines of the quartz-siderite association preserve the shape of the detrital quartz grains: X 150. natural transmitted light.

sometimes against calcite. In ores containing very small amounts of quartz, magnetite (euhedral grains) formed as a result of diagenetic reduction. Effect of diagenesis on the chemical composition of ores As indicated above, the diagenetic reactions which occur after deposition essentially do not involve exchange of the original Fe, A1,0,, P and Mn in solid phase with interstitial solutions. At the most, the role of these solutions is to take part in displacements of material, which are considerable in the case of formation of concretions and slight in the case of authigenic mineral formation. It is different for other elements and compounds, which may indeed undergo exchange with solutions present in the pores of the rock. These include particularly MgO, H,O, CO,, and CaO. The movements of these materials can considerably modify the mass relationships between the elements and greatly change the iron content of the ore. Thus, two ores with the same original iron content could be quite different if diagenesis affected them differently. In both cases, however, the relationships between the elements which were not exchanged will remain unaltered. This permits an

508 adequate comparison of different ores. In order to reduce them to identical conditions, it is advisable to compare them after an ignition loss test. The Fe"/Fe,,,,, ratio is used to determine the evolutionary stage of the ore. In the case of the Lorraine ore, exact knowledge of the changes and of the minerals that are present, has permitted the calculation of the chemical and mineralogical compositions solely on the basis of the determination of Fetora,,Fe" , S O z , A1,0,, and CaO contents (Bubenicek, 1963).

CONCLUSIONS

The present analysis of the role of diagenesis in the formation of iron ores shows the complexity and also the intensity of the phenomena which affect the sediments immediately after their deposition. Many Mesozoic and Tertiary deposits have not been so intensely modified by epigenesis as the older formations and they, therefore, give a clear picture of the fundamental process of diagenesis. All deposits which have been deeply buried or affected by an orogeny, and especially the older deposits, have been subjected to important modifications which frequently mask the earlier textures. It would appear, nevertheless, that the chemical characteristics established during diagenesis (particularly the reduction index FeZt/Fe,,,,,) are preserved for a long time if the deposit has not been subjected to weathering, recent or ancient. REFERENCES A N D BIBLIOGRAPHY Ailing, H.L., 1947. Diagenesis of the Clinton hematite ores of New York. Bull. Geol. Soc. A n ] . . 58: 99-1018. Amstutz. G.C.. 1962. L'origine des gites mineraux concordants dans les roches sedimentaires. Chron. Mines Outre-Mer Rech. Miniere, 308: 1 15- 126. Anderson. G.T. and Han, T.M., 1957. The relationship of Diagenesis, Melamorphisrn and Secondary Oxidation to the Concentrating Characteristics of the Negaunee Iron-Formation of /he Marquette Range. Unpubl. Rep., 4 pp. Arend. J.P.. 1933. Les particularites genetiques du bassin de Briey et leurs rapports avec la repartition, la constitution et les proprietes metallurgiques des minerais oolithiques. Reo. MPt. (Paris). 31: 43-53, 142-151, 188-199, 227-231. Augusthitis, S.S., 1962. Mineralogical and geochemical changes in the diagenetic and postdiagenetic phases of the Ni-Cr-ion oolitic deposit of Larymna/Locris, Greece. Chem. Erde, 2: 1-17. Baas Becking. L.G.M., Kaplan, I.R. and Moore, D., 1960. Limits of the natural environment in terms of pH and oxidation-reduction potentials. J. Geol., 68: 243-284. Berg. G.. 1924. Die Entstehung der sedimentaren Eisenerze. Geol. Rundsch., 15: 96-1 10. Bichrlonne. J. et Angot. P., 1939. Le Bassin Ferrifkre Lorraine, 1. Berger-Levrault, Nancy-Strasbourg. 464 pp.

509 Blondel, F.. 1955. Les types de gisements de fer. Chron. Mines Outre-Mer Re&. Mirii6re. 23 1: 226-246. Borchert. H., 1952. Die Bildungsbedingungen mariner Eisenerzlagerstatten. Chem. Erde. 60: 49-73. Braconnier, M.A., 1883. Description Geologique et Agronomique des Terrains de Meurthe-etMoselle. Berger-Levrault, Nancy-Paris, 436 pp. Braun, H., 1963. Zur Entstehung der marin-sedimentiiren Eisenerze. Borntrager. Berlin. 133 pp. Brown. J.S.. 1943. Supergene magnetite. Econ. Geol.. 38: 137- 148. Bubenicek, L., 1961. Recherches sur la constitution et la repartition des minerais de fer dans I'Aalenien de Lorraine. Sci. Terre, 8( 1-2): 5-204. Bubenicek, L., 1963. Les parageneses des minerais lorraines. R e v . Ind. Miner.. 45(7): 503-506. Bubenicek. L., 1964. L'oxydo-reduction en sedimentologie. Revue synthetique et critique. Bull. Bur. Rech. Geol. Miniere. 4: 36 pp. Bubenicek, L.. 1965. Les parageneses des minerais de fer oolithiques des terrains primaires de I'Ouest de la France. R e v . Ind. Miner.. 47: 1-23. Bubenicek, L., 1966. Geologie des gisements de fer: metallogenie ou geochimie? Minerul. Deposita. I( I ) : 43-55. Bubenicek, L.. 1968. Geologie des minerais de fer oolithiques. Mineral. Depositu. 3: 89- IOX. Bubenicek, L., 1971. Geologie du gisement de fer de Lorraine. Bull. Cent. Rech. Ptru-SNPA. 5(2): 223-320. Bushinskij, I., 1956. Sur la diagenese en rapport avec la genese des argiles refractaires. des minerais de fer sedimentaires et des bauxites. 1:v. A k a d . Nauk S . S . S . R . . Ser. Geol.. 1 1 : 3-115. Cailkre. S. et Kraut, F.. 1954. Les gisements de fer du bassin Lorraine. Meni. Mus. Nuti. Hist. Nut., Ser. C , 4( I ) : 175 pp. Caillere, S. et Kraut, F., 1956. Quelques remarques sur la genese du fer Ordovicien de Segre (Loire inferieure). C . R . Acad. Sci.. 238: 1499-1501. Castano, J.R. and Carrels. R.M.. 1950. Experiments on the deposition of iron with special reference to the Clinton iron ore deposit. Econ. Geo!.. 45: 755-770. Cayeux, L., 1909. Les Minerais de Fer Oolithique de France. I . Minerair de Fers Priniaires. Imprimerie Nationale, Paris, 294 pp. Cayeux, L., 1922. Les Minerais de Fer Oolithiques de France. E. Minerais de Fers Sec,onduire.r. Imprimerie Nationale, Paris, 105 1 pp. Correns. C.W.. 1947. Uber die Bildung des sedimentaren Eisenerze. Forsch. Fortschr.. 2 1 -23(4,5.6): 59-60. Correns, C.W., 1952. Zur Geochemie des Eisens. Congr. Geol. Int. lYme, Al,giers, 1952. C.R.. 2: 23-27. Coutry, G., 1959. Contribution a I'etude du mineral de fer de May-sur-Orne (Calvados), I . Bull. Soc. Geol. Fr.. 5(7): 500-510. Country, G.. 1961. Sur la sideritisation d'oolithes chloriteuses au sommet de la couche de mineral de fer de St. Remy (Calvados). C . R . Acad. Sci.. 252: 301-303. Deverin. L., 1945. Etude petrographique des minerais de fer oolithiques du Dogger des Alpes suisses. Mater. Geol. Suisse, Ser. GPotech., 13(2): 1-1 15. Dimroht, E. and Chauvel, J.J., 1973. Petrography of the Sokoman iron formation in part of the central Labrador Trough, Quebec, Canada. Geol. SOC.A m . Bull., 84: 1 11- 134. Dunham, K.C.. 1960. Syngenetic and diagenetic mineralization in Yorkshire. Proc. Yorks. Geol. Soc., 32(1 I): 229-284. Epprecht, W., 1946. Die Eisen- und Manganerze am Gonzen. Beitr. Geol. Schweiz.. Geotech. Ser.. 24: 1-128.

5 10 Formosova, L.H., 1959. Les Minerais de Fer du PreQrafNord, 1. Geol. Inst. Acad. Sci. USSR., Moscow, 444 pp. Garrels, R.M., 1960. Mineral Equilibria. Harper and Row, New York, N.Y., 254 pp. Gross, G.A., 1965/1967. Geology of Iron Deposits in Canada. 1. General Geology and Evaluation of Iron deposits, 181 pp.; 2. Iron deposits in the Appletrian and Greenwille regions of Canada, 11 pp. Geol. Sum. Can. Econ. Geol. Rep., 22. Grubb, P.L.C., 1971. Silicates and their paragenesis in the Brockman iron formation of Wittenoom Gorge, Western Australia. Econ. Geol., 66: 282-292. Hallimond, A.F., 1925. Iron ores, bedded ores of England and Wales. Geol. Sum. G . B., Mem. Geol. Surv. Spec. Rep. Miner. Resour., G.B., 29: 26-27. Harder, H., 1951. Uber den Mineralbestand und die Entstehung einiger sedirnentarer Eisenerze des Lias y . Heidelb. Beitr. Mineral. Petrogr., 2: 455-476. Harder, H., 1957, Zum Chemismus der Bildung einiger sedimentarer Eisenerze. Z . Dtsch. Geol. Ges., 109(1): 69-72. Harms, J.E., Whitehead, T.H. and Heaton, J.B., 1961. Syngenesis in some Australian iron formations. In: Syngenesis in Ore Deposition-Symp. A . N . Z . A . A . S . , Brisbane, pp. 1-10 (unpubl.). Hough, J.L.. 1958. Fresh-water environment of deposition of Precambrian banded iron formations. J . Sediment. Petrol., 28: 414-430. Hower, J., 1961. Some factors concerning the nature and origin of glauconite. A'm. Mineral., 46: 313-334. James, H.L., 1955. Sedimentary facies of iron formation. Econ. Geol., 49: 235-293. Kazakov, A.V., 1957. La glauconite. Tr. Inst. Geol. Nauk Akad. Nauk S . S . S . R . , Geol. Ser., 152(64): 39- 142. Kazakov, A.V.. Tikhomirova, M.M. et Plotnikova, V.J., 1957. Systerne FeO-C0,-H,O et conclusions sur la paragenese des siderites et phosphorites. Tr. Inst. Geol. Nauk S.S.S.R., Geol. Ser., 152(64): 59-71. Kolbe. H., 1958. Die Erzablagerungen im Salzgittergebiet. Geogr. Rundsch., lO(3): 92-99. Lemoalle, J. and Dupont, B., 1973. Iron bearing oolites and the present conditions of iron sedimentation in Lake Tchad (Africa). In: G.C. Amstutz and A. Bernard (Editors), Ores in Sediments. Springer, Berlin, pp. 167-178. Lindgren, W., 1933. Mineral Deposits. McGraw-Hill, New York, NY, 930 pp. Love, L.G., 1964. Early diagenetic pyrite in fine-grained sediments and the genesis of sulphide ores. In: G.C. Amstutz (Editor), Sedimentologv and Ore Genesis. Elsevier, Amsterdam, pp. 1 1- 19. Krotov, B.P., 1952. Sur le probleme d e la diagenese des sediments: diagenese et teleogenese. Dokl. Akad. Nauk. S . S . S . R . , 82(6): 973-976. McKinstry. H.E.. 1949. Mining Geology. Prentice-Hall, New York. NY, 680 pp. Mohr, P.A.. 1963. Geochemistry of authigenic magnetite from a sedimentary carbonate rock. Univ. CON.Addis Ababa, Fac. Sci., Contrib. Geophys. Obs. Ser. A , 3: 1 I . Oftedahl, C., 1958. A theory of exhalative-sedimentary ores. Geol. Foren. Stockholm Forh., 80(1): 19 pp. Ostroumov. E.A. and Shilov, V.M., 1956. Distribution of sulfides of iron and hydrogen in deep sediments in the northwestern Pacific Ocean. Geochemistry ( U .S. S . R . ) (English transl.). 1960(7): 669-683. Petranek, J.. 1964. Shallow-water origin of Early Paleozoic oolitic iron ores. In: L.M.J.U. van Straaten (Editor), Deltaic and Shallow Marine Deposits. Elsevier, Amsterdam, pp. 3 19-322. Pettijohn. F.J.. 1956. Sedimentary Rocks. Harper, New York, NY, 2nd ed., 718 pp. Popov, B.P.. 1955. Au sujet des carbonates et des silicates dans les minerais de fer de la

511 presqu'ile de Kertch. Tr. lnst. Geol. Nauk, Akud. Nuuk Ukr. R . S . R . . Ser. Prtrogr.. Mineral., Geokhim., 61: 97- 100. Poustovalov, L.V., 1940. Petrographie des Roches Sedimenraires. 1-2. Gostoptekhirdat. Moscow. Ramberg, H., 1952. The origin of Metamorphic and Merasomaric Rocks. 1. Univ. Chicago Press, Chicago, Ill.. 317 pp. Ramdohr, P., 1960. Die Erzmineralien und ihre Veru~uchsungen.Akademie Verlag. Berlin. 1089 PP. Routhier, P., 1963. Les Gisements Metalliferes. Geologie et Priricipes de Recher.che.s. Masson. Pans, 1282 pp. Schnelderhohn, H., 1962. Erzlagerstatten, 4. Fischer, Stuttgart, 37 1 pp. Seibold, E., 1955. Zum Phosphat. Eisen- und Kalkgehalt einiger Horizonte des suddeutschen Jura. Geol. Jahrh., 70( 1955): 577-610. Stanton, R.L., 1964. Textures of stratiform ores. Nature, 202(4928): 173- 174. Strachov, N.M., 1953. La diagenese des sediments et son importance pour la metallogenese sedimentaire. I n . Akad. Nauk S.S.S. R., Ser. Geol.. 5 : 12- 19. Taylor, J.H., 1949. Petrology of the Northampton Sand-Ironstone Formation. Geol. Surv. G. B., M e m . Geol. Suru. G .B., Engl., Wales. 1949: 11 1 pp. Taylor, J.H., 1955. Concentration in sediments. In: Natural Processes of Mineral Concentration. Inter-Univ. Geol. Congr., 3rd, Durham, Proc., pp. 15-20. Teodorovich, G.I., 1958 (1961). Authigenic Minerals i n Sedimentuq Rocks. Consultants Bur.. New York, NY, 120 pp. Tochilin, M.S., 1956. Geochemistry of authigenic siderites. Vopr. Mineral. Osud. O h : . , 3-4: 203-2 1 1. Tsu-Ming Han, 1968. Ore mineral relations in the Cuyuma sulfide deposit. Minnesota. Mineral. Depositu. (3): 109- 134. Williams, H., Turner, F.J. and Gilbert, C.M., 1954. Perrogruphx. Freeman, San Francisco. CA, 406 pp. Zobell, C.E.. 1942. Changes produced by microorganisms in sediments after deposition. J . Sediment. Petrol., 12: 127- 130. Zobell, C.E., 1946. Studies on redox potential of marine sediments. Bull. A m . Assoc. Pet. Geol., 30: 477-513.

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513 Appendix B MINERALOGY A N D PETROLOGY OF BURIAL DIAGENESIS (BURIAL METAMORPHISM) A N D INCIPIENT METAMORPHISM IN CLASTIC ROCKS (Chapter 5)

LITERATURE PUBLISHED SINCE 1976 * (Arranged by subject) HANAN J. KISCH

Changes in clay-mineral assemblages during burial diagenesis and incipient metamorphism: general reoiew

Hoffman, J. and Hower, J.. 1979. Clay mineral assemblages as low grade metamorphic geothermometers: application to the thrust faulted disturbed belt of Montana, U.S.A. In: P.A. Scholle and F.R. Schluger (Editors), Aspects of diagenesis. Soc. Econ. Paleonrol. Mineral., Spec. Publ., 26: 55-79. [Mineral assemblages (principally mixed-layer clays and zeolites) indicate low-grade metamorphism at 100"-2OO0C; this heating is ascribed to burial between thrust plates.] Lippmann, F., 1977. Diagenese und beginnende Metamorphose bei Sedimenten. Bull. Acad. Serbe Sci. Arts, CI. Sci. Nat. Math., 56(15): 49-67. Timofeev, P.P., Kossovskaya, A.G., Shutov, V.D., Bogolyubova. L.I. and Drits, V.A., 1974. New aspects of the study of stages of sedimentary rock development. Lirhol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 9(3): 3 18-336. [Review of diagenesis of clay minerals and of organic matter; very little concerning their correlation.] Experimental and thermodynamical studies and structural transformations of tlarious groups of clqr~minerals (smectite and kaolinite to illite or chlorite)

Lippmann, F., 1979. Stabilitatsbeziehungen der Tonminerale. Neues Jahrb. Mineral. Abh., 136(3): 297-309. Velde, B., 1977. Clays and Clay Minerals in Natural and Synthetic Systems. Elsevier. Amsterdam, 218 pp. [Extensive treatment of all major clay-mineral groups and zeolites.] Frank-Kamenetskij. V.A., Kotov, N.V., Gojlo, Eh. A. and Tomashenko. A.N., 1976. Strukturelle und genetische Wechselbeziehungen zwischen Schichtsilikaten (Phyllosilikaten) und einige Problemen der Tonmineralogie. Z . Angew. Geol., 22: 85-92 [Stepwise transformation of kaolinite and montmorillonite in aqueous and other solutions. Discussion of thermodynamical factors, stability-metastability of the phases formed, and controls by the structure of the starting materials.] Kotov, N.V., Frank-Kamenetsky, V.A. and Goilo, E.A., 1975. Crystal chemistry and thermodynamics of structural transformations of some layer silicates under hydrothermal conditions. Mineral. Pol., 6( 10): 3-27. [Experimental transformation of kaolinite and montmorillonite in the presence of Na, K, Ca and Mg chlorides, sulphates and carbonates under hydrothermal conditions. Inheritance of structure of initial minerals and stability range of the mixed-layer clays are discussed. A dehydration-ionic model is preferred.]

*

The references in this list (completed June 7, 1982, and added in proof) are not inserted in the index of this book.

5 14 Experrmentoi und thermodvnamical studies and structural transformations of dioctahedral three-layer clay minerals

Eberl. D.. 1978. Reaction series for dioctahedral smectites. Clays Clay Miner.. 26: 327-340. [Hydrothermal production of the eight reaction series relating to various dioctahedral clay minerals from beidellite and montmorillonite by making simple changes in interlayer and solution chemistry; significance of assumptions of stability or metastability of the mixed-layer phases for the significance of paragenesis in a reaction series.] Eberl. D.. 1979. Reaction series for dioctahedral smectite: the synthesis of mixed-layer pyrophyllite,’ smectite. In: M.M. Mortland, and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 375-383. [Synthesis from Ca- and Na-saturated montmorillonite in AI3+ solution between 320’ and 400°C.] Eberl, D. and Hower, J., 1976. Kinetics of illite formation. Geol. Soc. Am. Bull., 87: 1326-1330. [Large activation energies suggest that the alteration of smectite to illite involves breaking of chemical bonds in the 2 : 1 layers. The rate of the formation of illite from smectite on the Ocean floor is very slow.] Eberl. D. and Hower, J., 1977. The hydrothermal transformation of sodium and potassium smectite into mixed-layer clay. Clays Clay Miner., 25: 215-227. [Course of the reactions and the appearance of ordered interlayering in the mixed-layer phases are strongly affected by interlayer chemistry ( K versus Na); the difference in hydration energy may account for the fixation of K rather than Na in illite during burial diagenesis.] , Eberl, D., Whitney, G . and Khoury, H., 1978. Hydrothermal reactivity of srnectite. Am. Mineral., 63: 40 1-409. Robertson. H.E. and Lahann, R.W.. 1981. Smectite to illite conversion rates: effects of solution chemistry. Clays Clay Miner., 29(2): 129-135. [Reaction rate and the rate of ordering of mixed-layers were retarded by the addition of N a + , Ca2+ and M g 2 + . ] Velde, B. and Odin, G.S., 1975. Further information related to the origin of glauconite. Clays Clay Miner., 23(5): 376-381. [There is no mineralogical or chemical continuity between illite and glauconite when the K-content is 6 wt,% or greater. Low-grade metamorphism of illitic and glauconitic mica-smectite mixed-layers produces different mica phases.] at 300°C Velde. B. and Weir, A.H., 1979. Synthetic illite in the chemical system K,O-AI,O,-Si0,-H,O and 2 kb. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier. Amsterdam. pp. 395-404. [Limited solid solution in a partially-ordered 1 M rnicaceous mineral; compositions more pyrophylliterich than 80% muscovite contain illite-beidellite interstratification.] Also: -

Velde, 1977, Chap. 4 “Montmorillonites” (ref. p. 513).

.YRD methods for determination of percentage expandables

in

illite-smectite mixed-layers

Johns. W.D. and Kurzweil. H., 1979. Quantitative estimation of illite-smectite mixed phases formed during burial diagenesis. TMPM Tschermaks Mineral. Petrogr. Mitt., 26: 203-215. [Modifications to the methods of Reynolds and Hower (1970) and Perry and Hower (1970).] Rettke. R.C.. 1981. Probable burial diagenetic and provenance effects on Dakota Group clay mineralogy. J . Sediment. Petrol.. 51(2): 541-551. [Use of “saddle”/l7 A intensity ratio and of the I/S composite peak position to determine expandability of I/S mixed-layers.] Schultz. L.G.. 1978. Mixed-layer clay in the Pierre Shale and equivalent rocks, northern Great Plains region. U.S. Geol. Surv.. Prof. Pap., 1064-A: 28 pp, reflection for [Determination of the proportion of illite- and smectite-type layers using the S,,/I,, glycolated material; distinction between beidellite and montmorillonite using the expanding behaviour after the Li+ -200°C-glycerol treatment.]

5 15 Srodon, J., 1980. Precise identification of illite-smectite interstratifications by X-ray powder diffractlon. Clays Clay Miner., 28(6): 401-411. [Methods that take layer-spacing variability of the EG complex of dioctahedral smectites into account. as well as techniques fob quantifying the degree of layer ordering and minimizing the error due to the influence of domain size on the positions of the reflections.] Burial-diageneiic transformation of smectite to K-illite (including regular mixed-layers)

Blank, P. and Seifert, U.. 1976. Zur Untersuchung diagenetischer Tonmineralbildungen und deren experimentelle Modellierung. Z. Angew. Geol., 22( 12): 560-564. [Profiles in sedimentary sections compared with experimental data: importance of cation concentrations is stressed.] Boles, J.R. and Franks, S.G., 1979. Clay diagenesis in the Wilcox sandstones of southwest Texas: implications of smectite diagenesis on sandstone cementation. J . Sedrment. Peirol.. 49: 55-70. [Samples from depths of 975 to 4650 m (representing the temperature range 55°-2100C). Temperatures given for ( I ) disappearance of discrete smectite: (2) disappearance of kaolinite: and ( 3 ) replacement of calcite cement by ankerite. Smectites with high ( F e Mg)/AI ratios appear to resist conversion to illite until temperatures high enough to produce ordering are attained.] Eslinger, E. and Sellars, B., 1981. Evidence for the formation of illite from smectite during burial metamorphism in the Belt Supergroup. Clark Fork. Idaho. J . Sedrment. Petrol., 51( I ) : 203-216. [The ratio K-feldspar/plagioclase decreases, and the illite/quartz ratio inSreases downward in a 8500 m section. K-rich 1Md illite formed from K-poor smectite rather than from direct weathering of feldspar.] Eslinger, E., Highsmith, P.. Albers, D. and DeMayo, B., 1979. Role of iron reduction in the conversion of smectite to illite in bentonites in the Disturbed Belt, Montana. Clqvs Clur Miner.. 27(5): 327-338. [Increase in Fe2+/Fe3+ with increasing percentage of illite layers is tentatively attributed to a redox reaction involving the oxidation of organic matter.] Heling, D., 1974. Diagenetic alteration of smectite in argillaceous sediments of the Rhinegraben (SW Germany). Sedimentologv, 2 1 : 463-472. [Different temperatures of disappearance of smectite in different formations related to differences in permeability, and hence of availability of potassium ions.] Heling, D.. 1978. Diagenesis of illite in argillaceous sediments of the Rhinegraben. Clay Miner.. 13: 21 1 . [Alteration of smectite to illite depends primarily on temperature X time. Potassium is supplied by decomposition of feldspar, rather than from distant sources. Mite diagenesis is affected by the inherited layer charge of the initial smectite.] Jonas, E.C., 1975. Crystal chemistry of diagenesis in 2 : 1 clay minerals. In: S.W. Bailey (Editor). Proceedings of ihe International Clyv Conference 1975. Applied Publishing Ltd.. Wilmette. Ill.. pp. 3-13. Lahann, R.W.. 1980. Smectite diagenesis and sandstone cement: the effect of reaction temperature. J . Sediment. Petrol., 50(3): 755-760. [The temperature at which illitization proceeds may control the spatial distribution zones of cementation through the effect of temperature upon the distance of solution transport of silica. Percentage expandable layers versus T are given for four wells in the Gulf Coast.] Nadeau. P.H. and Reynolds, R.C., 1981. Burial and contact metamorphism in the Mancos Shale. Clays Clay Miner.. 29(4): 249-259. [Correlation of percentage expandable layers in illite-smectite mixed-layer with coal ranks and Laramide tectonic activity. Presence of carbonate inhibits illitization. The use of mixed-layered illite-smectite compositions to infer thermal regimes is misleading unless allowance is made for local chemical controls.] McDowell. S.D. and Elders. W.A., 1980. Authigenic layer silicate minerals in borehole Elmore 1. Salton Sea Geothermal Field, California. U.S.A. Contrrh. Mineral. Petrol.. 74: 293-3 10. [Decrease in percentage expandable layers in illite-smectite mixed-layer phase decreases from 10- 15% at 185'C (41 1.5 m depth) to 6% at 210' (494 m); no expandable layers below 725 m. The recrystallized white mica below 850m depth tends to a progressively more muscovitic composition.]

+

5 16 Rettke, R.C., 1981. Probable burial diagenetic and provenance effects on Dakota Group clay mineralogy. J. Sediment. Petrol., 51(2): 541-551. [Ordered illite-smectites are attributed to provenance differences, which gave lower-expandability illite-smectite a “head start” of diagenetic trends.] Schultz, L.G., 1978. Mixed-layer clay in the Pierre Shale and equivalent rocks, northern Great Plains region. U.S. Geol. Surv., Pro6 Pap., 1064-A: 28 pp. [Beidellite- and montmorillonite-type layers are distinguished; generally 20-60% illite layers. In the Montana disturbed belt much of the clay has been altered to regularly interlayered IS-ordered mixed layer with 60-85% illite layers.] Thorez, J. and Pirlet, H., 1979. Petrology of K-bentonite beds in the carbonate series of the Visean and Tournaisian stages of Belgium. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 323-332. [Mainly illite-smectite and illite-vermiculite mixed-layers; volcanic origin.] Yeh, H.-W., 1980. D/H ratios and late-stage dehydration of shales during burial. Geochim. Cosmochirn. A d a , 44 (2): 341-352. [Significant D/H fractionation between residual and expelled pore waters. The montmorillonite to illite transformation during burial diagenesis is considered to be the most important mechanism of late-stage dehydration.] Also: -

Aoyagi and Kazama, 1980 (ref. p. 529).

- Foscolos and Powell, 1979 (ref. p. 516). - Hoffman and Hower, 1976 (ref. p. 513). -

Kubler, 1980 (ref. p. 517).

- Powell et al., 1978 (ref. p. 517). - Velde, 1977, Chap. 5, “Illite, montmorillonite and mixed layered minerals in sequences of buried rocks ( P - T space)” (ref. p. 513). Kaolinite-group minerals: polyrypism and transformation to illite

Heling, D., 1980. Tonmineraldiagenese und Palaotemperaturen im gebleichten mittleren Buntsandstein am Westrand des Rheingrabens. Neues Jahrb. Mineral. Monatsh., 1980(1): 1- 10. [Near the western fault of the Rhinegraben the “Bunter” is bleached, and the clay fraction is almost entirely 2M illite (the Bunter normally has IM illite with some 50% kaolinite). This is ascribed to high paleothermal gradients and heat transfer by ascending waters along deep faults during periods of rapid subsidence (e.g., the Oligocene).] Rodionova, A.E. and Koval’skaya, M.S., 1974. Dickite distribution in coal-bearing formations of the Donets Basin. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.), 9(4): 75 1-755. [Two dickite-forming stages are recognized: (1) in late epigenesis, and (2) in the post-inversion stage (“regressive epigenesis”).] Relationship between burial-diagenetic modification of clay mineralogy (except illite crystallinity) and of organic matter, including hydrocarbon generation

Gindorf, L. and Paetz, H., 1979. Wechselbeziehungen zwischen Organiten und Anorganiten als Indikatoren geologischer Prozesse wahrend der superkrustalen Gesteinsbildung. Z. Geol. Wiss. (DDR), 7(2): 235-240. Foscolos, A.E. and Powell, T.G., 1979. Mineralogical and geochemical transformation of clays during burial-diagenesis (catagenesis): relation to oil generation. In: M.M. Mortland and V.C. Farmer (Editors), International Clay Conference 1978. Elsevier, Amsterdam, pp. 26 1-270. [First dehydration of the interstratified clays coincides with 0.5% R, vitrinite reflectance and occurs several thousand feet above the main phase of oil generation, whereas the second dehydration step takes place below the oil generating zone.] Heroux, Y.,Chagnon, A. and Bertrand, R., 1979. Compilation and correlation of major thermal maturation indicators. Am. Assoc. Pet. Geol. Bull., 63(12): 2128-2144. [Correlation chart of the most commonly used organic and mineral thermal maturation indicators.]

517 Kubler, B., 1980. Les premiers stades de la diagenese organique et de la diagenese minerale. Deuxieme partie: Zoneographie par les transformations mineralogiques, comparaison avec la reflectance de la vitrinite, les extraits organiques et les gaz adsorbes. Bull. Ver. Schweiz. Pet.-Geol. -Ing., 46( 110): 1-22. Powell, T.G., Foscolos, A.E., Gunther, P.R. and Snowdon, L.R.. 1978. Diagenesis of organic matter and fine clay minerals: a comparative study. Geochim. Cosmochim. Acta, 42: 1181-1 197. [Relationships in subsurface samples from Canadian Northwest Territories. Mixed-layer clays comprising smectite-vermiculite-illite are transformed during thermal diagenesis to smectitevermiculite-illite-chlorite. The first clay dehydration occurs at 0.5% R prior to hydrocarbon generation. Vermiculite is an intermediary in the smectite-illite transformation. and in the (common) presence of Ca2+ ions delays the second dehydration step to the zone where cracking of liquid hydrocarbons to gas occurs (between 1.O- 1.2 and 1.4% R o,l ).] Zhelinskii, V.M.. 1980. Catagenesis of terrigenous rocks and metamorphism of coals in south Yakutia. Lirhoi. Miner. Resour. (transl. from Lifol. Polezn. Iskop.), 15(2): 187- 199. [Lateral variation of relationship between alteration of minerals and coal rank is correlated to effects of magmatic bodies or hydrothermal processes.] Burial-diagenetic transformation of smectite and disappearance of kaolinite fwithouf rllire cn.sral1init.v) in sections with coal-rank data

Chudaev, O.V., 1978. Occurrence of clay minerals in flyschoid sediments of eastern Kamchatka. Lithol. Miner. Resour. (transl. from Litol. Polezn. Iskop.). 13(I): 89-97. [Upper part of volcano-sedimentary sequence, with montmorillonite and randdm chlorite-montmorillonite mixed-layers, and middle part, with corrensite-like minerals and “defective chlorite”, have respectively long-flame coal and gas coal, and are assigned to initial epigenesis; the lower part. with chlorite and hydromica, and coking coal, to the “zone of hypogenesis”.] Hutcheon, I., Oldershaw, A. and Ghent, E.D., 1980. Diagenesis of Cretaceous sandstones of Kootenay Formation at Elk Valley (southeastern British Columbia) and Mt Allan (southwestern Alberta). Geochim. Cosmochim. Acta, 44: 1425- 1435. [At Elk Valley the authigenic kaolinite-dolomite assemblage is associated with vitrinite reflectances from 0.8 to at least 1.6% R,,, oil; at Mt Allan kaolinite disappears at 265 m depth (at approx. 1.4% R o,l), and chlorite calcite appear. Reaction is considered to involve a CO, fluid which is immiscible with water under the extrapolated diagenetic conditions.] Ivanova, N.V., Volkova, A.N., Rekshinskaya, L.G. and Konysheva, R.A., 1980, Pyroclastic material in coal measures of the Donets Basin