The Triassic timescale

The Triassic Timescale

The Geological Society of London Books Editorial Committee Chief Editor

BOB PANKHURST (UK) Society Books Editors

JOHN GREGORY (UK) JIM GRIFFITHS (UK) JOHN HOWE (UK) RICK LAW (USA) PHIL LEAT (UK) NICK ROBINS (UK) RANDELL STEPHENSON (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) MAARTEN DE WIT (SOUTH AFRICA )

IUGS/GSL publishing agreement This volume is published under an agreement between the International Union of Geological Sciences and the Geological Society of London and arises from the ICS Sub Commission on Triassic Stratigraphy. GSL is the publisher of choice for books related to IUGS activities, and the IUGS receives a royalty for all books published under this agreement. Books published under this agreement are subject to the Society’s standard rigorous proposal and manuscript review procedures.

It is recommended that reference to all or part of this book should be made in one of the following ways: LUCAS , S. G. (ed.) 2010. The Triassic Timescale. Geological Society, London, Special Publications, 334. O’DOGHERTY , L., CARTER , E. S., GORICˇ AN , Sˇ. & DUMITRICA , P. 2010. Triassic radiolarian biostratigraphy. In: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 163– 200.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 334

The Triassic Timescale

EDITED BY

SPENCER G. LUCAS New Mexico Museum of Natural History and Science, USA

2010 Published by The Geological Society London

THE GEOLOGICAL SOCIETY The Geological Society of London (GSL) was founded in 1807. It is the oldest national geological society in the world and the largest in Europe. It was incorporated under Royal Charter in 1825 and is Registered Charity 210161. The Society is the UK national learned and professional society for geology with a worldwide Fellowship (FGS) of over 9000. The Society has the power to confer Chartered status on suitably qualified Fellows, and about 2000 of the Fellowship carry the title (CGeol). Chartered Geologists may also obtain the equivalent European title, European Geologist (EurGeol). One fifth of the Society’s fellowship resides outside the UK. To find out more about the Society, log on to www.geolsoc.org.uk. The Geological Society Publishing House (Bath, UK) produces the Society’s international journals and books, and acts as European distributor for selected publications of the American Association of Petroleum Geologists (AAPG), the Indonesian Petroleum Association (IPA), the Geological Society of America (GSA), the Society for Sedimentary Geology (SEPM) and the Geologists’ Association (GA). Joint marketing agreements ensure that GSL Fellows may purchase these societies’ publications at a discount. The Society’s online bookshop (accessible from www.geolsoc.org.uk) offers secure book purchasing with your credit or debit card. To find out about joining the Society and benefiting from substantial discounts on publications of GSL and other societies worldwide, consult www.geolsoc.org.uk, or contact the Fellowship Department at: The Geological Society, Burlington House, Piccadilly, London W1J 0BG: Tel. þ44 (0)20 7434 9944; Fax þ44 (0)20 7439 8975; E-mail: [email protected]. For information about the Society’s meetings, consult Events on www.geolsoc.org.uk. To find out more about the Society’s Corporate Affiliates Scheme, write to [email protected]. Published by The Geological Society from: The Geological Society Publishing House, Unit 7, Brassmill Enterprise Centre, Brassmill Lane, Bath BA1 3JN, UK (Orders: Tel. þ44 (0)1225 445046, Fax þ44 (0)1225 442836) Online bookshop: www.geolsoc.org.uk/bookshop The publishers make no representation, express or implied, with regard to the accuracy of the information contained in this book and cannot accept any legal responsibility for any errors or omissions that may be made. # The Geological Society of London 2010. All rights reserved. No reproduction, copy or transmission of this publication may be made without written permission. No paragraph of this publication may be reproduced, copied or transmitted save with the provisions of The Copyright Licensing Agency Ltd, Saffron House, 6 –10 Kirby Street, London EC1N 8TS UK. Users registered with the Copyright Clearance Center, 222 Rosewood Drive, Danvers, MA 01923, USA: the item-fee code for this publication is 0305-8719/10/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-296-0 Typeset by Techset Composition Ltd, Salisbury, UK Printed by Antony Rowe, Chippenham, UK Distributors North America For trade and institutional orders: The Geological Society, c/o AIDC, 82 Winter Sport Lane, Williston, VT 05495, USA Orders: Tel. þ1 800-972-9892 Fax þ1 802-864-7626 E-mail: [email protected] For individual and corporate orders: AAPG Bookstore, PO Box 979, Tulsa, OK 74101-0979, USA Orders: Tel. þ1 918-584-2555 Fax þ1 918-560-2652 E-mail: [email protected] Website: http://bookstore.aapg.org India Affiliated East-West Press Private Ltd, Marketing Division, G-1/16 Ansari Road, Darya Ganj, New Delhi 110 002, India Orders: Tel. þ91 11 2327-9113/2326-4180 Fax þ91 11 2326-0538 E-mail: [email protected]

Contents LUCAS , S. G. The Triassic timescale: an introduction

1

LUCAS , S. G. The Triassic chronostratigraphic scale: history and status

17

MUNDIL , R., PA´ LFY , J., RENNE , P. R. & BRACK , P. The Triassic timescale: new constraints and a review of geochronological data

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HOUNSLOW , M. W. & MUTTONI , G. The geomagnetic polarity timescale for the Triassic: linkage to stage boundary definitions

61

TANNER , L. H. The Triassic isotope record

103

TANNER , L. H. Cyclostratigraphic record of the Triassic: a critical examination

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ORCHARD , M. J. Triassic conodonts and their role in stage boundary definition

139

O’DOGHERTY , L., CARTER , E. S., GORICˇ AN , Sˇ. & DUMITRICA , P. Triassic radiolarian biostratigraphy

163

MC ROBERTS , C. A. Biochronology of Triassic bivalves

201

BALINI , M., LUCAS , S. G., JENKS , J. F. & SPIELMANN , J. A. Triassic ammonoid biostratigraphy: an overview

221

KU¨ RSCHNER , W. M. & WALDEMAAR HERNGREEN , G. F. Triassic palynology of central and northwestern Europe: a review of palynofloral diversity patterns and biostratigraphic subdivisions

263

CIRILLI , S. Upper Triassic –lowermost Jurassic palynology and palynostratigraphy: a review

285

KOZUR , H. W. & WEEMS , R. E. The biostratigraphic importance of conchostracans in the continental Triassic of the northern hemisphere

315

KLEIN , H. & LUCAS , S. G. Tetrapod footprints – their use in biostratigraphy and biochronology of the Triassic

419

LUCAS , S. G. The Triassic timescale based on nonmarine tetrapod biostratigraphy and biochronology

447

Index

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The Triassic timescale: an introduction SPENCER G. LUCAS New Mexico Museum of Natural History and Science, 1801 Mountain Road N. W., Albuquerque, New Mexico 87104 (e-mail: [email protected]) Abstract: German geologists began to study rocks now recognized as Triassic during the late 1700s. In 1823, one of those German geologists, a very astute mining engineer named Friedrich August von Alberti (1795–1878), coined the term ‘Trias formation’ for an c. 1 km thick, tripartite succession of strata in southwestern Germany – the Bunten Sandsteins, Muschelkalk and Keuper of the German miners. Alberti also recognized Triassic rocks outside of Germany, throughout much of Europe and as far away as India and the United States. By the end of the nineteenth century, Triassic rocks had been identified across Europe and Asia, and in North America, South America and Africa. Indeed, in 1895, the Austrian geologist Edmund von Mojsisovics (1839– 1907) and his collaborators published a complete subdivision of Triassic time based on ammonoid biostratigraphy and, in so doing, introduced many of the Triassic chronostratigraphic terms still used today. The twentieth century saw the elaboration of an ammonoid-based Triassic timescale, especially due to the work of Canadian palaeontologist E. Timothy Tozer (1928-). During the last few decades, work also began on developing a global magnetic polarity timescale for the Triassic, a variety of precise numerical ages tied to reliable Triassic biostratigraphy have been determined, and conodont biostratigraphy has become an important tool in Triassic chronostratigraphic definition and correlations. The current Triassic chronostratigraphic scale is a hierarchy of three series (Lower, Middle, Upper) divided into seven stages (Lower ¼ Induan, Olenekian; Middle ¼ Anisian, Ladinian; and Upper ¼ Carnian, Norian, Rhaetian) further divided into 15 substages (Induan ¼ upper Griesbachian, Dienerian; Olenekian ¼ Smithian, Spathian; Anisian ¼ Aegean, Bithynian, Pelsonian, Illyrian; Ladinian ¼ Fassanian, Longobardian; Carnian ¼ Julian, Tuvalian; Norian ¼ Lacian, Alaunian, Sevatian). Ammonoid and conodont biostratigraphies provide the primary basis for the chronostratigraphy. A sparse but growing database of precise radioisotopic ages support these calibrations: base of Triassic c. 252 Ma, base Olenekian c. 251 Ma, base Anisian c. 247 Ma, base Ladinian c. 242 Ma, base Jurassic c. 201 Ma. A U/Pb age of c. 231 Ma from the Italian Pignola 2 section is lower Tuvalian, and U/Pb ages on detrital zircons from the nonmarine Chinle Group of the western USA of c. 219 Ma are in strata of late Carnian (Tuvalian) age based on the biostratigraphy of palynomorphs, conchostracans and tetrapods. These data support placement of the Norian base at c. 217 Ma, and indicate that the Tuvalian is more than 10 million years long and that the Carnian and Norian are the longest Triassic stages. Magnetostratigraphic data establish normal polarity for all of the Triassic stage bases except Anisian and Ladinian. An integrated biostratigraphic correlation web for the marine Triassic consists of ammonoids, bivalves, radiolarians and conodonts, whereas a similar web exists for the nonmarine Triassic using palynomorphs, conchostracans and tetrapods. Critical to cross correlation of the two webs is the Triassic section in the Germanic basin, where a confident correlation of nonmarine biostratigraphy to Triassic stage boundaries has been achieved. The major paths forward in development of the Triassic timescale are: finish formal definition of all Triassic stage boundaries, formally define the 15 Triassic substages, improve the integration of the Triassic biostratigraphic webs and develop new radioisotopic and magnetostratigraphic data, particularly for the Late Triassic.

Today, the Subcommission on Triassic Stratigraphy (STS; part of the IUGS International Commission on Stratigraphy) advocates a Triassic chronostratigraphic scale of three series (never a subject of debate) and seven stages (much debated) (Fig. 1). The boundaries of the Triassic System are defined by global stratotype sections and points (GSSPs), and the numerical ages of those boundaries appear to be determined with a precision of about 1%. Nevertheless, much work remains to be done to refine the Triassic timescale. Precise numerical age control within the Triassic is generally sparse

and uneven, and a global polarity timescale for the Triassic is far from established. Chronostratigraphic definitions of most of the 15 Triassic substages widely used today remain unfinished, and many issues of marine biostratigraphy are still unresolved. In the nonmarine Triassic realm, much progress has been made in correlation, especially using palynomorphs, conchostracans and tetrapods (amphibians and reptiles), but many problems of correlation remain, especially the cross correlation of nonmarine and marine chronologies.

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 1 –16. DOI: 10.1144/SP334.1 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Britain (Warrington et al. 1980). By the 1860s Austrian geologist Edmund von Mojsisovics began constructing a detailed Triassic chronostratigraphy based on ammonoid biostratigraphy. In 1895, Mojsisovics and his principal collaborators, Wilhelm Waagen and Carl Diener, published a Triassic timescale that contains most of the stage and substage names still used today (Mojsisovics et al. 1895). Spath (1934) proposed a Triassic ammonoid-based biochronological timescale congruent with that of Mojsisovics et al. (1895). Tozer (e.g. 1965, 1967, 1984, 1994) proposed a Triassic ammonoid-based timescale based on North American standards, particularly in the Canadian Arctic islands and the Cordillera of British Columbia and Nevada. Distinctive features of Tozer’s timescale included proposal of four Lower Triassic stages (Griesbachian, Dienerian, Smithian and Spathian) and abandonment of the Rhaetian as the youngest Triassic stage. The STS began its work in the 1970s and now recognizes seven Triassic stages in three series (Fig. 1). Strata pivotal to the development of a Triassic timescale were originally spread across much of the Boreal and Tethyan periphery of Pangaea during the Triassic (Fig. 2). The 1990s saw the rise of Triassic conodont biostratigraphy so that four agreed on (or nearly agreed on) Triassic GSSPs use conodont events as defining features. Most of the bases of the Triassic stages have been (or will soon be) defined by GSSPs: Fig. 1. The Triassic chronostratigraphic scale.

This book reviews the state-of-the-art of the Triassic timescale, and this introductory chapter provides an overview of this book. It also presents a Triassic timescale based on the data presented in several chapters in this book.

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Triassic chronostratigraphy In this volume, Lucas (2010a) reviews the nearly two-century-long development of the Triassic chronostratigraphic scale, which is now a hierarchy of three series, seven stages and 15 substages developed during nearly two centuries of research (Fig. 1). The first geological studies of Triassic rocks began in Germany in the late 1700s and culminated when Alberti (1834) coined the term Trias for the Bunten Sandsteins, Muschelkalk and Keuper of southwestern Germany, an c. 1 km thick succession of strata between the Zechstein (Permian) and the Lias (Jurassic). Recognition of the Trias outside of Germany soon followed, and included early work on a similar succession of Triassic strata in Great

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The base of the Induan Stage (¼ base of Triassic, ¼ base of Lower Triassic) is defined by the lowest occurrence (LO) of the conodont Hindeodus parvus at the Meishan section in Guangxi, southern China (Yin 1996; Yin et al. 1996, 2001). The base of the Olenekian Stage may be defined by the LO of the conodont Neospathodus waageni at the Mud section in Spiti, India (Krystyn et al. 2007a), though this is still under discussion. The base of the Anisian Stage (¼ base of the Middle Triassic) may be defined by the LO of the conodont Chiosella timorensis at the Des¸li Caira section in Romania (Orchard et al. 2007). The base of the Ladinian Stage is defined by the LO of the ammonoid Eoprotrachyceras curioni at the Bagolino section in Italy (Brack et al. 2005). The base of the Carnian Stage (¼ base of the Upper Triassic) is to be defined by the LO of the ammonoid Daxatina canadensis at the Stuores Wiesen section in Italy (Mietto et al. 2007; Gaetani 2009). The base of the Norian Stage is the farthest from decision, with a GSSP located either at

INTRODUCTION

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Fig. 2. Triassic world map with areas pivotal to construction of a Triassic timescale indicated. Artwork by Matt Celeskey.

7.

8.

Black Bear Ridge in British Columbia, Canada or at Pizzo Mondelo in Sicily, and it probably will be based on a conodont event close to the base of the Stikinoceras kerri ammonoid zone, which has been the traditional Norian base in North American usage (M. Orchard, written communication, 2009). The base of the Rhaetian Stage is to be defined by the LO of the conodont Misikella posthernsteini at the Steinbergkogel section in Austria (Krystyn et al. 2007b). The base of the Hettangian Stage (¼ base of the Jurassic, ¼ base of the Lower Jurassic) is to be defined by the LO of the ammonoid Psiloceras spelae at the Kuhjoch section in Austria (Von Hillebrandt et al. 2007).

These GSSPs define boundaries of the seven Triassic stages recognized by the STS and also define the boundaries of the three Triassic Series and of the Triassic System. Most of the bases of the 15 Triassic substages (Fig. 1), however, still lack formal definition. They provide a more refined subdivision of Triassic time than do the stages, and should be the focus of future chronostratigraphic research.

Radioisotopic ages A precise and detailed numerical timescale does not yet exist for the Triassic. This is partly because of the relatively low level of Triassic volcanism recorded in fossiliferous rocks, which resulted in a dearth of datable volcanic ash beds (in contrast to

some of the other geological systems such as the Cretaceous, which had a much more extensive record of volcanism). Nevertheless, some important advances have been made in the last two decades. The Early and Middle Triassic have the best numerical age constraints, and these demonstrate that the Early Triassic only represents about five million years, whereas the Middle Triassic is about 10 million years long; the Late Triassic thus is more than two-thirds of the entire duration of the Triassic. In this volume, Mundil et al. (2010) review the Triassic numerical timescale to produce a significantly different calibration than that published in the most recent compilation by Ogg (2004). The differences mostly reflect the availability of new radioisotopic ages, but some of them also reflect different selection criteria and different approaches to attempting to eliminate the biases (both systematic and random) in the ages. The ages Mundil et al. (2010) use to calibrate the Triassic timescale are from U –Pb analyses applied to zircons with uncertainties at the permil level or better. According to their compilation, the age of the beginning of the Triassic is 252.3 Ma, and the end of the Triassic is 201.5 Ma. Robust age constraints also exist for the Induan– Olenekian boundary (251.2 Ma) and the Early–Middle Triassic (Olenekian–Anisian) boundary (247.2 Ma), so the Early Triassic is approximately five million years long. The AnisianLadinian boundary is constrained to 242.0 Ma by new U –Pb and 40Ar/39Ar ages reported by Mundil et al. Nevertheless, radioisotopic ages for the Late Triassic are scarce, and the only reliable and biostratigraphically controlled age is from a Carnian (lower

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Tuvalian) tuff dated at 230.9 Ma. This means that the Late Triassic is about 30 million years long.

Magnetostratigraphy The global polarity timescale for rocks of Late Jurassic, Cretaceous and Cenozoic age provides a valuable tool for evaluating and refining correlations that are based primarily on radioisotopic ages or biostratigraphy. However, there is no agreed geomagnetic polarity timescale (GPTS) for the Triassic, although a composite GPTS is now becoming available based on successions cobbled together from marine and nonmarine sections in North America, Europe, and Asia. Hounslow & Muttoni (2010) review Triassic magnetic polarity history in this volume. They note that Lower Triassic magnetostratigraphy is primarily calibrated by ammonoid biostratigraphy in Canada and Svalbard. In addition, extensive magnetostratigraphic studies of the Permian– Triassic and Olenekian –Anisian boundaries are calibrated by conodont biostratigraphy. Various magnetostratigraphic studies of nonmarine Lower Triassic strata validate and cross-correlate the marine-based ages into some nonmarine successions. The Middle Triassic magnetostratigraphic timescale is well constrained by conodont and ammonoid zonations from multiple Tethyan sections, and it is consistent with detailed data from several nonmarine Anisian sections. The middle Carnian is the only significant interval in the Triassic for which biostratigraphic calibration of the magnetostratigraphy is not well resolved. Problems with the Norian and early Rhaetian magnetostratigraphy focus on properly constraining the magnetostratigraphic correlation between nonmarine strata, such as the Newark Supergroup of eastern North America, and the polarity timescale based on marine Tethyan sections. Hounslow & Muttoni’s (2010) review concludes that average magnetozone duration is about 240,000 years for the Lower and Middle Triassic, and about twice that for the Upper Triassic. In sum, they recognize 133 valid magnetozones for the Triassic Period.

Isotope stratigraphy In this volume, Tanner (2010a) reviews the use of isotopes in Triassic stratigraphy. As he notes, measurements of d13C, d18O, d34S and 87Sr/86Sr provide information about the state of the water column in which deposition took place. The most widely studied isotope is d13C, and, indeed, the carbon isotope record for the Triassic System is now known in some detail, and it is complex. Thus, a pronounced negative excursion begins

below the base of the Triassic and continues into the lowermost Triassic. Isotopic instability characterizes most of the Lower Triassic, with positive and negative excursions continuing through the basal Middle Triassic. Unlike the Lower Triassic, relative isotopic stability characterizes much of the Middle and Upper Triassic, with rising values of d13C likely due to environmental recovery (after the end-Permian mass extinction) and increasing storage of organic carbon in terrestrial environments. A pronounced negative excursion near the Triassic –Jurassic boundary has been linked to significant biotic turnover. The causes of the Triassic carbon isotope excursions remain a topic of discussion, with the most likely mechanisms being outgassing during volcanic activity, changes in productivity, ocean anoxia, and seafloor methane releases. These processes evidently perturbed the global carbon cycle and forced episodic biotic extinctions. The construction of a global carbon isotope curve for the Triassic is thus well underway. This curve, with some judicious calibration, should become an increasingly important tool for Triassic correlation. However, isotope curves, like magnetostratigraphy, are not independent correlation tools and always need to be tied to biostratigraphic or radioisotopic data in order to be of value in correlation.

Cyclostratigraphy In this volume, Tanner (2010b) reviews the use of cycles in Triassic stratigraphy, and he notes that high frequency (fourth- and fifth-order) cyclicity is a common feature of sedimentary sequences in Triassic depositional settings. Tectonism and autocyclicity clearly drove some of this cyclicity, but many Triassic cycles have been attributed to orbital-forced variations in solar insolation at the Milankovitch frequencies (the precession, as well as the short and long eccentricity cycles at scales of tens of thousands to hundreds of thousands of years). This orbital forcing is thought to have controlled sedimentation through periodic changes in climate or sea-level. Examples of interpreted Milankovitch-frequency cyclicity throughout the Triassic record include much of the Germanic Triassic section, the Newark Supergroup of eastern North America, and parts of the Alpine Triassic. The cyclostratigraphy of these sections has been used as a tool for intrabasinal correlation and for chronostratigraphy. However, conceptual arguments and radioisotopic age data call some of these interpretations into question. At present, Triassic cyclostratigraphic studies remain far from the goal of developing a reliable, astronomically-calibrated Triassic timescale.

INTRODUCTION

Biostratigraphy The distribution of fossils in marine Triassic strata has provided the primary basis for construction of the Triassic timescale. The most important taxa in this regard are conodonts, radiolarians, bivalves and ammonoids. Nonmarine Triassic biostratigraphy has also been developed, based primarily on palynomorphs, conchostracans, tetrapod (amphibian and reptile) footprints and tetrapod body fossils.

Conodonts Conodonts are microscopic tooth-like structures composed of calcium phosphate that are abundant and widespread in Triassic marine strata. Although the biological source of conodonts was long unknown, they are now clearly associated with chordates. In this volume, Orchard (2010) reviews Triassic conodont biostratigraphy, which now plays a pivotal role in the delineation of a Triassic timescale. The base of the Triassic (base of the Induan) is defined by the LO of the conodont Hindeodus parvus; a parallel zonation is provided by Neogondolella species. The LO of Neospathodus waageni within a plexus of similar species is favoured to define the Olenekian base, accompanied by species of Borinella and Eurygnathodus (Krystyn et al. 2007a). The base of the Anisian is close to the LO of Chiosella, with Triassospathodus and Spathicuspus characterizing the late Olenekian, and Gladigondolella tethydis and Nicoraella restricted to the Anisian. The early evolution of Budurovignathus species provides a proxy for the base of the Ladinian, which is defined by an ammonoid event. The Carnian base, also defined by an ammonoid event, is close to the LO of Metapolygnathus (including M. polygnathiformis and M. tadpole). The base of the Norian is characterized by a conodont faunal turnover featuring many new species, as well as an abundance of M. primitius and M. echinatus in the basal Norian. The Rhaetian base will be defined by the evolution of Misikella in Europe, with coeval changes recognized in North American Epigondolella. Conodont extinction was long thought to mark the end of the Triassic, but recently published data indicate a minimal survival of conodonts into the basal Jurassic (Pa´lfy et al. 2007).

Radiolarians Radiolarians are marine zooplankton which secrete a skeleton of opaline silica. In the Modern oceans they form massive skeletal accumulations (radiolarian oozes) on the seafloor in deep waters (up to 4000 m deep). Their Triassic fossils are typically

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found in deep-marine deposits associated with chert horizons. In this volume, O’Dogherty et al. (2010) review the Triassic radiolarian record. They summarize 30 years of research on Triassic radiolarian biostratigraphy and present a correlation of radiolarian zonations currently used in Europe, Japan, Siberia and North America. O’Dogherty et al.’s up-to-date assessment of the stratigraphic ranges of the 281 valid Triassic genera of radiolarians indicates that they are useful in substage-level correlations. After the end-Permian extinction, the most severe extinction in radiolarian history, a long recovery until the early Anisian was followed by a rapid diversification. Maximum generic diversity was during the early Carnian when the first severe within Triassic extinctions took place. Diversity declined through the Norian and Rhaetian, culminated by a mass extinction of radiolarians at the Triassic –Jurassic boundary. One of the most complete Triassic radiolarian biozonations is for Japan, largely a function of the extensive Triassic record of radiolarites in Japan and the intensity of study (Sugiyama 1997). Extensive radiolarian biostratigraphy also exists for western Europe (Kozur & Mo¨stler 1994, 1996; Kozur et al. 1996) and for eastern Siberia (Bragin 1991). Further development of the Triassic record of radiolarians promises to make them a very robust tool in Triassic marine correlations.

Bivalves Late Palaeozoic seas were dominated by pelmatozoans, brachiopods and bryozoans, but molluscs dominated the Triassic seas (Vermeij 1977; Sepkoski 1981). Bivalves (pelecypods) were common Mesozoic molluscs that underwent a substantial Triassic diversification to dominate many level-bottom, reefal and pelagic settings (e.g. McRoberts 2001; Fraiser & Bottjer 2007). Earliest Triassic bivalve assemblages are mostly epifaunal pteriomorphs and detritus-feeding nuculoids, and they are very abundant as fossils. The Middle–Late Triassic saw a diversification of arcoid, mytiloid, trigonioid and veneroid genera. The thin-shelled bivalves Claraia, Peribositria, Enteropleura, Daonella, Aparimella, Halobia and Monotis (the so-called ‘flat clams’ because of their very thin shells and narrow valve convexity) are characteristic Triassic forms widely used in biostratigraphy. In this volume, McRoberts (2010) reviews the application of the ‘flat clams’ to Triassic biostratigraphy and biochronology, noting that these bivalves generally are widely distributed and have very high species turnover rates, making them excellent biostratigraphic indexes of portions of the Triassic.

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Their biostratigraphic value has long been recognized, and McRoberts (2010) reviews previous zonations and proposes his own, which is based on the first global summary of Triassic bivalve zonation. In McRoberts’ (2010) zonation, the Lower Triassic encompasses two to three Claraia zones that represent the Induan and lower Olenekian. In the upper Olenekian, species of Peribositria are useful zonal indexes. During the Middle Triassic, Enteropleura (middle Anisian) and Daonella (upper Anisian through Ladinian) have significant records in the circum-Pacific and Boreal realms. McRoberts divides the Upper Triassic into 8 to 13 bivalve zones based on the succession of species of Halobia, Eomonotis, and Monotis sensu lato.

Ammonoids Most prominent of the Triassic molluscs were ammonoid cephalopods, a group whose rapid diversification during the Triassic provides a fossil record that has long subdivided Triassic time. Most Triassic ammonoids were ceratitidans, with relatively simple suture lines, descended from only two ammonoid stocks that survived the end-Permian mass extinction: the otoceratids and the xenodiscids. Triassic ammonoid genera define at minimum three broad marine palaeobiogeographic provinces around the Pangaean periphery (Tethyan, Boreal and notal), but the ammonoid palaeobiogeography of Triassic Panthalassa was complex and remains little understood (e.g. Dagys 1988). Ammonoids have long been the workhorses of Triassic marine biostratigraphy, and most of the Triassic timescale was built on ammonoid biostratigraphy. In this volume, Lucas (2010a) reviews the historical development of the Triassic chronostratigraphic scale, which is largely a review of the application of ammonoids to Triassic chronostratigraphy. Also, in this volume, Balini et al. (2010) review Triassic ammonoid biostratigraphy. The study of Triassic ammonoids began during the late 1700s, and Mojsisovics et al. (1895) published an essentially complete Triassic chronostratigraphic scale based on ammonoid biostratigraphy. This scale introduced many of the Triassic stage and substage names still used today, and all terminology of stages and substages subsequently introduced has been based on ammonoid biostratigraphy. Early Triassic ammonoids show a trend from cosmopolitanism (Induan) to latitudinal differentiation (Olenekian), and the four Lower Triassic substage (Griesbachian, Dienerian, Smithian and Spathian) boundaries are globally correlated by widespread ammonoid biotic events. Middle Triassic ammonoids have provinciality similar to that of the Olenekian and provide a basis for recognizing six Middle Triassic substages. Late Triassic

ammonoids provide a basis for recognizing three stages divided into five substages. Significantly, the main uncertainty for the future of Triassic ammonoid biostratigraphy is not the decline of the ammonoids as a tool for dating and correlation of Triassic strata, but rather the dramatic decrease in the number of palaeontologists who study Triassic ammonoids, due to the lack of replacement of experienced specialists who started their activity in the 1950s and 1960s.

Other marine biostratigraphy Some other Triassic marine fossils have been useful in biostratigraphy, but have not provided the robust biostratigraphies of the Triassic given by radiolarians, conodonts, bivalves and ammonoids. These other fossils include foraminiferans, brachiopods and nautiloids. Triassic foraminiferans are particularly well studied in Europe and in the former Soviet Union, where a complete Triassic zonation based on foraminiferans has been proposed for the Caucasus (e.g. Efimova 1991; Vuks 2000, 2007). Much more sampling and study remains to determine the global biostratigraphic utility of Triassic foraminiferans. Triassic nautiloid cephalopods appear to have undergone relatively little change at the end of the Permian, but reached great diversity in the Triassic to suffer an extensive (but not complete) extinction near the end of the period. Some Triassic biostratigraphic zonations based on nautiloids have been proposed (e.g. Kummel 1953; Gradinaru & Sobolev 2006), but the rarity of nautiloids in most Triassic facies limits their value to correlation. Brachiopods did not suffer total extinction at the end of the Permian, although their numbers were greatly reduced, but they were relatively minor but persistent components of Triassic marine faunas. A diverse published biostratigraphy of Triassic brachiopods (e.g. Dagys 1974; Hoover 1991; Benatov 2001; Chen et al. 2005; and Shen et al. 2006 represent a small sampling) suggests a fair amount of provincialism and relatively slow turnover rates, which limit the broad applicability of brachiopods to problems of Triassic correlation.

Palynomorphs Spores and pollen are the microscopic reproductive structures of vascular plants. They have organic walls that resist pressure, desiccation and microbial decomposition, so they are often well preserved in sedimentary rocks, and Triassic strata are no exception. Because of their abundance (one plant may produce thousands of palynomorphs), durability and easy dispersal (often by wind), palynomorphs

INTRODUCTION

are found in both nonmarine and marine strata, so they provide an important means for cross correlation of nonmarine and marine strata based on shared palynomorph taxa. On the other hand, most palynomorphs are only dispersed within a few km or less of the plant that produces them, and any provincialization of the paleoflora hinders their use in broad scale correlation. Furthermore, plants are very environmentally sensitive, so palaeoenvironmental and facies restrictions of extinct plants affect the distribution of their palynomorphs. During the Triassic, the palaeoflora was provincialized into at least three or more provinces (see below), and correlations based on fossil plants and the palynoflora have even proven to be difficult between portions of Triassic Pangaea that were in relatively close proximity, such as Western Europe and eastern North America. In this volume, Ku¨rschner & Herngreen (2010) discuss the microfloral trends during the Triassic in the Germanic and Alpine domains, emphasizing diversity fluctuations and related palaeoenvironmental changes. They also propose a set of nine palynomorph zones (with several subzonal subdivisions) for the European (mostly Germanic) Triassic that they correlate to the Triassic marine timescale based largely on the correlation of the German section proposed by Kozur & Bachmann (2005, 2008) (Fig. 3). Cirilli (2010) reviews Upper Triassic –Lower Jurassic palynological assemblages and palynozonations for the Northern and Southern hemispheres. She examines the evolutionary progression of palynological assemblages in the Tethyan domain to conclude that there has been a gradual change in palynofloral composition from the Carnian to uppermost Rhaetian/lower Hettangian. According to Cirilli, the biostratigraphic resolution based on Upper Triassic palynological assemblages is rather low due to the rarity of palynomorphs associated with other means of age determination (i.e. ammonoids, conodonts, isotopes, palaeomagnetism), microfloristic palaeoprovinciality, palaeoenvironmental conditions and the varied preservational grade of Upper Triassic palynological assemblages.

Conchostracans Conchostracans are bivalved crustaceans that live in freshwater lakes and ponds. Their minute, drought-resistant eggs can be dispersed by the wind, and this guaranteed a broad geographic range to some conchostracan taxa across much of Triassic Pangaea. Triassic conchostracan biostratigraphy has been developed by several workers, most recently by Kozur & Weems (2005, 2007) in Europe and North America.

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In this volume, Kozur & Weems (2010) present a revised Triassic conchostracan biostratigraphy. Again, key to their correlation is the Germanic Basin Triassic section, which yields numerous conchostracan assemblages from the Buntsandstein and Keuper. According to Kozur & Weems, Triassic conchostracan zones often provide a stratigraphic resolution comparable to ammonoid and conodont zones of the marine Triassic. Kozur & Weems thus present a conchostracan zonation for the Late Permian –Early Jurassic of the Northern hemisphere and correlate it to the marine timescale. This conchostracan zonation is especially well developed for the Changhsingian to lower Anisian, upper Ladinian to Julian and Rhaetian to Hettangian intervals, but remains preliminary for most of the Middle and Upper Triassic.

Tetrapod footprints Fossil footprints of Triassic tetrapods, which have been studied since the early 1800s, are common in some Triassic nonmarine strata and have very broad palaeogeographic distributions. Furthermore, some Triassic nonmarine strata that lack or nearly lack a tetrapod bone record have an extensive footprint record. Therefore, various workers have used Triassic tetrapod footprints in biostratigraphy, beginning with the pioneering work of Haubold (1969, 1971a, b; Demathieu & Haubold 1972, 1974). In this volume, Klein & Lucas (2010) review the use of tetrapod footprints in Triassic biostratigraphy. They argue that several characteristic Triassic footprint assemblages and ichnotaxa have restricted stratigraphic ranges and thus represent distinct time intervals. Klein & Lucas thus identify five distinct Triassic tetrapod-footprint-based biochrons: (1) dicynodont tracks (Lootsbergian); (2) Protochirotherium (Synaptichnium), also includes Rhynchosauroides and Procolophonichnium (Nonesian); (3) Chirotherium barthii, also includes C. sickleri, Isochirotherium, Synaptichnium, Rotodactylus, Rhynchosauroides, Procolophonichnium, dicynodont tracks and Capitosauroides (Nonesian –Perovkan); (4) Atreipus-Grallator (“Coelurosaurichnus”), which also includes Synaptichnium, Isochirotherium, Sphingopus, Parachirotherium, Rhynchosauroides and Procolophonichnium (Perovkan –Berdyankian); and (5) Brachychirotherium, which also includes Atreipus-Grallator, Grallator, Eubrontes, Apatopus, Rhynchosauroides and dicynodont tracks (Otischalkian–Apachean). Tetrapod footprints prove useful for Triassic biostratigraphy and biochronology, but compared to the tetrapod body fossil record with eight biochrons, the five footprint-based biochrons provide less temporal resolution. Nevertheless, in

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Fig. 3. Summary of the Germanic Basin Triassic section correlated to the Triassic chronostratigraphic scale (modified from Kozur & Bachmann 2005, 2008).

INTRODUCTION

nonmarine Triassic strata where body fossils are rare, tetrapod footprints can be useful for biostratigraphy and biochronology.

Tetrapods Triassic tetrapod (amphibian and reptile) fossils have long been used in nonmarine biostratigraphy, with a tradition extending back to at least the 1870s. Lucas (1990) advocated developing a global Triassic timescale based on tetrapod evolution, and Lucas (1998) presented a comprehensive global Triassic tetrapod biochronology that divided the Triassic into eight time intervals (land-vertebrate faunachrons: LVFs) based on tetrapod evolution. In this volume, Lucas (2010b) presents the current status of the Triassic tetrapod-based timescale. The Early Triassic tetrapod LVFs, Lootsbergian and Nonesian, have characteristic tetrapod assemblages in the Karoo basin of South Africa, the Lystrosaurus assemblage zone and the lower twothirds of the Cynognathus assemblage zone. The Middle Triassic LVFs, Perovkan and Berdyankian, have characteristic assemblages from the Russian Ural foreland basin, the tetrapod assemblages of the Donguz and the Bukobay svitas (‘formations’). The Late Triassic LVFs, Otischalkian, Adamanian, Revueltian and Apachean, have characteristic assemblages in the Chinle basin of the western USA, the tetrapod assemblages of the Colorado City Formation of Texas, Blue Mesa Member of the Petrified Forest Formation in Arizona, and the Bull Canyon and the Redonda formations in New Mexico. Since the Triassic LVFs were introduced, subdivision of several of them has been proposed: Lootsbergian can be divided into three sub-LVFs, Nonesian into two, Adamanian into two and Revueltian into three. However, the broad correlation of most of these sub-LVFs remains to be demonstrated. Records of nonmarine Triassic tetrapods in marine strata, palynostratigraphy, conchostracan biostratigraphy, magnetostratigraphy and radioisotopic ages provide some basis for correlation of the LVFs to the standard global chronostratigraphic scale. These data indicate that Lootsbergian ¼ uppermost Changhsingian, Induan and possibly earliest Olenekian; Nonesian ¼ most of the Olenekian; Perovkan ¼ most of the Anisian; Berdyankian ¼ latest Anisian? and Ladinian; Otischalkian ¼ early –earliest late Carnian; Adamanian ¼ late Carnian; Revueltian ¼ early –middle Norian; and Apachean ¼ late Norian – Rhaetian.

Other nonmarine biostratigraphy Some other nonmarine Triassic fossils have been employed in biostratigraphy, including charophytes, megafossil plants, ostracods, bivalves and fishes.

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None of these groups has provided what can be considered a robust global or even provincial biostratigraphy, but all have some potential to aid in Triassic correlations. Charophytes are the calcified egg cases (gyrogonites) of characeous algae, and they have been documented from some Triassic lacustrine deposits (e.g. Kaesler & Feist 2005). However, too little is known of the Triassic charophyte record to allow its use in biostratigraphy. Furthermore, much of the variation in gyrogonite morphology is ecophenotypic, and therefore more a function of environmental variation than a consistent evolutionary signal. Thus, I suspect that the long-ranging charophyte genera now known from the Triassic (e.g. Stellatochara) will not segregate into temporally successive species useful in biostratigraphy. During the Permian and Triassic, there was a complex and prolonged global change from pteridophyte-dominated floras of the Palaeozoic to the gymnosperm-dominated floras that characterized much of the Mesozoic, which is when the arborescent lycopods and sphenopsids of the Permian gave way to Triassic floras dominated by seed ferns, ginkgophytes, cycads, cycadeoids and conifers. Distinct Gondwanan and Laurasian floras can be recognized, and within Laurasia two or three provinces are recognized during the Triassic – a more boreal Siberian province and a more equatorial Euramerican province. The Triassic Laurasian floras were dominated by primitive conifers, ferns, cycads, bennettitaleans and sphenopsids. Conifers were the dominant large trees, whereas the other plant types formed the understory. In coastal settings, stands of the lycopsid Pleuromeia were dominant. Gondwana floras of the Triassic were dominated by a wide range of seed ferns, especially the genus Dicroidium. These floras were generally composed of only a few (no more than ten) genera. Dicroidium was dominant in a variety of vegetation types, from heath to broad-leaved forest to dry woodland. Other important elements of Gondwana floras were conifers and some Laurasian groups of cycadaleans and ginkgos. Near the end of the Triassic, the Dicroidium flora declined and was replaced by a cosmopolitan conifer –benettitalean flora. Megafossil plant biostratigraphy of the Triassic has been developed at regional and more global scales. At the regional scale, a good example is Ash (1980, 1987), who proposed a succession of three floral zones useful in correlating North American Upper Triassic palaeofloras. Despite subsequent range extensions and stratigraphic refinements (Axesmith & Kroehler 1988; Lucas 1997, 2006), these zones do allow broad correlation of lower Chinle Group and lower Newark Supergroup palaeofloras, but the zonal index taxa extend no farther palaeogeographically. A cautionary tale

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Fig. 4. Early Triassic timescale.

INTRODUCTION 11

Fig. 5. Middle Triassic timescale.

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Fig. 6. Late Triassic timescale.

that demonstrates the facies dependence of megafossil plants is provided by Ash (1980), who considered the Santa Clara flora of Sonora, Mexico to be Rhaeto –Liassic, whereas intercalated marine invertebrates clearly indicate it is of Carnian age (Alencaster de Cserna 1961). At a broader scale, Dobruskina (1994) reviewed the Triassic megafossil plants of Eurasia, comparing them to the global record. She recognized at least three floral provinces (‘phytochoria’) during the Triassic –Siberian, European –Sinian and Gondwanan provinces. At the scale of her review, three temporally successive global floral zones can be recognized: Induan–Anisian, Ladinian–Carnian and Norian –Rhaetian. Ostracods are bivalved crustaceans having calcareous shells. Most nonmarine Triassic ostracods

are darwinulaceans: simple, bean-shaped and unornamented. Although they have provided the basis for some Triassic biostratigraphy (e.g. Pang 1993), their taxonomy strikes me as unreliable: there are not enough morphological characters preserved to reliably recognize species. Nevertheless, some other nonmarine Triassic ostracods, such as Lutkevichinella, do have ornamentation and may be more readily separated into reliable species of use to Triassic ostracod biostratigraphy. Nonmarine bivalves have a diverse record in Triassic strata but they have not been studied on a global basis. Like charophytes and ostracods, ecophenotypic variation is a problem for Triassic nonmarine bivalve taxonomy. A recent example of a nonmarine Triassic bivalve biostratigraphy stems from a revision of the extensive unionid bivalve

INTRODUCTION

record of the Upper Triassic Chinle Group in the western USA by Good (1998). He recognized two zones applicable within the Chinle basin (tetrapods define at least four zones in the Chinle), but the broader correlation of these zones is undemonstrated. A robust nonmarine biostratigraphy based on Triassic fishes does not exist and probably never will because nonmarine fishes are typically very limited in their distribution by particular lithofacies, so that their record is facies controlled and characterized by endemism. Regional biostratigraphy has been proposed for some nonmarine Triassic fish successions, such as in the Upper Triassic strata of North America: Chinle Group and Newark Supergroup. But, correlation within and between these successions is imprecise at best (Olsen et al. 1982; Huber et al. 1993; Milner et al. 2006).

A Triassic timescale The Triassic timescale presented here (Figs 4– 6) is based primarily on data in the chapters in this book. The numerical ages are from Mundil et al. 2010, and are only placed at the stage boundaries with precise numerical age control. Thus, no numerical ages are assigned to the bases of the Carnian, Norian or Rhaetian stages. The magnetostratigraphy is from Hounslow & Muttoni 2010, but has been simplified by eliminating less reliable magnetozones. The ammonoid zones follow Balini et al. (2010), the conodont zones are from Kozur (2003) and Orchard (2010), and the radiolarian zones for Japan (the most complete and broadly applicable Triassic zonation) are from O’Dogherty et al. (2010). Conchostracan zonation is that of Kozur & Weems (2010), the palynomorphs are the Germanic basin zonation of Kuerschner & Herngreen (2010), and the land –vertebrate faunachrons are from Lucas (2010b). Some imprecision and uncertainty exists in the correlation of some of the biostratigraphic zonations to each other. Thus, the ammonoid and conodont zonations are fairly precisely matched to each other, but the match to the radiolarian zonation is much less precise (see O’Dogherty et al. 2010). The conchostracan, palynomorph and tetrapod biozonations are readily matched to each other, but are less precisely correlated to the marine biostratigraphy. The absence of horizontal lines between some of the zonal indicators, especially in the radiolarian and the conchostracan zonations, reflects this imprecision. The Early Triassic (Fig. 4) and the Middle Triassic (Fig. 5) timescales are essentially scaled vertically to the radioisotopic ages. However, the Late Triassic (Fig. 6) is not. Here, I reject the ‘long Norian’ of Muttoni et al. (2004), which

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places the Norian base at c. 228 Ma. Instead, I follow arguments articulated by Kozur & Weems (2005, 2007, 2010) and Lucas (2010b) that the nonmarine biostratigraphy confirms a Norian base near the base of the Passaic Formation of the Newark Supergroup (eastern North America). This, coupled with an age of c. 219 Ma in Chinle Group (American SW) upper Carnian strata, places the Carnian –Norian boundary close to the 217 Ma estimated by Olsen & Kent (1999) from the Newark cyclostratigraphy. Nevertheless, such a conclusion creates a very long Tuvalian, at least 10 million years long, because of the c. 230 Ma age reported by Furin et al. (2006) from lower Tuvalian marine strata in Sicily. Clearly, there are problems reconciling the radioisotopic age data, magnetostratigraphy and biostratigraphy across the Carnian –Norian boundary. These problems, and similar issues, make any Triassic timescale, such as the one presented here, a work in progress that will continue to be modified and refined in the light of new data. I thank all the contributors to this volume for their perspicacity and patience. I also thank John Gregory, Norman Silberling and Lawrence Tanner for their reviews of this manuscript.

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INTRODUCTION K OZUR , H. & W EEMS , R. E. 2007. Upper Triassic conchostracan biostratigraphy of the continental rift basins of eastern North America: its importance for correlating Newark Supergroup events with the Germanic basin and the international geologic timescale. New Mexico Museum of Natural History and Science Bulletin, 41, 137–188. K OZUR , H. W. & W EEMS , R. E. 2010. The biostratigraphic importance of conchostracans in the continental Triassic of the northern hemisphere. In: L UCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 315– 417. K OZUR , H. W., K RAINER , K. & M O¨ STLER , H. 1996. Radiolarians and facies of the Middle Triassic Loibl Formation, South Alpine Karawanken Mountains (Carinthia, Austria). Geologisch– Pala¨ontologische Mitteilungen Innsbruck, Sonderband, 4, 195– 269. K RYSTYN , L., B HARGAVA , O. N. & R ICHOZ , S. 2007a. A candidate GSSP for the base of the Olenekian Stage: mud at Pin Valley; Himachal Pradesh (W. Himalaya), India. Albertiana, 35, 5– 29. K RYSTYN , L., B OQUEREL , H., K UERSCHNER , W., R ICHOZ , S. & G ALLET , Y. 2007b. Proposal for a candidate GSSP for the base of the Rhaetian Stage. New Mexico Museum of Natural History and Science Bulletin, 41, 189–199. K UMMEL , B. 1953. American Triassic coiled nautiloids. U.S. Geological Survey Professional Paper, 250, 1–104. K U¨ RSCHNER , W. M. & H ERNGREEN , W. 2010. Triassic palynology of central and northwestern Europe: a review of palynofloral diversity patterns and biostratigraphic subdivisions. In: L UCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 263– 283. L UCAS , S. G. 1990. Toward a vertebrate biochronology of the Triassic. Albertiana, 8, 36– 41. L UCAS , S. G. 1997. Upper Triassic Chinle Group, western United States: a nonmarine standard for Late Triassic time. In: D ICKINS , J. M. ET AL . (eds) Late Paleozoic and Early Mesozoic Circum-Pacific Events and Their Global Correlation. Cambridge University Press, Cambridge, 209– 228. L UCAS , S. G. 1998. Global Triassic tetrapod biostratigraphy and biochronology: Palaeogeography, Palaeoclimatology, Palaeoecology, 143, 347–384. L UCAS , S. G. 2006. Sanmiguelia from the Upper Triassic Chinle Group and its biostratigraphic significance. New Mexico Museum of Natural History and Science Bulletin, 37, 407–409. L UCAS , S. G. 2010a. The Triassic chronostratigraphic scale: history and status. In: L UCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 17–39. L UCAS , S. G. 2010b. The Triassic timescale based on nonmarine tetrapod biostratigraphy and biochronology. In: L UCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 447– 500. M C R OBERTS , C. A. 2001. Triassic bivalves and the initial marine Mesozoic revolution: a role for predators? Geology, 29, 359– 362. M C R OBERTS , C. A. 2010. Biochronology of Triassic bivalves. In: L UCAS , S. G. (ed.) The Triassic

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V ERMEIJ , G. J. 1977. The Mesozoic marine revolution: evidence from snails, predators and grazers. Paleobiology, 3, 245–258. V ON H ILLEBRANDT , A., K RYSTYN , L. & K UERSCHNER , W. M. (with contributions from B OWN , P. R., M C R OBERTS , C., R UHL , M., S IMMS , M., T OMASOVYCH , A. & U RLICHS , M.) 2007. A candidate GSSP for the base of the Jurassic in the Northern Calcareous Alps (Kuhjoch section, Karwendel Mountains, Tyrol, Austria. International Subcommission on Jurassic Stratigraphy Newsletter, 34, 2 –20. V UKS , V. J. 2000. Triassic foraminifers of the Crimea, Caucasus, Mangyshlak and Pamirs (biostratigraphy and correlation). Zentralblatt fu¨r Geologie und Pala¨ontologie Teil I, 11–12, 1353– 1365. V UKS , V. J. 2007. New data on the Late Triassic (Norian-Rhaetian) foraminiferans of the western Precaucasus (Russia). New Mexico Museum of Natural History and Science Bulletin, 41, 411–412. W ARRINGTON , G., A UDLEY -C HARLES , M. G. ET AL . 1980. A correlation of Triassic rocks in the British Isles. Geological Society Special Report, 13, 1– 78. Y IN , H. (ed.) 1996. The Palaeozoic–Mesozoic Boundary: Candidates of the Global Stratotype Section and Point of the Permian–Triassic Boundary. Wuhan, China University of Geoscience Press. Y IN , H., S WEET , W. C. ET AL . 1996. Recommendation of the Meishan section as global stratotype section and point for basal boundary of Triassic System. Newsletters on Stratigraphy, 34, 81–108. Y IN , H., Z HANG , K., T ONG , J., Y ANG , Z. & W U , S. 2001. The global stratotype section and point (GSSP) of the Permian–Triassic boundary. Episodes, 24, 102–114.

The Triassic chronostratigraphic scale: history and status SPENCER G. LUCAS New Mexico Museum of Natural History and Science, 1801 Mountain Road NW, Albuquerque, NM 87104 – 1375 USA (e-mail: [email protected]) Abstract: The Triassic chronostratigraphic scale is a hierarchy of three series, seven stages and 15 substages developed during nearly two centuries of research. The first geological studies of Triassic rocks began in Germany in the late 1700s and culminated in 1834 when Friedrich August von Alberti coined the term ‘Trias’ for the Bunten Sandsteins, Muschelkalk and Keuper, a thick succession of strata between the Zechstein and the Lias. Recognition of the Trias outside of Germany soon followed, and by the 1860s Austrian geologist Edmund von Mojsisovics began constructing a detailed Triassic chronostratigraphy based on ammonoid biostratigraphy. In 1895, Mojsisovics and his principal collaborators, Wilhelm Waagen and Carl Diener, published a Triassic timescale that contains most of the stage and substage names still used today. In 1934, Leonard Spath proposed a Triassic ammonoid-based biochronological timescale that differed little from that of Mojsisovics and his collaborators. In the 1960s, E. Timothy Tozer proposed a Triassic ammonoid-based timescale based on North American standards, and his timescale included proposal of four Lower Triassic stages (Griesbachian, Dienerian, Smithian and Spathian). The work of the Subcommission on Triassic Stratigraphy began in the 1970s and resulted in current recognition of seven Triassic stages in three series: Lower Triassic–Induan, Olenekian; Middle Triassic–Anisian, Ladinian; Upper Triassic –Carnian, Norian and Rhaetian. The 1990s saw the rise of Triassic conodont biostratigraphy so that four intervals that have agreed on Triassic GSSPs use conodont occurrences as defining features: bases of Induan, Olenekian, Anisian and Rhaetian. The bases of the Ladinian and Carnian are defined by ammonoid events. The base of the Norian remains undefined, but will most likely be defined by conodonts. Except for the Rhaetian, the Middle and Upper Triassic stages and substages have been fairly stable for decades, but there has been much less agreement on Lower Triassic chronostratigraphic subdivisions. Issues in the development of a Triassic chronostratigraphic scale include those of: stability and priority of nomenclature and concepts; disagreement over and changing taxonomy; the use of ammonoid v. conodont biostratigraphy; differences in the perceived significance of biotic events for chronostratigraphic classification; disagreements about the utility of relatively short stages; correlation problems between the Tethyan and Boreal realms (provinces); and competing standards from the Old and New worlds. Most of these issues have been resolved in the recognition of three Triassic series and seven stages. Further development of the Triassic chronostratigraphic scale needs to focus on definition and characterization of the 15 Triassic substages as these will provide a much more detailed basis for subdivision of Triassic time than do the seven stages.

The Mesozoic Era begins with the approximately 50– million-year-long Triassic Period, a major juncture in earth history when the vast Pangean supercontinent completed its assembly and began its fragmentation, and the global biota diversified and modernized after the end-Permian mass extinction, the most extensive biotic decimation of the Phanerozoic (e.g. Lucas 1999; Benton & Twitchett 2003; Lucas & Orchard 2004; Erwin 2006). The temporal ordering of geological and biotic events during Triassic time thus is critical to the interpretation of some unique and pivotal events in Earth history. This temporal ordering is mostly based on the Triassic chronostratigraphic scale: a relative scale of series, stages and substages that has been developed and refined for nearly two centuries. Here, I offer a brief history of that scale (see Zittel 1901 and Tozer 1984, for more in-depth treatments of some aspects

of this history). My goals in presenting this history of the Triassic chronostratigraphic scale are both to explain the origin and development of the currently used Triassic chronostratigraphic terminology and to present some analysis and evaluation of this process and of the current Triassic chronostratigraphic scale. In this article, GSSP ¼ ‘Global Stratotype Section and Point’ and STS ¼ IUGS Subcommission on Triassic Stratigraphy.

History Recognition of a distinctive interval in Earth history (originally identified as a distinct succession of stratified rocks) that corresponds to the current concept of Triassic began in Germany more than 200 years ago. This early work was culminated by Alberti’s

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 17– 39. DOI: 10.1144/SP334.2 0305-8719/10/$15.00 # The Geological Society of London 2010.

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(1834) monograph in which he coined the term Trias. The 200-year-long history of the development of a Triassic relative timescale (the standard global chronostratigraphic scale) has been discussed by various authors, most notably Zittel (1901), Silberling & Tozer (1968) and Tozer (1984). This history can be divided into six distinct phases: 1. The initial studies of the Triassic strata of Germany, culminated by Alberti’s (1834) recognition of the Trias ‘formation’. 2. Extension of the term Trias to marine rocks and fossils outside of its type section in Germany, particularly southward into the Alps of Austria –Italy. 3. Recognition of subdivisions of Triassic time based on ammonoids, primarily in the Alps and in what is now Pakistan. 4. The 1934 Triassic timescale of British Museum palaeontologist Leonard Spath, a purely biochronological sheme based on his perception of the temporal succession of Triassic ammonoids. 5. The New World timescale of Canadian palaeontologist E. Timothy Tozer, first published in the 1960s, developed in part in collaboration with American geologist Norman J. Silberling, that divided Triassic time based on ammonoid successions in the Canadian Arctic islands (especially Ellesmere and Axel Heiberg Islands) and the western American Cordillera, mostly in British Columbia and Nevada. 6. The timescale still being developed by the STS, which is part of the International Commission on Stratigraphy of the International Union of Geological Sciences.

Alberti names the Trias Rocks and fossils in central Europe now considered to be Triassic in age have been studied since the late 1700s and were particularly well studied in southwestern Germany (Fig. 1). Thus, in the 1760s, the German scholar George Christian Fu¨chsel (1722– 1773) made specific published reference to rocks that later became part of the type Triassic (he even used the term ‘Muschelkalk’). Two other German scholars also contributed much to early understanding of the type Triassic: Johann Gottlob Lehmann (1719–1767), who first used the term Bunter (‘variegated’) for the sandstones below the Muschelkalk, and Leopold von Buch (1774–1853), who in 1822 Fig. 1. The type of Alberti’s (1834) Trias is an approximately 1.1–km-thick section of strata exposed in southwestern Germany that German geologists had early divided into the mostly nonmarine, sandstone-dominated Buntsandstein (originally Bunter or Bunter Sandsteins),

Fig. 1. (Continued) a marine limestone, dolomite and shale interval, the Muschelkalk, and a variegated interval of mostly nonmarine claystones and sandstones, the Keuper. The generalized lithologic section here is after Aigner et al. (1999).

TRIASSIC CHRONOSTRATIGRAPHIC SCALE

christened the Keuper (a corruption of a term used by miners in the Coburg region of Germany) to apply to the claystones and sandstones above the Muschelkalk (Fig. 1). Thus, the German geologists of the late 1700s and early 1800s recognized that a thick succession of strata lay between the Zechstein (the ‘firm ground’ of German miners, who mined the ‘copper slate’ [Kupferschiefer] above it) and the marine strata that came to be called Lias or Jura. One of these geologists was Friedrich August von Alberti (1795–1878), a salt mining engineer and mine manager who had extensive knowledge of the geology of southwestern Germany (for a biography see Hansch 1998 and references cited therein). The rocks he came to call Trias in 1834 are extensively exposed in southwestern Germany (Fig. 1), and Alberti had a detailed, firsthand knowledge of them. Alberti’s work was set in a tradition alien to modern geologic paradigms and terminology (cf. Berry 1987). Thus, to Alberti a ‘formation’ was a widely distributed aggregation of rocks of diverse lithology; its modern equivalent would be a ‘System’. Nevertheless, in modern terms, Alberti’s work can be simply seen as his recognition that a thick succession of stratified rocks is present between the Zechstein (Permian, in later terms) and the Lias (base of the Jurassic), which he recognized as a distinct formation (System), the Trias (Triassic). Indeed, Alberti (1834) did not recognize the Trias based on its lithologic homogeneity. Thus, he made it clear that the Triassic strata included sandstone, gypsum, dolomite, limestone, clay and other kinds of rocks. Rather, it was the fossil content of the Trias that indicated a single ‘formation’. Thus, to Alberti, Triassic fossils were distinct from those of the Zechstein below and the Lias above. Furthermore, by Alberti’s (1834, p. 314 –319) tabulation of German Triassic fossils, many genera of plants, clams, gastropods, crinoids, fishes and reptiles were found in two, three or four of the Triassic stratal groups that he recognized, giving the Trias a distinct palaeontological integrity and identity. Alberti’s (1834) monograph title (Fig. 2) makes clear his purpose: Beitrag ¼ contribution, Verbindung ¼ unification (union) and Gebilde ¼ things (formations or structures). Part I (p. 17 –157) and Part II (p. 158 –298) are the bulk of Alberti’s monograph and describe the Bunter Sandstein, Muschelkalk and Keuper in southwestern Germany, providing details of outcrops, lithologies, fossils and other features – the two sections amount to a ‘regional geology’ of the Triassic of southwestern Germany. Part III (p. 299– 339) provides the conclusions or deductions (‘Folgerungen’) from the first two parts, culminated by a section (p. 323– 324) titled (in the table of contents, on p. xix) as

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Fig. 2. Title page of Alberti’s (1834) monograph in which the name Trias is proposed.

‘Bunter Sandstein, Muschelkalk und Keuper sind Eine Formation. – Trias’. Whoever more closely examines the preceding analysis and tabulates all the fossils of the three hitherto separate formations; whoever further examines the transition of the different forms one into the other, and, indeed, considers the entire structure of the mountains and the markedly different character of the fossils of the Zechstein from those of the Lias, will realize that the Bunter sandstone, Muschelkalk and Keuper are the result of a single period, their fossils, to use Elie de Beaumont’s words, are the thermometer of a geological period; that separation into three formations is not appropriate, and that it is more in accord with the concept of a formation to unite them into a single formation, which I shall provisionally name Trias (Alberti, 1834, p. 323–324; English translation modified from Berry 1987).

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The monograph concludes with a bibliography (p. 340 –366) of 269 titles. This bibliography and the text make it clear that Alberti recognized Triassic rocks in countries outside of Germany, including England, France, Austria, Italy, Poland, Russia, the United States and India. This broad geographic distribution of the ‘Trias formation’ served to insure the ready acceptance of the Triassic as part of the geological timescale.

Alpine marine Triassic Alberti’s type Triassic in southwestern Germany is a sandwich of dominantly nonmarine red beds (Buntsandstein and Keuper) with a restricted marine middle portion (Muschelkalk) (Fig. 1). Even in the nineteenth century, the broadscale correlation of nonmarine strata was recognized as more difficult than the correlation of marine fossil assemblages, which are often similar in very disparate regions. Thus, recognition of Muschelkalkequivalent marine strata, based largely on their content of ceratites (ammonoids), became key to extension of the Trias outside of Germany. This extension took place between about 1834 and 1867, and a central figure was the Austrian geologist Franz Ritter von Hauer (1822– 1899). Hauer provided the first descriptions of many ammonoid taxa that later became key Triassic index fossils (Tozer 1984). Indeed, Hauer (1850) rather quickly recognized and correlated marine strata in the Alps to the Germanic Triassic. Other workers, especially the German geologists Leopold von Buch, Karl Wilhelm von Gu¨mbel (1823–1898) and Ferdinand von Richthofen (1833–1905) also contributed much to the understanding of Triassic strata in the Alps (see Zittel 1901, p. 464 –483 for a detailed history). The Alps contain a relatively complete section of Triassic marine strata, so extension of the Triassic into the Alpine marine strata became central to further subdivision and correlation of Triassic time. Indeed, the work of Hauer and his contemporaries set the stage for making strata in the Alps standards for many of the subdivisions of Triassic time. Furthermore, during this time interval (1830s– 1860s), as Tozer (1984) documents, marine Triassic fossils were identified in far-flung localities across the globe, including Siberia, Japan, Peru and western North America. Thus, although Hauer and his contemporaries did not produce a Triassic timescale, they laid the foundation for the advances that immediately followed.

Mojsisovics The subdivision of Triassic time owes more to the Austrian geologist Edmund von Mojsisovics

(1839–1907) than to any other geologist. Mojsisovics 30-year career began in the 1860s (Rosenberg 1958; Tozer 1984). He worked for the Geological Survey (the Geologiches Reichsanstalt) of the Hapsburg monarchy, collecting and studying Triassic ammonoids throughout much of central and southern Europe, and publishing on ammonoids sent to him from locations as remote as the Olenek River in eastern Siberia (e.g. Mojsisovics 1869, 1874, 1882a, b, 1888, 1892, 1893, 1902). Mojsisovics provided detailed and lavishly illustrated monographic documentation of the Triassic ammonoids he studied (Fig. 3). Recognition of subdivisions of Triassic time based on ammonoids by Mojsisovics and his collaborators, particularly Carl Diener (1862–1928) and Wilhelm Heinrich Waagen (1841–1900), produced most of the stagelevel terminology of Triassic time still used today. This work was culminated by Mojsisovics et al. (1895), the singlemost important article written on the Triassic timescale. This article built largely on Mojsisovics’ and Diener’s work in Europe (primarily Austria, Italy and Bosnia), and Waagen’s work in the Salt Range of what is now Pakistan. It coined the names of most of the marine stages and sub-stages recognized today (Fig. 4). This timescale was refined subsequently, especially by the addition of Bittner’s (1892) Ladinian, but remained the basic Triassic timescale until at least the 1960s. Mojsisovics et al. (1895) began with an introduction (p. 1271 –1278) that credited Hauer and others who developed the initial understanding of the Alpine marine Triassic. Mojsisovics et al. (1895) then acknowledged the presence of Triassic formations in ‘Ostindien’ and Siberia, and also made it clear they were aware of Triassic ammonoid records in Timor, New Caledonia, New Zealand, Peru, the western USA, Canada, Japan and Spitzbergen, which indicates a remarkable breadth of knowledge of the Triassic ammonoid record (compare, for example, Kummel’s 1979 summary of the global record of Triassic ammonoids). Mojsisovics et al. (1895) noted that a critical point was to correlate the German Muschelkalk to part of the Alpine Triassic. They also pointed out the inadvisability of applying the terms Buntsandstein and Keuper as time terms for marine strata outside of Germany. Their goal was explicitly stated as establishment of a classification of the ‘pelagic’ Triassic based on cephalopod faunas, similar to that which had been established by Albert Oppel (1831– 1865) and Friedrich August von Quenstedt (1809–1889) for the Jurassic. Waagen and Diener received separate authorship for the Early Triassic part of the timescale (Mojsisovics et al. 1895, p. 1278–1296). They explained the origin of the names for time subdivisions, such as

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Fig. 3. Plate 20 of Mojsisovics (1893) monograph on the ammonoids of the Triassic ‘Mediterraneanprovinz’. Figure 1 is ‘Trachyceras longobardicus’, and Figure 2 is ‘Trachyceras pseudo-Archelaus’.

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Fig. 4. The Triassic timescale of Mojsisovics et al. (1895).

Skythisch (Fig. 4), and then listed the succession of ammonoid assemblages in the Salt Range of Pakistan that characterized each subdivision. They also discussed the Dinarian, treating it as Early Triassic, though it includes the Anisian. Waagen and Diener ended their discussion by reviewing the distribution of Early Triassic ammonoid faunas. Mojsisovics (p. 1296–1302) covered the ‘Obere Trias’ by presenting an updated review of Mojsisovics (1893). In the 1895 paper, he also coined the terms Cordevol, Jul, Tuval, Lac, Alaun and Sevat, subdivisions of Carnian and Norian time, most of which are still used today as substages. He then reviewed the distribution of Upper Triassic ammonoid faunas. The broad influence of the timescale of Mojsisovics et al. (1895) is well seen in the work of American palaeontologists Alpheus Hyatt (1838–1902) and, particularly James Perrin Smith

(1864–1931). They monographed the North American Triassic ammonoid assemblages that Smith collected (Hyatt & Smith 1905; Smith 1914, 1927, 1932), fitting them to the timescale developed by the Austrians, and filling in some of the gaps (notably in the upper Lower Triassic) in the ammonoid successions.

Spath’s biochronological timescale British Museum palaeontologist Leonard F. Spath (1888–1957) produced a comprehensive Triassic timescale in the context of his monographic description of the British Museum’s collection of Triassic ammonoids (Spath 1934, 1951). Spath’s (1934) scheme (Fig. 5) was a purely biochronologic one, based on the concepts of British palaeontologist Sydney S. Buckman (1860–1929), who divided Jurassic time based on his perceptions of ammonoid

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Fig. 5. Spath’s (1934) Triassic timescale.

evolution and succession, to some extent (especially in his later work) without actual regard for stratigraphic succession (see Arkell 1933, p. 19 –37 for a review of Buckman’s concepts and work). Thus, Spath defined a succession of ages named after ammonoids (for example, Tropitan, named after the widespread Late Triassic ammonoid genus Tropites) that were seen of as intervals of time (biochrons) that correspond to the temporal ranges of ammonoid taxa. Nevertheless, stratigraphic position did play an important role in Spath’s scheme, as Spath (1934, tables 3– 5) listed the ammonoid zones he assigned to each age and referred to specific strata and localities as ‘equivalents’. Indeed, his Triassic timescale built heavily on that of Mojsisovics et al. (1895) and synthesized their work with that of Smith and others. Although Spath did no stratigraphic work on Triassic ammonoids, he was well aware of the limitations of their record. He thus urged caution by stating that ‘until there is considerably more information about detailed sequences in the Trias, and especially its lower beds, the scheme here put forward cannot be considered to be more than an approximation to a more detailed chronology’ (Spath 1934, p. 22). Nevertheless, Spath’s Triassic

timescale had relatively limited influence. Thus, in a rare example, it was incorporated into the work of Kiparisova & Popov (1956, p. 844), who made clear the correspondence between their newly named Induan and Olenekian stages and Spath’s ammonoid biochrons.

Tozer’s American timescale Beginning in the 1960s, Canadian palaeontologist E. Timothy Tozer (1928-), in part collaborating with American geologist Norman J. Silberling (1928-), assembled a Triassic timescale based on North American ammonoid zones (e.g. Tozer 1965, 1967, 1971, 1974, 1978, 1984; Silberling & Tozer 1968). This work built heavily on the pioneeering efforts of Frank McLearn (1885–1964) in British Columbia, and Siemon Muller (1900– 1970) in Nevada, who had deciphered much of the ammonoid succession in the American Cordillera. Tozer added Lower Triassic ammonoid successions from the Queen Elizabeth Islands of Arctic Canada, especially Ellesmere and Axel Heiberg Islands, to the American record. The Silberling & Tozer (1968) timescale captured the early phase of this work, and Tozer continued to refine it, culminated

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by publication of his 1994 monograph on Canadian Triassic ammonoids (Fig. 6). Key components of Tozer’s Triassic timescale were that it: (1) used most of the Middle and Upper Triassic stage names of Mojsisovics et al. (1895), with the addition of Bittner’s Ladinian; (2) but, it defined Triassic stage boundaries based on North American ammonoid localities; (3) it rejected

the Rhaetian as a distinctive stage; and (4) new stage names were coined to create a fourfold division of the Lower Triassic. Tozer’s timescale, especially the fourfold subdivision of the Lower Triassic, found wide acceptance in the English language literature on the subdivision of Triassic time, though few abandoned the Rhaetian (e.g. Kummel 1979; Harland et al. 1982, 1990).

Fig. 6. Tozer’s (1994) Triassic timescale showing the distribution of ammonoids (black dots) in strata in the Canadian Arctic Islands and Cordillera.

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The STS timescale Conceived in 1968, and beginning its meetings in the 1970s (Tozer 1985), the Subcommission on Triassic Stratigraphy (STS), as part of the International Commission on Stratigraphy, was primarily designed to establish a global Triassic timescale based on GSSP definitions of the bases of the Triassic stages (e.g. Gaetani 1996). Tozer (1984), in his history of the Triassic timescale, had labelled his North American timescale as one of ‘emancipation’. He ended his work with a chapter on ‘worldwide reconciliation’, arguing that the North American standard could be reconciled with that in the Alps and elsewhere in the Tethys. However, instead of an immediate reconciliation, the STS began its published discussion (in the STS journal Albertiana) with a lively debate over the Tozer timescale – particularly over whether or not to recognize the Rhaetian as a separate stage, which Tozer had abandoned as a ‘substage’ of the Norian (e.g. Tozer 1986a, 1988, 1990; Ager 1988, 1990; Dagys 1988a; Silberling & Nichols 1988; Krystyn 1990). After initial acceptance in 1984 of most aspects of the Tozer timescale, in 1991, the STS agreed on a stage nomenclature of the Triassic in which key aspects of Tozer’s timescale were rejected (Fig. 7). Indeed, the STS has

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even recently arrived at a formal definition of the Rhaetian Stage (Fig. 8). The 1990s witnessed the rise of conodont biostratigraphy and the professional retirement of the proponents of the North American ammonoidbased timescale. Thus, instead of using ammonoiddefined stage boundaries – which had a century long history – at least some of the Triassic stage boundaries were redefined using conodonts. Indeed, the first GSSP approved by the STS, the base of the Induan (base of the Triassic) was defined by the LO of the conodont Hindeodus parvus at the Meishan section in southern China (Yin et al. 1996, 2001; also see Yin 1996). Subsequent GSSP proposals based on conodont-defined boundaries have been for the bases of the Olenekian, Anisian, Carnian, Norian and Rhaetian (e.g. Shevyrev 2003, 2006; Krystyn et al. 2007a, b; Orchard et al. 2007). Only the bases of the Ladinian and Carnian are defined by ammonoid datums (Balini 2003; Brack et al. 2005; Gaetani 2009).

Subdivisions of the Triassic The Triassic has always been subdivided into three series: Lower, Middle and Upper. No other serieslevel subdivisions of the Triassic have been

Fig. 7. STS Triassic stage nomenclature in 1985 and in 1992 (see Visscher 1985, 1992).

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Fig. 8. Possible GSSP levels (1, 2, 3) and proxies for the base of the Rhaetian Stage at the Steinbergkogel, Austria, section (from Krystyn et al. 2007b).

proposed. Some workers (e.g. Harland et al. 1982) have used the term ‘Scythian’ to refer to the Lower Triassic series, but this has not been widely used. Thus, no formal terms are used for the Triassic series. Marine sections critical to definition of Triassic chronostratigraphic subdivisions are those in the Alps of central and southern Europe, the Aegean islands, Transcaucasia, India –Pakistan, southern

China, eastern Siberia, the Cordillera of western North America and the Queen Elizabeth Islands of the Canadian Arctic (Fig. 9).

Lower Triassic The ‘type’ Lower Triassic is the German Buntsandstein, a complex lithosome that consists mostly of

Fig. 9. The Triassic world with regions labelled that have been critical to the development of the Triassic chronostratigraphic timescale (artwork by Matt Celeskey).

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nonmarine red-bed siliciclastic strata (Fig. 1). Biostratigraphic and magnetostratigraphic data now indicate that the Buntsandstein ranges in age from latest Permian to earliest Middle Triassic (early Anisian) (e.g. Kozur & Bachmann 2005). However, given that the type Lower Triassic is mostly nonmarine strata, it has not been used as a standard to subdivide Early Triassic time. There has been more disagreement about the stage-level subdivisions of the Lower Triassic than there has been about subdividing the Middle and the Upper Triassic (e.g. Tozer 1967, 1978, 1984; Guex 1978; Zakharov 1987; Kozur 1992). Indeed, as Tozer (1978) noted, about 20 chronostratigraphic terms have been proposed to refer to all or part of the Lower Triassic (Fig. 10). The STS now divides the Lower Triassic into two stages, the Induan and the Olenekian. However, other schemes of subdividing the Lower Triassic exist, particularly in North America, where at least three Lower Triassic stages have been recognized (Fig. 10). Most notable are the stages proposed by Kiparisova & Popov (1956), who divided the Lower Triassic into the Induan and the Olenekian stages. Tozer (1967), based on studies in the Canadian Arctic (especially Ellesmere and Axel Heiberg islands), divided the Lower Triassic into four stages based on ammonoid biostratigraphy. Here, I briefly review proposed subdivisions of the Lower Triassic that have current usage. Other proposed subdivisions, such as Werfenian (Mojsisovics 1882a), Gangetian, Gandarian, Jakutian and Brahmanian (Mojsisovics et al. 1895) or Ellesmerian (Kozur 1972), have never gained acceptance and wide usage, so they are not discussed in detail here.

Scythian Mojsisovics et al. (1895) proposed Scythian as a subdivision of the Triassic, deriving the name from Skythia (an old Greek term that referred to what is now southern Russia, the eastern Ukraine and Kazakstan), where strata were known to yield Early Triassic ammonoids such as Tirolites. The apparent type locality is Mount Bogdo, east of the Volga River. However, beginning with Waagen and Diener (in Mojsisovics et al. 1895), the entire Lower Triassic was assigned to the Scythian, used either as a stage (Kummel 1957) or as a series (e.g. Harland et al. 1982, 1990). Today, the term Scythian is little used as it is merely a synonym of Lower Triassic.

Induan Though working in Siberia, Kiparisova & Popov (1956, p. 844) chose to call the oldest stage of the Triassic the Induan, using the Otoceras woodwardi

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ammonite zone of the Himalayas as the base of the stage, and the Ceratite Sandstone of the Salt Range of western Pakistan as the top of the stage. Kiparisova & Popov (1964) later excluded the Ceratite Sandstone from the Induan. The name Induan is for the Indus River in India–Pakistan. In 1991, the STS decided to use the term Induan as the name of the oldest Triassic stage (Visscher 1992), so that the Induan is now used on the current ICS timescale as the oldest Triassic stage (Ogg 2004a, b; Ogg et al. 2008). Its base has been formally defined as the LO (lowest occurrence) of the conodont Hindeodus parvus at the Meishan section in southern China (Yin et al. 1996, 2001). There is also broad agreement that the Induan is equivalent to the upper Griesbachian and the Dienerian of the Tozer timescale (e.g. Ogg 2004a) (Fig. 10).

Olenekian Kiparisova & Popov (1956, p. 844) introduced the term Olenekian to refer to the younger stage of the Lower Triassic. The name comes from the Olenek beds on the Olenek River in northeastern Siberia. Kiparisova & Popov (1956) considered it equivalent to the ‘Upper Eo-Trias’ of Spath (1934), but later (Kiparisova & Popov 1964) changed the stage to include older strata. Vavilov & Lozovsky (1970) restricted the term Olenekian to be equivalent to the Spathian, but it has generally been considered equivalent to the Smithian and Spathian of Tozer (e.g. Shevyrev 2002; Ogg 2004a) (Fig. 10). In 1991, the STS decided to adopt the Olenekian as the younger stage of the Lower Triassic (Visscher 1992). Recent lengthy debate about a GSSP definition of the Olenekian base may be resolved by the choice to use the LO of the conodont Neospathognaothus waageni as the defining criterion, with a GSSP located in the Mud section in India (Krystyn et al. 2007a).

Griesbachian Tozer (1965, 1967) introduced the term Griesbachian Stage for four ammonoid zones (Otoceras concavum, O. boreale, Ophiceras commune and Proptychites striatus zones). The name is for Griesbach Creek on Axel Heiberg Island in the Canadian Arctic. Subdivision into two substages, Gangetian (lower Griesbachian ¼ O. concavum and O. boreale zones) and Ellesmerian (upper Griesbachian ¼ Ophiceras commune and P. striatus zones) has never gained wide usage. However, among ammonoid workers, the term Griesbachian still is in wide usage. Griesbach (1880) first proposed the LO of the Tethyan ammonoid Otoceras woodwardi as the base of the Triassic, and Tozer (1967, 1986b, 1994)

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Fig. 10. Various proposed subdivisions of the Lower Triassic (modified from Tozer 1974).

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equated the lowest occurrence of the boreal ammonoid O. concavum with the base of the Triassic. However, the lowest occurrence of O. concavum was subsequently shown to be older than the lowest occurrence of O. woodwardi (Krystyn & Orchard 1996). Furthermore, the conodont-based definition of the base of the Triassic is the lowest occurrence of Hindeodus parvus, which is in the lower part of the O. boreale zone. Thus, the lower part of the Griesbachian is regarded as Permian (Orchard & Tozer 1997a, b).

Dienerian Tozer (1965) introduced the Dienerian Stage for two ammonoid zones, the Proptychites candus and Vavilovites sverdrupi zones. The name is for Diener Creek on Ellesmere Island, Canada. Subsequent work by Guex (1978) induced Tozer (1981) to reduce the Dienerian to a substage of Guex’s Nammalian Stage, but the Nammalian has never gained wide usage, whereas the Dienerian remains a recognizable subdivision of the Lower Triassic (see below).

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Therefore, Guex (1978, p. 115) proposed the Nammalian Stage to encompass the Dienerian and Smithian stages. The name is for Nammal Gorge in the Salt Range of Pakistan, where Nammalian strata are the Lower Ceratite Limestone, Ceratite Marls, Ceratite Sandstone and the Upper Ceratite Limestone. Tozer (1981, 1984) initially accepted this modification of his fourfold stage division of the Lower Triassic, but by the 1990s Nammalian has largely disappeared from usage because the Dienerian and Smithian do remain recognizable in many Lower Triassic sections (e.g. Tozer 1994; Brayard & Bucher 2008; Bru¨hwiler et al. 2008; Shigeta et al. 2009). Incidentally, Rostovcev & Dagys (1984) argued that the name Nammalian was preoccupied by Gee’s (1935) use of the term Nammal to refer to Eocene strata in the Salt Range. Indeed, the Nammal Formation is a formal lithostratigraphic name still used to refer to an Eocene lithosome in the Salt Range and Trans-Indus ranges of Pakistan (Shah 1977).

Smithian

Middle Triassic

Tozer (1965, 1967) named the Smithian for two ammonoid zones in the Blind Fiord Formation on Smith Creek, Ellesmere Island: the Euflemingites romunderi and the Wasatchites tardus zones. After Guex (1978), Tozer (1981) regarded the Smithian as a substage of the Nammalian Stage. However, the Smithian remains a useful and widely recognized stage in ammonoid biostratigraphy (e.g. Jenks 2007; Lucas et al. 2007; Brayard & Bucher 2008; Bru¨hwiler et al. 2008; Shigeta et al. 2009).

The STS voted in 1984 and in 1991 to use two stages to divide the Middle Triassic: the Anisian and Ladinian (Fig. 7). Unlike the Lower Triassic, this part of the Triassic chronostratigraphic scale is fairly noncontroversial – only two stages have been advocated for the Middle Triassic and few competing (overlapping) stage terms have been proposed. The Anisian and Ladinian stages have also been subdivided into substages based on ammonoid biostratigraphy (Fig. 11).

Spathian

Anisian

Tozer (1965) named the Spathian for Spath Creek on Ellesmere Island. It originally encompassed two ammonoid zones in the Blaa Mountain Formation: the ‘Olenekites’ pilicatus and the Keyserlingites subrobustus zones. Silberling & Tozer (1968) recognized ‘beds’ containing Prohungarites and Subcolumbites as distinct zonal intermediates between the two zones. The Spathian is well recognizable in North America and elsewhere (Fig. 10), and recent work (e.g. Guex et al. 2005a, b) is revealing a further diversity and succession of Spathian ammonoid taxa.

Waagen and Diener (in Mojsisovics et al. 1895) introduced the term Anisian for fossiliferous limestones near Gross Reifling along the Enns (Latin name ‘Anisus’) River in Austria. The term was intended to replace the term ‘Alpine Muschelkalk’, which Mojsisovics had previously used to refer to two ammonoid zones, the ‘Ceratites’ binodosus and the ‘C.’ trinodosus zones. The Anisian was originally composed of two substages – Balatonian (¼ Balatonites balatonicus ammonoid zone) and the Bosnian (¼ Paraceratites trinodosus ammonoid zone). These terms have long been abandoned in favor of a fourfold subdivision of the Anisian into Aegean (lower), Bithynian and Pelsonian (middle) and Illyrian (upper) (Fig. 11). Arthaber (1905) redefined the Anisian of Waagen and Diener to include older strata and, indeed, Pia (1930) considered the type Anisian to only encompass the upper Anisian, and identified as lower Anisian the ‘Hydasp’, a term Mojsisovics

Nammalian Guex (1978) argued that although the Griesbachian – Dienerian boundary is readily recognized by ammonoid biotic events (notably the appearance of abundant meekoceratines), the Dienerian-Smithian boundary is less distinctive (also see Tozer 1974).

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(Steiermark, Austria), which was the stratotype Anisian of Waagen and Diener (in Mojsisovics et al. 1895). This section includes the ‘Binodosus –Fauna’ of the Tiefengraben locality and the ‘Trinodosus –Fauna’ from nearby Gamsstein, as well as the Rahnbauerkogel locality. Assereto (1974, p. 36) introduced a third (and oldest) subdivision of the Anisian, the Aegean substage, for the Paracrochordiceras beds on Chios Island in the Greek Aegean Sea. Assereto (1974, p. 35) also introduced the Bithynian as a substage of the Anisian older than the Pelsonian, basing it on two ammonoid zones (Nicomedites osmani and Anagymnotoceras ismindicum) on the Turkish Kokaeli Peninsula (‘Bithynia’). However, not all workers (e.g. Tozer 1981, 1984) regard the Bithynian as distinct from the Aegean. Either threefold or fourfold subdivisions of the Anisian are widely recognized (e.g. Mietto & Manfrin 1995; Monnet & Bucher 2005a, b).

Ladinian

Fig. 11. The current Triassic chronostratigraphic scale. Note that the Griesbachian substage as originally defined extends downward into the Permian, and that no substages of the Rhaetian have been proposed.

et al. (1895) used for a stage between the Scythian and the Anisian (Fig. 4). However, Hydasp has generally not been used to refer to the lower Anisian, neither has the term Kularian, introduced by Archipov et al. (1971) to refer to the lower and middle Anisian. Pia (1930) divided the upper Anisian into the Pelsonian and Illyrian, regarding the strata with ‘Ceratites’ binodosus zone ammonoids (Mojsisovics 1882a) at Cimego, Dont and Monte Cucco in the southern Alps as the type strata of the Pelsonian. He considered the Prezzo Limestone of Lombardy, which contains ammonoids of the ‘C.’ trinodosus Zone, as the type strata of the Illyrian (also see Assereto & Casati 1965). In effect, Pelsonian is the Latin name for the Balatonian, and Illyrian is the Latin name for the Bosnian (Pia 1930; Assereto 1974; Tozer 1984). Summesberger & Wagner (1972) provided a detailed description of the section at Grossreifling

When Mojsisovics (1869) named the Carnian and Norian, he considered the Norian to be older than the Carnian. However, in 1892 he reversed that order, and introduced the term Juvavian to refer to the pre-Carnian stage (Juvavian replaced his original Norian). At the same time, having established the proper sequence of Carnian and Norian strata in the Hallstatt area of Austria, Bittner (1892) introduced the name Ladinian for the stage immediately before the Carnian, to refer to the Buchenstein and Wengen beds (and possibly the St. Cassian Beds) of the Italian Dolomites. He named it for the Ladin people, who inhabit that region. Bittner’s Ladinian was then explicitly applied to strata that Mojsisovics originally had equated with his type Norian, resulting in an acrimonoius argument between Mojsisovics and Bittner (Zittel 1901; Tozer 1984). This issue was easily resolved more than a century ago by recognizing three temporally successive stages – Ladinian, Carnian and Norian – and abandoning the term Juvavian. The Ladinian is typically divided into two substages: Fassanian (lower) and Longobardian (upper). Mojsisovics (in Mojsisovics et al. 1895, p. 1297) introduced the term Fassanian (his ‘unternorisch’ [Fig. 4], which is equivalent to Bittner’s lower Ladinian) for the Val di Fassa in northern Italy where the Buchenstein Beds and Marmolada Limestone yielded ammonoids characteristic of the substage. Mojsisovics (in Mojsisovics et al. 1895, p. 1298) also introduced the Longobardian (his ‘obernorsich’ [Fig. 4], which is equivalent to Bittner’s upper Ladinian) for the Wengen Beds of the Dolomites of northern Italy. The name was derived from the Langobard people of the region, which

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has prompted the suggestion that the name should be properly spelled ‘Langobardian’ (Pia 1930), although this has never been followed. In Europe, the Fassanian and Longobardian substages are frequently used and are equivalent to the terms lower and upper Ladinian used by many workers in North America (Fig. 11). The base of the Ladinian is one of the few completely ratified GSSPs of the Triassic timescale. It is the lowest occurrence of the ammonoid Eoprotachyceras curionii in the Buchenstein Formation near Bagolino in northern Italy (Brack et al. 2005).

Upper Triassic The Upper Triassic is now divided into three stages: Carnian, Norian and Rhaetian. However, the status of the Rhaetian has been contentious: at times the stage has been recommended for abandonment and at other times it was assigned to the Jurassic System. At present, the STS recognizes three stages – Carnian, Norian and Rhaetian – as the Upper Triassic (Figs 7 & 11).

Carnian Mojsisovics (1869, p. 127) introduced the term Carnian Stage for ammonoid-bearing strata in the Austrian state of Ka¨rnten (Carinthia). He erroneously regarded it as younger than the Norian (see above). Mojsisovics (1874) assigned three ammonoid zones to the Carnian: Tropites subbullatus, Trachyceras aonoides and Trachyceras aon zones. Later, Mojsisovics (in Mojsisovics et al. 1895) divided it into three substages: Cordevolic (¼ Aon zone), Julic (¼ Aonoides Zone) and Tuvalic (¼ Subbullatus Zone). Tozer (1984) regarded the type locality of the Carnian as vague, as it was stated to refer to the Trachyceras and Tropites beds of the Hallstatt Limestone, but also included localities at Raibl, Bleiberg and St. Cassian. Lieberman (1980) proposed the Raibl section as the stratotype of the stage. Tozer (1984) and some others (see Fig. 7) have spelled the name ‘Karnian’, but this spelling has not been widely adopted. The Carnian Stage is typically divided into two substages named by Mojsisovics (in Mojsisovics et al. 1895): Julian (lower) and Tuvalian (upper). However, Mojsisovics (in Mojsisovics et al. 1895) also recognized a third Carnian substage, the Cordevolian, still used by some workers. Based on the St. Cassian Beds, Cordevolian derives its name from the Cordevol people who lived in the type area in northern Italy (Mojsisovics et al. 1895, p. 1298). Krystyn (1978) discussed the original definition of the Cordevolian and argued that it essentially referred to the same time interval as the Julian (also see Tozer 1967, 1974).

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The Julian was based on the Raibl Formation in the Julian Alps (southern Alps) by Mojsisovics (in Mojsisovics et al. 1895, p. 1298), and has come to be viewed by most workers as the lower Carnian (cf. Krystyn 1980; Tozer 1984, 1994). Mojsisovics (in Mojsisovics et al. 1895, p. 298) took the name Tuvalian from the Tuval Mountains between Hallein and Berchtesgaden (Bavaria –Austria), which was the Roman name for the area between Hallein and Berchtesgarden in Austria. He based it on the Tropites subbullatus ammonoid zone. Krystyn & Schlager (1971) suggested using the section at Feuerkogel near Aussee, Austria as the Tuvalian stratotype as well as the place to define the base of the Norian, in large part because the original ammonoids of Mojsisovic’s stratotype Tuvalian came from syntectonic fissure fills at Rappolstein. The term ‘Tuvalian’ has come to be used by most workers to refer to the entire upper Carnian (e.g. Krystyn & Schlager 1971; Tozer 1984, 1994). A GSSP for the base of the Carnian stage has recently been agreed on (Gaetani 2009). It is the LO of the ammonoid Daxatina canadensis at the Parti di Stuores/Stuores Wiesen section in northern Italy (Mietto et al. 2007a, b).

Norian Mojsisovics (1869, p. 127) named this stage for the Roman province of Noria, which was south of the Danube and included what is now the area of Hallstatt, Austria. He based the stage on the Hallstatt Limestone of the Salzkammergut in Austria, strata containing ‘Ammonites’ (Pinacoceras) metternichi. As well recounted by Tozer (1984), Mojsisovics originally thought the Norian was between the ‘Alpine Muschelkalk’ and the Carnian. When that mistake was discovered, Mojsisovics (1892) moved the term Norian to refer to pre-Carnian Hallstatt strata and named the Juvavian Stage, which is now regarded as synonymous with the Norian. As noted above, this caused an acrinomius debate with Bittner (1892), who argued to retain Norian as originally defined and proposed Ladinian to refer to the time interval before the Carnian (Zittel 1901, p. 494–497; Tozer 1984). Adding further to the confusion, Mojsisovics also provided no type section for the Juvavian, but instead referred to a succession of ammonoid zones (Mojsisovics 1902), a succession critiqued by Kittl (1903) and Diener (1921, 1926). The stratotype of the Norian has been considered to be the Bicrenatus Lager at Sommeraukogel, Hallstatt (Zapfe 1971; Krystyn & Schlager 1971; Krystyn et al. 1971). The Norian is generally divided into three substages: Lacian (early), Alaunian (middle) and Sevatian (upper). Mojsisovics (in Mojsisovics et al. 1895, p. 1298) used the term ‘Lacian’ to refer to the ‘lower

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Juvavian’. He took the name from the Roman name Lacia, which referred to the Salzkammergut area in Austria, and based it on the Cladiscites ruber and Sagenites giebeli ammonoid zones of the Hallstatt Limestone. As Tozer (1974) stressed, technically the Lacian was based on upper Norian ammonoids, so it is not a designation for the lower Norian as now recognized. However, this technicality has been largely ignored, and Lacian is frequently used to refer to the lower Norian substage. Mojsisovics (in Mojsisovics et al. 1895, p. 1298) named the Alaunian substage for the Alauns, a people who lived around the Hallein, Austria area during Roman times. He based it on what is now the Cyrtopleuirites bicrenatus ammonoid zone, and it is well accepted as the name of the middle Norian substage (Fig. 11). Mojsisovics (1895, p. 1298) named the Sevatian substage for a Celtic people who lived between the Inn and Enns Rivers in Austria. It was based on the Pinacoceras metternichi and Sirenites argonautae ammonoid zones in the Hallstatt area. The term is used by many workers to refer to the upper Norian, though Tozer (1984), who did not recognize the Rhaetian, did not use it. Problems with the Sevatian have largely been associated with defining a Rhaetian base. The now agreed-on GSSP assigns more than one ammonoid zone to the Rhaetian, so this ‘long Rhaetian’ encompasses part of the upper Sevatian of some previous usages (Krystyn et al. 2007b).

Rhaetian Gu¨mbel (1859, 1861, p. 116) used the term ‘Rha¨tische Gebilde’ to refer to the uppermost Triassic strata (Ko¨ssen beds) in the Bavarian Alps. The name was either for the Roman province of Rhaetium or the ratische Alpen. No type locality was specified, but Gu¨mbel did refer to the ‘Schichten der Rhaetavicula contorta’ (beds with the bivalve R. contorta). Thus, to Mojsisovics et al. (1895), the Rhaetian was the ‘Zone der Avicula contorta’ (Fig. 4). Lengthy debate about the Rhaetian (e.g. Pearson 1970; Ager 1987; also see above) has focused on three issues: (1) whether or not the stage should be assigned to the Jurassic; (2) whether or not the stage should be recognized or just subsumed into the Norian as advocated by Tozer (1967) and Silberling & Tozer (1968) and adopted by Zapfe (1974), Palmer (1983) and some others; and (3) how to define the Rhaetian base. Thus, for example, Slavin (1961, 1963) proposed that the Rhaetian beds of northwestern Europe, which lack ammonoids, should be placed in the Jurassic as a new Bavarian Stage, a proposal that met no acceptance. The STS recognizes a distinct Rhaetian, which is the youngest Triassic stage (Figs 7 & 11). The GSSP

candidate is the Steinbergkogel section near Hallstatt, Austria (Krystyn et al. 2007b). The recently agreed on redefinition of the Rhaetian is of a ‘long’ Rhaetian that includes two ammonoid zones with its base very close to or at the base of the Paracochloceras suessi zone (Fig. 8). However, the favoured definition of the Rhaetian base is the LO of the conodont Misikella posthersteini.

Other Triassic chronostratigraphic scales Current stratigraphic practice seeks to recognize a single global stage for each interval of time, and each series and system base corresponds to the base of a stage. Furthermore, the definition of stages is now based on the GSSP concept and the practice of integrated stratigraphy that applies multiple data sets to the definition of chronostratigraphic units (e.g. Salvador 1994; Remane et al. 1996; Walsh et al. 2004). However, the provinciality of fossil taxa compounded by limitations of facies distributions (rarely is any taxon or facies global in extent) have prevented universal recognition and use of a single chronostratigraphic terminology. Indeed, there remains great value in provincial stages, which Cope (1996) has aptly called the ‘secondary standard’ in stratigraphy. The Triassic has a variety of secondary standards, and some of the disagreement about a single stage nomenclature of the Triassic System reflects this. Furthermore, there are also provincial Triassic chronostratigraphies, such as that for New Zealand (e.g. Carter 1974; Campbell 1985; Cooper 2004). Here, I do not review these provincial scales, but note that their regional utility will guarantee their continued use.

Discussion This review of the historical development of the Triassic chronostratigraphic scale reveals some issues of stratigraphic philosophy and procedure that have been central to developing a Triassic timescale, and will continue to be factors in its future refinement. Here, I briefly comment on these issues, offering my own perspective on how the Triassic chronostratigraphic scale can be further refined.

Stability and priority There are no strict rules of priority in chronostratigraphic nomenclature, as there are in zoological nomenclature. However, most stratigraphers do attempt to employ the oldest term proposed for a time interval, redefining those terms as needed in the light of current data. This explains the continuity

TRIASSIC CHRONOSTRATIGRAPHIC SCALE

of usage of the Middle and Upper Triassic stage names since the work of Mojsisovics et al. (1895) and Bittner (1902). However, although Induan and Olenekian of Kiparisova & Popov (1956) are relatively old names for Lower Triassic stages, the Brahmanian of Waagen and Diener (in Mojsisovics et al. 1895) is a much older name and could have been used instead of Induan, as Kozur (1973, 1975) suggested (Fig. 10). The fact that the name with priority, Brahmanian, has been little used, using often-employed Induan is a step towards stability in the chronostratigraphic nomenclature. Pointing out the defects in the original definitions of the Triassic stages (Tozer 1984 does an excellent job of this) is historically interesting, but in no way detracts from the utility of terms that have been in use for more than a century, as long as those terms are precisely defined. It seems that after about a quarter century of study and debate by the STS, the seven Triassic stage names they endorse merit continued use. I realize that none of these stages is perfect: all have their defects. Furthermore, I do not believe that current definitions of their bases should be immutable: they, too, will be subject to further scientific testing and, if deemed advisable, redefinition. However, in the interest of stability, the seven Triassic stage names are long used and widely published terms that should continue to be used. Triassic chronostratigraphic research now needs to focus on definition of the boundaries of the Triassic substages, as these will provide an even more detailed basis for the subdivision of Triassic time than do the stages.

Taxonomy Arguments over biostratigraphy commonly reduce to arguments over taxonomy. This is because biostratigraphic correlation relies on what I call the ‘index taxon hypothesis’ – taxon A is the same age wherever its fossils are present. This hypothesis is usually tested by comparing the distribution of taxon A to that of other taxa, or by using another method of correlation (e.g. radioisotopic dates, magnetostratigraphy) to try to demonstrate diachroneity in the distribution of taxon A. However, palaeontological taxonomists sometimes decide that taxon A is no longer a single taxon, so they split it into two or more taxa, thus reducing (or eliminating) its utility as an index taxon. These kinds of taxonomic changes destabilize existing biostratigraphy and exasperate many geologists who are forced to contend with resulting changes in the chronostratigraphy (see Shaw 1969 for an excellent expression of this). However, taxonomic changes will always be with us, as new data, techniques and epistemologies are employed in palaeontological taxonomy. To my mind, in Triassic

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biostratigraphy this favors the ammonoids, whose long history of taxonomic study has tested many taxonomic concepts. Triassic conodonts, in contrast, have a much younger and less tested taxonomy. I predict that further testing of their taxonomy will produce destabilizing changes that will modify existing Triassic conodont biostratigraphy, and this will impact conodont-based chronostratigraphic definitions. Indeed, as an example, recent taxonomic revisions have altered the taxonomy of the Hindeodus species across the Permo-Triassic boundary (Nicoll et al. 2002; Metcalfe et al. 2007).

Ammonoid v. conodont biostratigraphy The Triassic chronostratigraphic scale was built on ammonoid biostratigraphy. However, in the 1990s, a movement to define Triassic stage boundaries with conodonts began, and at present four stage boundaries (bases of the Induan, Olenekian, Anisian and Rhaetian) are defined (or will soon be defined) by the LOs of conodont taxa, and it is likely that the base of the Norian will also be defined by a conodont datum. I view this as problematic because it broke with more than a century of literature that relied on a Triassic chronostratigraphic scale based on ammonoid-defined boundaries (which are not always the same as the conodontdefined boundaries). By thus abandoning priority, conodont-based definitions have not served the stability of the Triassic timescale. This is most clear with the Norian base, where possible conodont-based definitions are significantly different from the existing ammonoid-based definition (cf. Ogg 2004a). Furthermore, I see various problems with conodont-based Triassic biostratigraphy, including: (1) as already noted, the relative youth of Triassic conodont taxonomy, which remains unstable for many taxa; (2) reworking of conodonts, which is not easily recognized and rarely addressed (Macke & Nichols 2007); (3) problems of facies restrictions, diachroneity and provinciality, which do affect Triassic conodont distributions (Clark 1984); and (4) the invisibility of conodonts on outcrop, so that they cannot be used in the field to determine the ages of strata. Against this, I recognize that the ubiquity of conodonts in some strata, like other microfossils, presents an advantage over macrofossils, whose occurrences are much more limited. However, the ammonoid-based Triassic timescale is underpinned by nearly two centuries of collecting and taxonomic work on Triassic ammonoids. It provides macrofossil-based age assignments that can be used in the field where ammonoids occur and (outside of the Hallstatt facies) suffers from few reworking issues. Moreover, the species-level evolution through time of ammonoids is commonly obvious from the morphology of their shells,

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which record their entire life history, whereas conodont evolution is necessarily interpreted from morphological change in a particular element among the tooth-like elements that constitute their only commonly fossilized record. To a large extent, Triassic conodont-based GSSPs were an answer to longstanding disagreements over taxonomy and correlation among ammonoid specialists. They were also an important part of developing an integrated chronostratigraphic scale. As a relatively newly studied taxonomic group, Triassic conodonts did not have the perceived ‘excess baggage’ of ammonoids – a long history of taxonomic changes and disagreements, known provinciality (Tozer 1981; Dagys 1988b) and demonstrably diachronous distributions of some taxa. Furthermore, the ubiquity and perceived cosmopolitanism of conodonts as well as the retirement in the 1990s of the main Triassic ammonoid workers fueled the rise of Triassic conodont biostratigraphy. Nevertheless, I am certain that further studies of Triassic conodonts will reveal that they, too, have all of the ‘excess baggage’ of the ammonoids and are not inherently superior biostratigraphic tools with which to refine Triassic chronostratigraphy.

Biotic events Typically, the beginnings of stages have been defined by biotic events. GSSP definitions of stage bases now focus on one biotic event, marked by the LO of a single taxon, in order to unambiguously identify a single point in time at a single location. However, many of the Triassic stages were originally conceived of as turning points in ammonoid evolution: typically the appearance of new ammonoid families and/or the beginnings of new evolutionary diversifications of ammonoids. A good example is the appearance of trachyceratid ammonoids (they subsequerntly underwent a major Late Triassic evolutionary diversification), which is now used to define the base of the Ladinian (Brack et al. 2005). Arguments about the definition and recognition of Triassic stages have often been arguments over the timing and magnitude of a biotic event – is it significant enough to mark the beginning of a new stage? Indeed, most of the varied schemes for subdividing the Lower Triassic (Fig. 10) reflect differences of opinion on just such issues. Also, one of Tozer’s principal reasons for rejecting the Rhaetian Stage was that it was not linked to a significant ammonoid evolutionary event and thus not ‘worthy’ of recognition as a stage. Stage recognition also relies heavily on correlateability, both of the stage boundaries and of the stage itself. Currently, there are few problems assigning marine strata to the Triassic stages based

on biostratigraphic criteria. More difficult are substage correlations and, as I stated above, these need to be the focus of further refinement of the Triassic chronostratigraphic scale.

Short v. long stages Stages are the lowest formal unit usually recognized in the chronostratigraphic hierarchy, though, as noted here, substages have long been used to subdivide most of the Triassic stages. The Triassic chronostratigraphic scale advocated here recognizes the seven stages endorsed by the STS divided into 15 substages (Fig. 11). Arguments against recognizing the Rhaetian as a stage were in part based on it being ‘too short’ and not biotically distinctive enough (at least in terms of ammonoid evolution) to merit stage recognition. Yet, the Rhaetian as now defined is readily recognized across most of Europe, and the definition of its base identifies it as a relatively short but distinct interval at the end of Triassic time (Krystyn et al. 2007b). Arguments for a threefold or fourfold division of the Lower Triassic were rooted in recognizing key ammonoid biotic events: apperance of Otoceras at the beginning of the Griesbachian, appearance of meekoceratines at the beginning of the Dienerian, etc. (Tozer 1978). The global correlateability of these events is demonstrable. I believe the best stages are short stages, because they discriminate short intervals of time. Ironically, I think the Rhaetian is the very best Triassic stage because it represents a relatively short interval of time (as does the Induan). Problematic are the long Triassic stages – Carnian and Norian – and their substage subdivisions are a necessary refinement of the scale. In retrospect, I also see no good reason why four Lower Triassic stages could not have been recognized and ratified. But, the fourfold division of the Lower Triassic is now at the substage level. The path forward will be to formally define and better correlate the 15 Triassic substages.

Boreal v. Tethyan The Triassic chronostratigraphic scale is built on fossils collected from marine strata deposited around the periphery of the Pangean supercontinent. Most of these rocks were deposited in the paleoequatorial Tethyan region, extending from the European Alps to southern China (Fig. 9). Boreal marine rocks from Arctic Canada, and other relatively high paleolatitude strata from eastern Siberia and the mixed palaeolatitude strat of the American Cordillera also play a large role in the Triassic timescale. This has created a problem of provinciality among some of the fossil taxa (particularly some of the ammonoids: Tozer 1981; Dagys

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1988b) used for stage definition and characterization. Thus, key ammonoid taxa known in Tethys are not known in North America (or vice versa), and there is suspected diachroneity of some taxa across the provinces. This provinciality was, as noted above, one of the deficits in using ammonoids in Triassic chronostratigraphy and helped advance the apparently more cosmopolitan conodonts into a pivotal role in chronostratigraphic definitions. Thus, provinciality has created some problems for Triassic timescale construction, as it always does anywhere in the fossil record. Nevertheless, the Triassic is a relatively cosmopolitan world compared to some other time intervals, such as the Carboniferous, and I see provinciality as a minor issue in the further refinement of the Triassic chronostratigraphic scale.

Old World v. New World chronostratigraphic standards Mojsisovics et al. (1895) produced a nearly finished Triassic timescale based on ammonoid biostratigraphy using Old World standards (sections/outcrops). Subsequent correction, refinement and elaboration built those standards and others in the Old World into a robust Triassic chronostratigraphy capable of providing global correlations, particularly in ammonoid-bearing strata. Tozer, however, advocated using a new set of North American standards to subdivide Triassic time. In particular, Tozer (1971, 1984) criticized the Alpine-Mediterranean Triassic stages because of: (1) incorrect ordering of some of the ammonoid zones due to insufficient stratigraphic data; (2) condensed or mixed deposits that yield fossils of different ages from the same ‘bed’; and (3) missing ammonoid zones (time intervals without fossil representation). Much of this criticism focused on the ‘Hallstatt facies’, a relatively thin succession of mostly red micritic limestone in which condensation of macrofossils (especially ammonoids) is common, and local syndepositional dissolution has resulted in mixed fossil assemblages. This facies ranges in age from Smithian to Norian and extends from Hallstatt in Austria eastward through the Balkans into eastern Europe and is even known in Timor. Ammonoid preservation in this facies is exquisite, so many classic Triassic ammonoid localities are in Hallstatt facies strata. Tozer’s North American-based Triassic timescale introduced new stage names for the Lower Triassic, but other than the Rhaetian, Tozer’s scheme used all of the Middle and Upper Triassic stage names introduced for European standards. Predictably, Tozer’s suggestion to base the Triassic timescale on North American standards found little

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acceptance in Europe, although it was essential to unscrambling the faunal succession recorded in Hallstatt limestones and filling gaps in European sections. Priority has generally been used to place the Middle and Late Triassic GSSPs in Europe. Nevertheless, the GSSPs for the bases of the Anisian, Ladinian, Carnian and Norian, however defined, have been (or will be) placed at essentially the levels where they were drawn in North America by Tozer (1967) and Silberling & Tozer (1968). The Lower Triassic strata in western Europe never provided good standards for stage-level subdivisions, so these stage boundaries have been defined in Asia. The Canadian Arctic stratotypes of Tozer’s Lower Triassic stages were deemed too difficult to access, and this was one of the primary reasons why they were not accepted as the standard Lower Triassic stages. Nevertheless, Tozer’s North American standards have much to recommend them in terms of unambiguous superposition and extensive outcrop belts from which many data can be obtained. Certainly, his Lower Triassic stages are viable substages, and the North American record will continue to add significant data to refine Triassic chronostratigraphy, as it has recently done in the Anisian through the work of Monnet & Bucher (2005a, b).

Prospectus The two century history of the development of the Triassic chronostratigraphic scale has produced a scheme of three series, seven stages and 15 substages (Fig. 11). Many of the contentious issues surrounding this chronostratigraphic scale have been resolved, though not always for the ‘best’ reasons nor always to the satisfaction of all workers. Despite this, I believe that the seven stages of the Triassic currently recognized do a good job of honoring priority and stability in Triassic chronostratigraphic nomenclature. Formal definition of the stage boundaries is nearly complete. The frontier of Triassic chronostratigraphy is in definition and characterization of the 15 Triassic substages: this is the way forward. I am grateful to many colleagues on the STS who have taught me much about the Triassic timescale during the last two decades. Karl Krainer generously helped with the Austrian literature, and Jim Jenks provided Figure 3. Marco Balini, Jim Jenks, Jim Ogg and Norm Silberling provided helpful reviews of the manuscript.

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The Triassic timescale: new constraints and a review of geochronological data ´ ZSEF PA ´ LFY2, PAUL R. RENNE1,3 & PETER BRACK4 ROLAND MUNDIL1*, JO 1

Berkeley Geochronology Center, 2455 Ridge Road, Berkeley, CA 94709, USA

2

Research Group for Paleontology, Hungarian Academy of Sciences-Hungarian Natural History Museum, POB 137, Budapest, H-1431 Hungary

3

Department of Earth and Planetary Science, University of California, Berkeley, CA 94720, USA 4

Departement Erdwissenschaften, ETH Zu¨rich, CH-8092 Zu¨rich, Switzerland *Corresponding author (e-mail: [email protected])

Abstract: A review of geochronological data underlying the geological timescale for the Triassic yields a significantly different timescale calibration than that published in the most recent compilation (Geologic TimeScale 2004). This is partly due to the availability of new radio–isotopic data, but mostly because strict selection criteria are applied and complications arising from biases (both systematic and random) are accounted for in this contribution. The ages for the base and the top of the Triassic are constrained by U– Pb ages to 252.3 and 201.5 Ma, respectively. These dates also constrain the ages of major extinction events at the Permian– Triassic and Triassic– Jurassic boundaries, and are statistically indistinguishable from ages obtained for the Siberian Traps and volcanic products from the Central Atlantic Magmatic Province, respectively, suggesting a causal link. Ages for these continental volcanics, however, are mostly from the K –Ar (40Ar/39Ar) system, which requires accounting and correcting for a systematic bias of c. 1% between U –Pb and 40Ar/39Ar isotopic ages (the 40Ar/39Ar ages being younger). Robust age constraints also exist for the Induan– Olenekian boundary (251.2 Ma) and the Early– Middle Triassic (Olenekian–Anisian) boundary (247.2 Ma), resulting in a surprisingly short duration of the Early Triassic, which has implications for the timing of biotic recovery and major changes in ocean chemistry during this time. Furthermore, the Anisian–Ladinian boundary is constrained to 242.0 Ma by new U– Pb and 40Ar/39Ar ages. Radio– isotopic ages for the Late Triassic are scarce, and the only reliable and biostratigraphically-controlled age is from an upper Carnian tuff dated to 230.9 Ma, yielding a duration of more than 35 Ma for the Late Triassic. All of these ages are from U–Pb analyses applied to zircons with uncertainties at the permil level or better. The resulting compilation can only serve as a guideline and must be considered a snapshot, resolving some of the issues mainly associated with inaccurate and misinterpreted data in previous publications. However, further advances will require revision of some of the data presented here.

The Triassic Period marks an unusual time in Phanerozoic Earth history in that its beginning and end are marked by extinction events and episodes of continental scale volcanism (Siberian Traps and Central Atlantic Magmatic Province or CAMP, respectively) and in that continents are arranged in a unique assembly (Pangaea). Low faunal and floral diversity in the Early Triassic, a protracted biotic recovery, global perturbations in the oceanic carbon cycle and reduced carbonate sedimentation are only some of the striking phenomena following the end-Permian extinction. Each of these is in need of an accurate time frame to help a better understanding and eventually numerical modelling of the underlying processes. The large-scale Middle Triassic biotic radiation, high levels of Late Triassic faunal diversities, and the end-Triassic mass

extinction and its relationship to CAMP are also the subjects of much current research that require a reliable time frame. The reconstruction of these geological events calls for reliable and high-resolution dating of rocks to infer the age and tempo of processes. This is, in most cases, provided by radio–isotopic dating techniques. In principle it is possible to produce radio– isotopic ages at or even below the permil level of resolution because the actual age depends to the first order on the ability to precisely measure isotopic ratios, which is possible even for samples as small as fragments of single crystals. The achievement of permil age resolution is largely restricted to U –Pb and 40Ar/39Ar methods, upon which the modern Phanerozoic timescale is largely built, and is due mainly to advances in

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 41– 60. DOI: 10.1144/SP334.3 0305-8719/10/$15.00 # The Geological Society of London 2010.

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mass spectrometry and sample preparation (Fig. 1). Despite the potential for permil age resolution using these methods, there are many pitfalls when radio – isotopic ages are used and taken at face value, and it is not always clear when and to what extent ‘highresolution’ ages are also accurate and have the inferred (or desired) geological meaning. In many (arguably most) cases, the search for information about geological time begins with timescale compilations and often users fail, or they are not equipped, to review the underlying data from the original publications. Unfortunately, timescale compilations mostly fail to account for complications related to precision and accuracy of published radio – isotopic ages, although some of the issues have been discussed in detail (Villeneuve 2004). This contribution aims at discussing and clarifying some of the aspects related to radio–isotopic ages from the Triassic in more detail than is usually provided by timescale compilations. In doing so, we hope to provide more transparency so that radio– isotopic ages can be better evaluated and realistically used by the reader. The resulting compilation and a comparison with previous timescales (Harland et al. 1990; Gradstein et al. 1994; Odin 1994; Gradstein et al. 2004) are summarized in Figure 2.

Radio – isotopic ages from previous studies Before discussing modern age data (i.e. results generated after 1995) for the Triassic it is useful to briefly review some older age data that served as a basis for timescale constructions until the mid 1990s. This review mainly serves as a reference frame in order to emphasize and contrast how dramatically the resolution of radio –isotopic ages has changed, often, however, with unquantifiable consequences for the robustness (in the sense of accuracy) of more recently obtained highly precise data. Moreover, the data below are listed for completeness, as it may be worth revisiting some of the samples. All ages have been recalculated using the decay constants given in Steiger & Ja¨ger (1977). It is evident that the database for the Triassic timescale had been extremely scarce until about a decade ago. Most of the then-available age determinations were ambiguous either with respect to their stratigraphic level, or their analytical reliability (or both). In the following, such age data are listed according to their assumed stratigraphic age in ascending order (uncertainties are at the 1s level unless quoted otherwise). Pb –alpha dates of 230 + 40 and 290 + 45 Ma for the Koipato rhyolite porphyry that intrudes the Rochester Rhyolite within the Early Triassic Koipato Group in the Humboldt Range (Nevada) are associated with large uncertainties but bear

potential for the application of state of the art U – Pb zircon analyses (Silberling 1973; Wallace 1960). The Neara Volcanics in the Toogoolawah Group in SE Queensland (Australia) consist of 5000 m of sediment intercalated with andesites and trachytes. Based on palynogical evidence the Toogoolawah Group is thought to be of Anisian to Ladinian age (De Jersey 1972). From two different horizons within these volcanics, Irwin (1976) reported two conventional K –Ar ages of 242 + 5 Ma (whole rock) and 239 + 5 Ma (hornblende). The Kin Kin Beds of the Gympie Basin (Queensland, Australia) of probably Early Triassic age (Smithian), based on the occurrence of ammonites, are intruded by the Goomboorian diorite, yielding a K –Ar age of 236 + 7 Ma (Green & Webb 1974; Murphy et al. 1976). Reptile assemblages in the Puesto Viejo Fm. in Argentina suggest an Early to Middle Triassic (probably Scythian) age. Intercalated acidic and basic volcanics have been dated by K –Ar techniques at 237 + 4 Ma (Valencio et al. 1975). Of particular significance for previous timescale compilations were K –Ar and 40Ar/39Ar ages of Hellmann & Lippolt (1981) determined for K – feldspars from bentonites of the biostratigraphically well-constrained Anisian/Ladinian Grenzbitumenzone at Monte San Giorgio in Ticino (southern Switzerland). One of these horizons was re-examined in Mundil et al. (1996c, sample MSG.09, see also below for new ages). The feldspars dated by Hellmann & Lippolt (1981) included both volcanic sanidines (assumed to be primary) and secondary authigenic K –feldspars. High-sanidine phenocrysts yielded an average K –Ar age of 232 + 9 Ma (combining results from several horizons and taking calibration errors into account) and phenocrysts composed of high-sanidine cores with a secondary overgrowth of K –feldspar had an average K –Ar age of 235 + 6 Ma. Analyses performed by 40Ar/39Ar-techniques yielded plateau ages of 232 + 8 Ma and 237 + 7 Ma, respectively. Wholly authigenic feldspars yielded an average K –Ar age of 226 + 8 Ma and a 40Ar/39Ar-plateau age of 225 + 9, which was interpreted to represent the age of alteration transforming original volcanic tuffs into bentonites. On the basis of these results Forster & Warrington (1985) proposed an age of 235 + 5 Ma for the Anisian –Ladinian boundary, however, the use of an obscure standard as fluence monitor for 40Ar/39Ar analyses in Hellmann & Lippolt (1981) makes evaluation of the latter age difficult. In New Zealand, the Etalian –Kaihikuan stage boundary (possibly equivalent in age to the Anisian –Ladinian boundary) has been dated by the 40Ar/39Ar-laser technique on single biotite crystals and the U –Pb technique on multi-grain zircon

TRIASSIC RADIO –ISOTOPIC AGES

samples (Retallack et al. 1993). The 40Ar/39Ar biotite age spectra are discordant, but yield a weighted mean plateau age of 242.8 + 0.6 Ma, which is complemented by 238U/206Pb ages of 236.8 + 2 Ma to 243.1 + 2 Ma from zircons displaying variable degrees of Pb loss. These data have not been considered in current timescales. The discordant biotite age spectra resemble those arising from interlayer alteration (e.g. Min et al. 2001), in which case the plateau ‘ages’ may be spuriously old. Another minimum age for the Middle Triassic is provided by Webb & McDougall (1967), who dated small granite plutons in the Maryborough Basin (Queensland, Australia). They report Rb –Sr isochron ages of 226 + 16 Ma and K –Ar ages of 222 + 2.5 Ma. The intrusion ages are considered to be bracketed by Early to Middle Triassic and Early Jurassic sediments, respectively. The Somerset Dam Gabbro cutting the Neara Volcanics (see above) was dated by Webb & McDougall (1967) using the K –Ar techniques to 219 Ma (hornblende) and 211 Ma (plagioclase), respectively (no errors quoted). Dykes of a swarm with an average K –Ar age of 223 Ma (hornblende, no errors quoted) are thought to postdate Anisian – Ladinian successions. The Predazzo and Monzoni igneous complexes in the Southern Alps (N Italy) postdate Early Ladinian platform sediments as dykes related to the intrusive bodies and related dykes cut basinal and platform carbonates of Ladinian age. Numerous age dates have been reported for these complexes by Borsi & Ferrara (1968a, b), Ferrara & Innocenti (1974) and Ko¨hler (in Masch & Huckenholz 1993), but several discrepancies exist when Rb –Sr and K –Ar data are compared. The phlogopite K –Ar ages from veins in the aureole of the Monzoni complex were considered to provide the best estimate of the age of this complex at 231 + 8 Ma. This value is consistent with Rb –Sr cooling ages of 235 + 5 Ma for biotite from a monzonite and from a contact rock sample (Ko¨hler, 1990 pers. commun.). The Predazzo igneous complex – in particular the late granitic member – has been re-examined by 40 Ar/39Ar techniques (Laurenzi et al. 1994) and with U –Pb single zircon techniques. Laurenzi et al. (1994) report an 40Ar/39Ar age of 234.7+0.44 Ma for biotite; the zircons yield a 206Pb/238U age of 237.2 + 0.4/21.0 Ma (2s error: Mundil et al. 1996a). Gehrels et al. (1987) give a U –Pb zircon age of 225 + 3 Ma for a supposedly earliest Norian rhyolite of the Hyd Group in Annette and Gravina islands (SE Alaska). Long & Lehman (1993) mention a Rb –Sr model age on biotite of c. 210 Ma for Upper Triassic sandstone (upper part of the Cooper Canyon Formation, SW USA).

43

The majority of the age values presented above are from intrusions with sometimes ambiguous relationships to sediments. Unavoidably this has a profound impact on the quality of the resulting timescale. Based on some of the data listed above, Forster & Warrington (1985) estimate the following ages for Triassic stage boundaries (ages refer to the base of a given stage). 250 + 5 Ma for the base of the Triassic (base of the Scythian stage); 242 + 5 Ma for the Anisian; 235 + 5 Ma for the Ladinian; 230 + 5 Ma for the Carnian; 220 + 5 Ma for the Norian; 210 + 5 Ma for the Rhaetian; and 205 + 5 Ma for the Triassic–Jurassic boundary (base of the Jurassic). Other timescales (Harland et al. 1990; Gradstein et al. 1994; Odin 1994) of similar vintage rely on largely the same data and there is no particular reason why we have chosen this example from the Forster & Warrington (1985) timescale. What all timescales have in common is the assignment of several percent uncertainties (5 to 10 Ma) to stage boundaries as a result of uncertain stratigraphic levels and analytical ambiguities of the underlying data listed above. In view of the quality of the underlying data the ages for stage boundaries by Forster & Warrington (1985) in their timescale is remarkably robust in the sense that their uncertainties, although fairly large, overlap with our current best estimates.

Recent advances and unresolved problems in geochronology In the past c. 15 years, scores of new radio-isotopic ages have been produced, most of which are associated with very small uncertainties (at the permil level). If associated with tight stratigraphic constraints, these ages are remarkably well-suited for timescale calibration purposes. The new ages are exclusively U –Pb or 40Ar/39Ar ages on volcanic products, either from ash-falls intercalated within sedimentary sections or from continental flood basalts (CFBs). It is to be expected that the quality of data, in this case radio–isotopic ages from different isotopic systems, will improve further as new analytical techniques are discovered and applied, and the knowledge about primary input values such as decay constants is also growing. The accuracy of radio-isotopic ages mainly depends on two prerequisites: (1) a closed isotopic system of the analyzed mineral or rock, that is, no gain or loss of the parent and/or daughter element; and (2) the quality of the calibration of reference materials (natural or artificial standards) and decay constants. Prior to a major breakthrough in U –Pb zircon dating (CA – TIMS, or annealing followed by chemical abrasion, a pre-treatment developed by Mattinson (2005),

44

R. MUNDIL ET AL.

effectively resulting in zircon crystals having a closed isotopic system) the first criterion was often not met in previous studies, resulting in ages that are slightly too young (due to the loss of the radiogenic daughter isotope), and it has often either been ignored or was not recognized (see Mundil et al. (2001) for discussion). The second prerequisite imposes a problem if ages from different isotopic systems are compared. For example, recent studies show evidence that the currently used decay constant for 40K is substantially miscalibrated by as much as 1%, which vastly exceeds the currently attainable analytical precision of c. 0.1% (Renne et al. 1998a; Min et al. 2000; Mundil et al. 2006). This miscalibration and the use of an inaccurate value (mostly arising from the decay constant miscalibration) for the age of the fluence monitor used for 40Ar/39Ar dating (here 28.02 Ma for Fish Canyon sanidine) requires a correction if 40Ar/39Ar ages are compared to ages from different isotopic systems. A recent study by Kuiper et al. (2008) shows similar findings (overlapping within uncertainties) using a calibration based on orbital tuning. It is important to note this bias is mainly due to a miscalibration of the electron capture branch of 40K decay, and therefore changes with time (i.e. it is smaller for Proterozoic and older ages). Details about how 40 Ar/39Ar data are affected can be found in numerous references (e.g. Renne et al. 1998b; Min et al. 2000; Kuiper et al. 2005; Mundil et al. 2006; Bachmann et al. 2007). However, only recently have systematic errors related to inaccurate decay constants been accounted for, which is negligible for U –Pb ages (at the permil level) but substantial for K/Ar and 40 Ar/39Ar ages (at the percent level). In particular, research aimed at studying the synchroneity of events has to take this complication into consideration. The individual data listed in the following section are in stratigraphic order and any complications arising from the two above-mentioned points will be considered. Uncertainties are given at the 2s or 95% confidence level.

Modern radio – isotopic ages for the Triassic Age of the Permian – Triassic boundary The Permian –Triassic transition is marked by the most severe biotic crisis in the Phanerozoic, although as per definition the boundary has been agreed upon by the first appearance of the conodont species Hindeodus parvus at the GSSP (Global Stratotype Section and Point: Yin et al. 2001), shortly after the main pulse of the extinction. A series of volcanic ashes suitable for radio-isotopic dating is

intercalated within marine sediments of Permian to Triassic age in central and south China. The ash bed closest to the mass extinction, and just predating the biostratigraphic boundary (bed 25 of Yin et al. 2001), was initially dated by U –Pb zircon SHRIMP techniques to 251.2 + 3.4 Ma. A later study (Black et al. 2003) evaluating the robustness of natural zircon standards used for the U –Pb SHRIMP method concludes that the SL13 standard used in Claoue´-Long et al. (1991) yields ages that are on average about 1.0% younger than those based on currently used standards. A conversion factor, however, cannot be applied because of demonstrable Pb/U heterogeneity within SL13. Therefore, the accuracy of the SHRIMP age for bed 25 is biased by an unquantifiable magnitude. A subsequent study used the U –Pb ID– TIMS method (Isotope Dilution Thermal Ionization Mass Spectrometry) applied to dissolved single- and multi-crystal zircon samples from several ash layers of the Meishan section (Bowring et al. 1998). The precision of radio– isotopic ages of small samples from the ID –TIMS method depends at the first order on the ability to precisely measure isotopic ratios of low intensity ion beams, which in principle is possible to better than 1 permil uncertainty with modern detectors. We now know that this analytical precision cannot always be equated with accuracy. In summary, the interpolated age for the base of bed 27a (the base of the Triassic) from the Bowring et al. (1998) study is 251 Ma based on ages of 250.7 + 0.3 Ma and 251.4 + 0.3 Ma of bracketing layers. Mundil et al. (2001), using an approach similar to the latter study on samples from the same layers, failed to provide a conclusive data set but deduced that the ages from Bowring et al. (1998) are affected by Pb loss and therefore likely too young (by a small but indeterminable amount). This is because the averaging effects from the use of multi-crystal samples (employed to provide large, precise ion beams in the mass spectrometer) of zircons mask real, Pb-loss-related age dispersion within a zircon population. In addition to this unresolved complication, the majority of the ages of individual analyses were excluded from the calculation of the precise mean. Since the magnitude of Pb loss for zircons from each individual ash layer is unknown and likely to be variable, it would be imprudent and misleading to conclude that the age differences can still be used for stratigraphic purposes, such as the derivation of sedimentation or extinction rates (as is suggested in Erwin 2006). As mentioned above, it is now clear that many of the studies employing the air abrasion technique (Krogh 1982) to zircons were only partially successful in eliminating Pb loss as displayed by age dispersion of single-crystal analyses of abraded

TRIASSIC RADIO –ISOTOPIC AGES

zircons (e.g. Mundil et al. 1996c, 2001). Despite these concerns, the interpolated age of 251 + 0.4 Ma from Bowring et al. (1998) has been assigned to the Permian–Triassic boundary by the most recent timescale compilation (Ogg 2004). A subsequent study presented single zircon ID – TIMS data from ash layers of the Shangsi section (Sichuan province), which is tightly correlated with the Meishan section, and an age for bed 25 in Meishan (Mundil et al. 2004). This study used the newly developed CA–TIMS technique (Mattinson 2005) applied to individual crystals. Not surprisingly, the closed system zircons yielded older and more accurate ages than those of previous studies. Zircons from bed 25 in Meishan yielded a weighted mean age of 252.4 + 0.3 Ma. This is resolvably older than the age for this bed from Bowring et al. (1998), and thus places the boundary at a slightly older age. New U –Pb zircon analyses on the same samples using the CA –TIMS technique (Crowley et al. 2006) appear to confirm the results from Mundil et al. (2004). The U –Pb ages in Mundil et al. (2004) are confirmed by an array of 40 Ar/39Ar ages from the same ash layers that are consistently younger by about 1% when compared to the U – Pb ages (Renne et al. 2004), except for the youngest age (sample SH32 in Mundil et al. (2004), see below). The 40Ar/39Ar ages in Renne et al. (2004) confirm 40Ar/39Ar ages of 249.9 + 0.2 Ma and 250 + 0.5 Ma for bed 25 in Meishan and an ash bed just above the Permian– Triassic boundary in Shangsi (Renne et al. 1995). Accounting for the bias yields corrected 40Ar/39Ar ages consistent with the U – Pb ages in Mundil et al. (2004).

Age of the Siberian Trap Although radio –isotopic ages from units of the Siberian Trap volcanics are of limited value for the purpose of timescale calibration, it is useful to discuss them here as they have a profound bearing on testing the hypothesis that CFB (Continental Flood Basalt) events and mass extinctions are causally linked. Effusive volcanic products from this type of volcanism are mostly basaltic and contain extremely small amounts of U-bearing mineral phases suitable for U – Pb dating. The method of choice is therefore 40Ar/39Ar dating applied to plagioclase from basalts. Initial geochronological studies (both U –Pb and 40Ar/39Ar) were associated with relative large uncertainties (Baksi & Farrar 1991; Dalrymple et al. 1991; Campbell et al. 1992). In their study from 1995, Renne et al. (1995) recalculated precise and undistinguishable 40 Ar/39Ar ages of samples near the bottom and top of the main pulse of the Siberian volcanics (Renne & Basu 1991) to 250.0 + 1.6 Ma (the

45

recalculated age is due to a revised age of the neutron flux monitor Fish Canyon sanidine). A subsequent 40Ar/39Ar study by Venkatesan et al. (1997) reported ages indistinguishable from those of Renne et al. (1991). Accounting for the already mentioned systematic bias of 1% yields ages of 252.5 Ma for the U –Pb normalized 40 Ar/39Ar ages in Renne et al. (1995), indistinguishable from the U –Pb age for the Permian–Triassic boundary and mass extinction in Mundil et al. (2004). The conclusions and ages from the Renne et al. (1995) study were confirmed by a number of 40 Ar/39Ar ages in Reichow et al. (2002), which also find further evidence for the main pulse at 249.4 + 0.5 Ma from mostly subsurface samples to the west of the area were the Siberian CFBs crop out, essentially doubling their area of spatial extent. U – Pb studies by Kamo et al. (1996, 2003) yield zircon, baddeleyite and perovskite ages between 250.2 and 251.7 Ma (all with uncertainties between 0.3 and 0.4 Ma) for various volcanic and intrusive units whose chronostratigraphic relationship to the main Siberian Traps are not always unambiguous. None of the analyses were performed on CA pre-treated crystals, and for reasons given above (and despite their high analytical quality), the ages may be therefore considered as minimum ages. In addition, the U –Pb perovskite age is sensitive to common Pb corrections, which introduces an additional source of ambiguity not fully accounted for.

Early Triassic ages As mentioned above, the U –Pb zircon age for an ash in the lowermost Triassic (252.5 + 0.2 Ma in Mundil et al. 2004) may be compromised by re-deposited zircons, as indicated by a greater than 1% difference to a 40Ar/39Ar age of 249.3 + 0.3 Ma for the same horizon (Renne et al. 2004). We therefore conclude that the ‘U –Pb normalized’ 40Ar/39Ar age of 251.8 Ma is more reliable, which has a small effect on the age of the Permian-Triassic boundary chosen here. The oldest reliable Early Triassic U –Pb age using the CA– TIMS technique is from analyses of three single zircons from an ash within the lower Olenekian Kashmirites densistriatus beds yielding an age of 251.2 + 0.2 Ma (Galfetti et al. 2007), placing the Induan–Olenekian boundary at a slightly older age and allowing only 1 Ma for the duration of the Induan stage. Employing the same pre-treatment to some of the zircons of their study (air abrasion to the remaining ones), Ovtcharova et al. (2006) present a series of ages for Lower to Middle Triassic ash beds from fossiliferous sections in South China. An ash layer within the Tirolites beds (Olenekian stage, Lower Triassic) is dated to

46

R. MUNDIL ET AL.

250.6 + 0.5 Ma. An age of 248.1 + 0.4 Ma is given for zircons from an ash bed in the Neopopanoceras haugi zone within the Olenekian and below the Lower to Middle Triassic boundary. The Lower– Middle Triassic (i.e. Olenekian –Anisian) boundary is bracketed by the latter age and an age of 246.8 + 0.4 Ma from within the lower part of the Middle Anisian. The youngest age from their dataset is from within the upper part of the Middle Anisian (Balatonites shoshonensis zone) and is dated to 244.6 + 0.5 Ma. Uncertainties arising from the calibration of the mixed tracer solution were supposedly included but unfortunately this augmentation was applied to the individual zircon ages rather than to the final weighted mean age, so that its effect is spuriously reduced and contributes to the apparent coherence of the data. This produced, in this case, only a very minor effect with insignificant impact on the accuracy of the data and the conclusions. A striking result of this study is the finding that the entire Early Triassic is extremely short (4.5 + 0.6 Ma), which has a profound impact on the duration of the observed perturbations in ocean chemistry (Payne et al. 2004) and their interpretation, essentially allowing less time for them to occur than previously thought. In a subsequent study, Lehrmann et al. (2006) present four U –Pb ages for ash layers bracketing the Lower to Middle Triassic boundary. All of the analyses are on single zircons that were either pre-treated with the CA technique or air abraded. All four data sets show dispersion of the individual 206Pb/238U ages, indicating that both older antecrysts from a previous magmatic cycle and younger ages due to Pb loss may be present. The extracted ages are from a subjectively chosen ‘coherent’ subset resulting in extremely small uncertainties of +0.1 Ma for each of the four layers. In detail, Lehrmann et al. (2006) give ages of 247.38 + 0.10 Ma and 247.32 + 0.08 Ma for two ash layers in the uppermost Olenekian, and ages of 247.13 + 0.12 Ma and 246.77 + 0.13 Ma within the lowermost and Middle Anisian, respectively. In a discussion, yet another younger age of 246.30 + 0.07 Ma is given (no data table is provided), also based on air-abraded and chemically abraded crystals (Bucher et al. 2007; Lehrmann et al. 2007; Ramezani et al. 2007). This ash is from a different section but supposedly slightly younger than the youngest ash in Lehrmann et al. (2006), which yielded 246.8 + 0.1 Ma. The data, however, show substantial dispersion, as do the ones in Lehrmann et al. (2006). Given the complexity of the individual zircon data and the fact that the data selection includes ages that are clearly affected by Pb loss, the uncertainties are probably underestimated. In general there is agreement with the two studies (Lehrmann et al.

2006; Ovtcharova et al. 2006), placing the Early – Middle Triassic boundary at c. 247.2 Ma. Although the constraining ages are significantly older than in previous compilations, they result in a duration of the Early Triassic of c. 5 Ma, which is comparable to the duration deduced for the GTS2004. Estimates for the latter were based on utilizing a composite standard from graphical correlation of Lower Triassic sections (Sweet & Bergstrom 1986) and by proportional scaling of ammonite zones according to subzones.

Middle Triassic ages We present new 206Pb/238U zircon and 40Ar/39Ar sanidine ages (MSG.09) from a volcanic ash within a series of latest Anisian fossiliferous dolostones and bituminous shales (Grenzbitumenzone) sampled at Monte San Giorgio (Switzerland; 458540 3500 N, 88560 3200 E, WGS84) (Table 1). The volcanic ash corresponds to bed 71 in Rieber (1973) and is the same one dated to 241.2 + 0.8 Ma (Mundil et al. 1996c) using U –Pb on zircon and 226 + 7 Ma (Hellmann & Lippolt 1981), respectively, using conventional K –Ar techniques on a feldspar mixture. This example of a previously analyzed ash fall demonstrates the shortcomings associated with previous techniques resulting in inaccurate ages. The assemblage of age-diagnostic fossils (ammonoids and various species of the bivalve Daonella) comprises species indicative of the Reitzi and the Secedensis (Nevadites) Zones (Brack & Rieber 1993). The ash fall horizons of the Grenzbitumenzone consist of a microcrystalline matrix of illite – montmorillonite and contain phenocrysts of alkali feldspar, quartz and accessory zircon. Because earlier studies show the effects of Pb loss (Mundil et al. 1996c), the zircons were annealed and chemically abraded following the methods described by Mattinson (2005) and modified by Mundil et al. (2004). The zircon residues were small (1.1 mg average) and low in U concentration (70 ppm average) but six out of eight age analyses were sufficiently precise to yield a 206 Pb/238U age of 242.1 + 0.6 Ma (MSWD 1.5, uncertainties are given at the 2s level, Fig. 1b). The remaining two analyses were imprecise (.1% uncertainty on 206Pb/238U) and thus rejected (with an insignificant effect of ,0.05% on age and uncertainty). The age is slightly older than (but within error of) the one by Mundil et al. (1996c), most likely due to elimination of Pb loss effects employing chemical abrasion. Analytical and data reduction are described in detail in the supporting online material of Mundil et al. (2004). Alkali feldspars were extracted from the light, non-magnetic fraction of residue from zircon

TRIASSIC RADIO –ISOTOPIC AGES

separation. Crystals were selected individually from the 100 to 180 micron fraction in mineral oil to avoid grains with impurities or overgrowths. Following ultrasonic rinsing in acetone followed by alcohol, the sample was cleaned ultrasonically in 3.5% HF. The sample was irradiated for 20.0 hrs. (BGC irradiation # 354PR) in the cadmium-lined CLICIT facility of the Oregon State University TRIGA reactor, together with grains of the Fish Canyon sanidine (FCs) standard in the same well of an Al disc similar to those used by Renne et al. (1998b). Individual analyses of 10 single grains of FCs yielded a J value of 0.005131+0.000010 based on the age of 28.02 Ma for FCs (Renne et al. 1998b). Approximately 5 mg of MSG.09 sanidine were analyzed by incremental heating with a CO2 laser whose beam was broadened with an integrator lens. Heating was achieved by maintaining constant laser power for 60 s per step. The resulting gas was purified by 180 s of gettering time for each step. Relative abundances of Ar isotopes from each heating step were determined by peak-hopping (10 cycles of magnetic field switching) with an MAP 215 –50 sector mass spectrometer using a Balzers electron multiplier detector in analog mode. Relative abundances were obtained by regression of peak-height versus time data to an initial equilibration time using linear or parabolic fits. Mass discrimination was monitored by 20 air pipettes bracketing the sample and standards, yielding a mean value of 1.00425 + 0.00298 per atomic mass unit based on a power law correction. Backgrounds were measured between every three heating steps

47

or fusions (standards), yielding values comparable to those reported by Renne et al. (1998b). Average values and their standard deviations were used to make the background corrections. Isotope data (Table 2) corrected for backgrounds, mass discrimination and radioactive decay were corrected for interfering reactions on K and Ca (Renne et al. 2005) and Cl (Renne et al. 2008). Ages in Table 2 are calculated relative to the 40K decay constants of Steiger & Ja¨ger (1977). The plateau age of 239.5 + 0.5 Ma (Fig. 1c) was determined by computing the inverse variance weighted mean of (40Ar*/39ArK) for all plateau steps, and applying the age equation (and error propagation) to the resulting value. The plateau age uncertainty does not include contributions from decay constants or the age of the standard. Taken at face value, the U –Pb and 40Ar/39Ar ages for MSG.09 do not overlap but are in agreement if the latter is corrected for the 1.0 + 0.3% bias as suggested in Mundil et al. (2006) based on 11 data pairs from rapidly-cooled volcanic rocks of different ages. The base of the Ladinian stage (first appearance of Eoprotrachyceras curionii at the GSSP of Bagolino, Italy: Brack et al. 2005) is bracketed by primary volcaniclastic layers, some of which have been dated by U –Pb single zircon techniques (Mundil et al. 1996b, c). Interpolating zircon ages from five airborne volcanic layers within fossiliferous sediments in the Southern Alps yielded an age of c. 240.5 Ma for the AnisianLadinian boundary, however, all of the zircons in that study were pre-treated by the air abrasion technique, which is now known to be not entirely

Fig. 1. Radio– isotopic ages for bed 71 (Rieber 1973) at Monte San Giorgio (MSG.09), constraining the Anisian– Ladinian boundary. (a) Concordia diagram showing previous U– Pb single zircon data from Mundil et al. (1996c) as well as data from zircons subjected to mild HF leaching (described in Mundil et al. 2001) affected by Pb loss; (b) Concordia diagram showing new U–Pb single zircon data using the CA–TIMS method (Mattinson 2005) and yielding a coherent age; (c) 40Ar/39Ar apparent age spectrum from step heating experiment on alkali feldspar separate. Note the 1% age bias compared to the U–Pb zircon age. See Tables 1 (U –Pb) and 2 (40Ar/39Ar) for analytical data.

48

Table 1. Analytical data for U –Pb age of a volcanic ash within a series of latest Anisian fossiliferous dolostones and bituminous shales (Grenzbitumenzone) sampled at Monte San Giorgio Isotopic ratios

Sample

ppm U

Pb(b) c. (pg)

ThðcÞ U

0.3 0.7 0.6 0.5 0.9 0.8 2.3 1.5

46 90 82 64 79 30 46 112

0.9 0.9 0.9 0.8 1.3 0.8 0.8 0.9

0.53 0.67 0.73 0.78 0.88 0.55 0.58 0.65

206

Pb 204 Pb

ðdÞ

56 190 149 113 154 86 337 485

207

ðdÞ

Pb 206 Pb

0.05974 0.05158 0.05314 0.05381 0.05109 0.05231 0.05119 0.05131

+ % 23.8 6.1 8.1 10.7 7.9 15.1 3.3 2.2

207

Pb 235 U

ðeÞ

0.3184 0.2749 0.2824 0.2846 0.2702 0.2762 0.2702 0.2703

+ % 25.4 6.5 8.5 11.4 8.3 16.0 3.5 2.4

206

Pb 238 U

ðeÞ

+ %

206 (f)

r

207

Pb 238 U

0.038644 1.83 0.87 244.4 + 4.5 0.038658 1.12 0.45 244.5 + 2.7 0.038535 0.64 0.77 243.7 + 1.6 0.038356 0.74 0.86 242.6 + 1.8 0.038357 0.51 0.85 242.6 + 1.2 0.038287 0.98 0.88 242.2 + 2.4 0.038278 0.27 0.71 242.2 + 0.7 0.038206 0.25 0.57 241.7 + 0.6 Weighted mean age 242.1 + 0.6

Pb 235 U

280.6 + 71.2 246.6 + 16.1 252.5 + 21.6 254.3 + 28.9 242.8 + 20.1 247.6 + 39.6 242.8 + 8.5 242.9 + 5.8 MSWD 1.5

207

Pb

206 Pb

594 + 516 267 + 140 335 + 183 363 + 242 245 + 181 299 + 345 249 + 76 255 + 52

(a)

Sample weight is calculated from crystal dimensions and is associated with as much as 50% uncertainty (estimated). Total common Pb including analytical blank (analytical Pb blank is 0.8 + 0.3 pg per analysis). Blank composition is Pb/204Pb ¼ 38.07 + 1.56 (all 2s of population), and a 206Pb/204Pb-207Pb/204Pb correlation of þ0.9. (c) Present day Th/U ratio calculated from radiogenic 208Pb/206Pb and age. (d) Measured value corrected for tracer contribution and mass fractionation (0.15 + 0.09%/amu). (e) Ratios of radiogenic Pb versus U; data corrected for mass fractionation, tracer contribution and common Pb contribution. (f) Correlation coefficient of radiogenic 207Pb/235U versus 206Pb/238U. Uncertainties of individual ratios and ages are given at the 2s level and do not include decay constant errors. Ratios involving 206Pb are corrected for initial disequilibrium in 230Th/238U adopting Th/U ¼ 4 for the crystallization environment. Weighted mean 206Pb/238U age (in italics) is calculated from individual ages in bold. (b)

208

206

Pb/204Pb ¼ 18.55 + 0.63,

207

Pb/204Pb ¼ 15.50 + 0.55,

R. MUNDIL ET AL.

MSG09.Z29 MSG09.Z23 MSG09.Z26 MSG09.Z27 MSG09.Z25 MSG09.Z28 MSG09.Z24 MSG09.Z22

mg(a) zirc.

Isotopic ages

TRIASSIC RADIO –ISOTOPIC AGES

successful in eliminating Pb loss (see MSG.09), so that the latter must be regarded as a minimum age. The same applies to four U –Pb zircon ages (241.1 + 0.5 Ma, 241.2 + 0.4 Ma, 240.5 +0.5 Ma and 240.4 + 0.4 Ma) from ash layers within different sub-zones of the Anisian Reitzi zone within the Forra´s Hill section near Felso´´o¨rs (Hungary), which suggest an age younger than 241 Ma for the Anisian-Ladinian boundary (Pa´lfy et al. 2003). Also, the analyses in that study are on multi-crystal samples that typically mask any complications arising from Pb loss (see above). In this context, it is important to mention that three U –Pb zircon ages of 242.6 + 0.7 Ma, 241.2 + 0.7/20.6 Ma (LAT30) and 241.7+1.5/20.7 Ma from ash layers within the uppermost Anisian of the Latemar platform are also too young by a small but unknown amount. This suspicion was subsequently confirmed with a coherent age of 242.8 + 0.2 Ma for LAT30 on CA-treated zircons from an ash layer within the uppermost Reitzi zone in the Latemar platform, which is in agreement with the new age result for MSG.09, suggesting that the Anisian-Ladinian boundary is c. 242.0 Ma (Brack et al. 2007). This is further substantiated

49

by a minimum age of 243 Ma (U –Pb on CA pretreated zircons) for an ash layer in the underlying Paraceratites trinodosus zone at Monte Bivera (Brack et al. 2007). Reliable radio-isotopic ages for the Ladinian stage are scarce, and three U –Pb single zircon ages between 238 Ma and 239 Ma for the Gredleri and Archelaus zones (Mundil et al. 1996c) must be considered minimum ages because the crystals were not subjected to the CA pre-treatment and therefore not devoid of Pb loss (see above). This is in agreement with a recent age of 239.3 + 0.2 Ma from CA pre-treated zircons from a volcanic tuff within the transition from the Gredleri to the Archelaus zone (Bru¨hwiler et al. 2007) as well as with a U– Pb zircon age of 237.3+0.4/21.0 Ma for the Late Ladinian granites at Predazzo (Brack et al. 1997; Mundil et al. 1996a).

Late Triassic ages A U –Pb single zircon age of 230.9 + 0.3 Ma for an ash bed within marine fossiliferous deposits of Late Carnian age (Metapolygnathus nodosus zone) is the only reliable constraint for the entire early portion of

Fig. 2. Compilation of Triassic timescales. Suggested stage boundaries from this study are based on weighted mean 206 Pb/238U single zircon ages (mostly applying the CA–TIMS technique: Mattinson, 2005) from biostratigraphicallycalibrated volcanic deposits. Uncertainties of the individual ages include uncertainties from the calibration of mixed tracer used in ID–TIMS analyses.

50

Table 2. Isotope data corrected for backgrounds, mass discrimination and radioactive decay with Triassic isotopic ages calculated relative to the 40K decay constants of Steiger & Ja¨ger (1977) Power (W)

Ar (mol)

8.42E-14 2.90E-13 2.97E-13 3.66E-13 3.42E-13 2.75E-13 2.14E-13 1.53E-13 1.19E-13 9.15E-14 5.15E-14 3.64E-14 3.33E-14 3.09E-14 1.66E-14

40

Ar (nA) + %

2.3531 8.0941 8.3096 10.2404 9.5470 7.6987 5.9787 4.2730 3.3365 2.5572 1.4397 1.0166 0.9310 0.8627 0.4639

0.16 0.07 0.07 0.06 0.08 0.07 0.19 0.11 0.22 0.12 0.13 0.20 0.18 0.17 0.36

39

Ar (nA) + %

0.08157 0.29290 0.30069 0.36991 0.34588 0.27771 0.21579 0.15442 0.12033 0.09188 0.05179 0.03649 0.03335 0.03117 0.01677

0.21 0.21 0.14 0.16 0.27 0.15 0.46 0.25 0.73 0.25 0.28 0.37 0.42 0.46 1.00

38

Ar (nA) + %

1.001E-03 3.508E-03 3.698E-03 4.561E-03 4.164E-03 3.360E-03 2.656E-03 1.836E-03 1.408E-03 1.156E-03 6.246E-04 4.450E-04 4.007E-04 3.899E-04 2.212E-04

3.5 1.7 1.6 1.5 1.5 1.9 1.6 2.3 2.7 3.3 4.8 6.3 6.4 6.9 11.5

37

Ar (nA) + %

1.19E-03 3.42E-03 3.17E-03 4.02E-03 3.61E-03 3.03E-03 2.20E-03 1.57E-03 1.39E-03 9.42E-04 5.70E-04 2.53E-04 4.44E-04 1.27E-04 1.60E-04

22 8 9 7 8 9 12 17 18 28 45 101 59 202 159

38

Ar (nA) + %

3.2E-05 2.3E-06 1.2E-05 4.1E-05 1.8E-05 2.5E-05 2.4E-05 1.3E-05 2.5E-05 1.9E-05 5.9E-06 2.8E-06 1.5E-05 6.0E-06 1.9E-06

43 635 123 37 152 98 85 155 93 102 337 724 129 333 1067

40

Ar*/39ArK + % % 40Ar* 28.73 27.63 27.62 27.65 27.59 27.70 27.67 27.65 27.67 27.77 27.76 27.84 27.78 27.62 27.64

0.23 0.19 0.17 0.17 0.21 0.18 0.32 0.22 0.47 0.25 0.32 0.43 0.46 0.49 0.93

99.6 100.0 100.0 99.9 99.9 99.9 99.9 99.9 99.8 99.8 99.9 99.9 99.5 99.8 99.9

Age 248.1 + 0.9 239.2 + 0.8 239.1 + 0.7 239.4 + 0.7 238.8 + 0.9 239.7 + 0.7 239.5 + 1.3 239.3 + 0.9 239.5 + 1.9 240.3 + 1.0 240.3 + 1.3 240.9 + 1.7 240.4 + 1.8 239.1 + 2.0 239.2 + 3.8

Lab#: 32542. Sample: MSG09 sanidine, c. 5 mg, 100 –180 micron fraction. J: 0.005131 + 0.000005. Standard: Fish Canyon sanidine (28.02 Ma). Disc.: 1.00425 + 0.00125 per AMU (Power law; atmospheric 40Ar/36Ar ¼ 295.5). Heating device: CO2 laser with integrator lens. Notes: Relative abundances of Ar isotopes and their uncertainties in nanoamperes (nA) of amplified ion beam current, corrected for background, mass discrimination, and radioactive decay. Ages are calculated using the decay constants of Steiger & Ja¨ger (1977), and interference corrections of Renne et al. (2005) for Ca- and K-derived Ar isotopes, and Renne et al. (2008) for Cl-derived isotopes. Age uncertainties neglect systematic errors arising from 40K decay constants or the age of the standard. Uncertainties are given at the 2s level.

R. MUNDIL ET AL.

2.0 3.0 3.5 4.0 4.5 5.0 5.5 6.0 7.0 8.0 9.0 10.0 12.0 17.0 20.0

40

TRIASSIC RADIO –ISOTOPIC AGES

the Late Triassic regarding timescale calibration (Furin et al. 2006). A complementary 40Ar/39Ar age of 227.8 + 0.3 Ma in tetrapod bearing terrestrial sediments in Argentina (Rogers et al. 1993) may prove useful for the correlation of marine and terrestrial events (accounting for the 1% bias the latter age translates to U– Pb normalized age of 230 Ma). An U –Pb single zircon age from impact melt rock of the Manicouagan impact structure yields an age of 215.5 Ma (Ramezani et al. 2005), slightly older than but more robust than the result from an earlier study on multi-crystal samples (Hodych & Dunning 1992). The age is of little stratigraphic value but may prove useful for research aimed at studying potential effects of the impact event on land and in the ocean. The lack of geochronological data useful for calibration in the Late Triassic timescale may be mitigated by a combination of a few new radio – isotopic ages from marine sediments in combination with magnetostratigraphy (Muttoni et al. 2004), and correlation with the detailed terrestrial record of the Newark basin, which is convincingly calibrated by orbital cycles (Kent et al. 1995; Olsen & Kent 1996). However, constraints from radio-isotopic ages are urgently needed in order to scale the Late Triassic, large portions of which are uncalibrated (Fig. 2). The end of the Triassic is again marked by a substantial extinction event as well as a CFB event of similar magnitude as the Siberian Traps, namely the Central Atlantic Magmatic Province (CAMP).

Age of the Triassic – Jurassic boundary There is as yet no ratified GSSP for the base of the Jurassic, although traditionally the lowest occurrence of the Psiloceras ammonite genus is considered a useful marker. An age of 199.6 + 0.3 Ma (Pa´lfy et al. 2000) for an ash bed within marine sediments just below the Triassic –Jurassic boundary in the Queen Charlotte Islands has to be considered a minimum age since the analyses were performed on multi-crystal samples of air abraded zircons. Suspicions that the latter age may be too young were confirmed by U –Pb single zircon analyses from ash layers within the Hettangian and Sinemurian which yielded an extrapolated age of the boundary of .201 Ma (Mundil & Pa´lfy 2005; Mundil et al. 2005; Pa´lfy & Mundil 2006). This is now confirmed by a recent study with an U –Pb single zircon age of 201.6 + 0.2 Ma from an ash layer situated between last and first occurrences of the boundary-defining Choristoceras and Psiloceras ammonites in Peru (Schaltegger et al. 2008). As in Ovtcharova et al. (2006), the uncertainty associated with tracer calibration has been applied to individual ages rather

51

than the final weighted mean age, which means that the accuracy has been (slightly, in this case) overstated.

Age of the Central Atlantic magmatic province There is a longstanding debate about a causal link between CAMP volcanism and the Triassic –Jurassic extinction summarized in Knight et al. (2004), also see Whiteside et al. (2007) and Marzoli et al. (2008). Although radio–isotopic ages for the basalts are of marginal value for the purpose of timescale calibration, they will be discussed here since they are essential for the correlation of events in terrestrial and marine environments. Detailed geochronological research facilitated the recognition that volcanic products related to the CAMP are widespread and now crop out on four continents (Marzoli et al. 1999; Hames et al. 2000; Marzoli et al. 2004) and yield new 40 Ar/39Ar ages, summarized in Knight et al. (2004). The latter study concludes that the peak of the volcanic activity occurred at an 40Ar/39Ar age of just before 200 Ma, which translated to a U –Pb normalized age of slightly younger than 202 Ma. Only few U –Pb ages exist for the products of CAMP volcanism (for reasons discussed above). U –Pb analyses of multi-crystal analyses from the Palisades and Gettysburg sill both yielded ages of c. 201 Ma, although both data sets are affected by considerable scatter (Dunning & Hodych 1990) and almost certainly suffer from Pb-loss effects. A recent study by Schoene et al. (2006) yielded an extremely precise age of 201.27 + 0.03 Ma for zircons from the North Mountain basalt, the lowest flow unit of CAMP within the Newark Supergroup in the Fundy Basin of eastern North America. The U –Pb normalized 40Ar/39Ar ages in combination with the U –Pb ages show that the age of the extinction recorded in marine sediments and those of CAMP volcanism are indistinguishable (at the permil level), suggesting a causal link (Pa´lfy 2003).

Summary and conclusions With the exception of large portions of the Late Triassic, the geochronological database for the calibration of the Triassic timescale is now substantially improved compared to previous compilations. Some of the data discussed herein were not yet available for the most recent published compilation (GTS 2004; Gradstein et al. 2004; Ogg 2004) but it is evident that every single age for stage boundaries suggested here is different from the ages given in the GTS 2004. The ages for stage boundaries are listed in Table 3. Here, we only used ages from

52

Table 3. Isotopic ages for Triassic stage boundaries 206 Pb/238U age(a)

+ 2s

(b)

(c)

40

Ar/39Ar age

Normalized 40 Ar/39Ar age(d)

+ 2s(e)

201.6 , 230 +0.2

+0.3

230.9 , 236 +0.1 +0.1 +0.3 239.3

+0.2 +0.4 +0.5

242.0 242.1

+0.6 +0.8 +0.9

242.8

+0.6 +0.4 +0.5

244.6

+0.5

246.3

+0.1 +0.1 +0.3

246.8

+0.4

246.8

+0.1 +0.2 +0.3

247.1

+0.1 +0.2 +0.3

247.2

+0.6

+0.5

239.5

241.9

+0.5

Rhaetian –Hettangian (Triassic – Jurassic) Uppermost Rhaetian (Utcubamba Valley, Peru) Upper Carnian (Pignola 2, Italy) Ladinian –Carnian Basal upper Ladinian (Flexenpass, Austria) Anisian – Ladinian Uppermost Anisian (M. San Giorgio, Switzerland) Uppermost Anisian (Latemar, Italy) Upper middle Anisian (Jinya, China) Middle Anisian (Guandao, China) Middle Anisian (Jinya, China) Lower Anisian (Guandao, China) Lowermost Anisian (Guandao, China) Olenekian – Anisian

Mineral

Pretreatment (f)

Reference

Zircon

CA

Schaltegger et al. 2008

Zircon

CA

Furin et al. 2006

Zircon

CA

Bru¨hwiler et al. 2007

Zircon, sanidine Zircon

CA

This study

CA

Brack et al. 2007

Zircon

CA and air abrasion Ovtcharova et al. 2006

Zircon

CA and air abrasion Ramezani et al. 2007

Zircon

CA and air abrasion Ovtcharova et al. 2006

Zircon

CA and air abrasion Lehrmann et al. 2006

Zircon

CA and air abrasion Lehrmann et al. 2006

R. MUNDIL ET AL.

201.5

Stage (or boundary) and locality

247.3

+0.1 +0.1 +0.3

247.4

+0.1 +0.1 +0.3

248.1

+0.4

+0.5

250.6

+0.5

+0.6

251.2

+0.2

+0.2

251.3 252.3 +0.3 +0.6 +0.6

252.2

+0.4 +0.7 +0.7

252.5

+0.3 +0.6 +0.6

(a)

Weighted mean 206Pb/238U age and interpolated (or estimated) age for the stage boundary (in Ma). 2s error including uncertainty on tracer calibration. (c) 2s error including uncertainty on tracer calibration and l238U. (d) Normalized to U –Pb decay according to Mundil et al. (2006). (e) Not including uncertanties on l40K and the age of the standard. (f) Zircons subjected to annealing/chemical abrasion (CA –TIMS, Mattinson 2005) or air abrasion (Krogh 1982). (b)

Zircon

CA and air abrasion Lehrmann et al. 2006

Zircon

CA and air abrasion Lehrmann et al. 2006

Zircon

CA and air abrasion Ovtcharova et al. 2006

Zircon

CA and air abrasion Ovtcharova et al. 2006

Zircon

CA

Galfetti et al. 2007

Zircon

CA

Mundil et al. 2004

Zircon

CA

Mundil et al. 2004

Zircon

CA

Mundil et al. 2004

TRIASSIC RADIO –ISOTOPIC AGES

252.4

Uppermost Olenekian (Guandao, China) Uppermost Olenekian (Guandao, China) Uppermost Olenekian (Jinya, China) Upper Olenekian (Jinya, China) Lower Olenekian (Jinya/ Waili, China) Induan – Olenekian Changsingian – Induan (Permian – Triassic) Uppermost Changsingian (bed 25 Meishan, China) Uppermost Changsingian (Shangsi, China) Uppermost Changsingian (Shangsi, China)

53

54

R. MUNDIL ET AL.

biostratigraphically-controlled volcanic deposits. The difference from the GTS 2004 appears puzzling but understandable in view of the conflicting information prior to its publication. A good example is the Middle Triassic, where relatively robust radio-isotopic ages from different studies (Mundil et al. 1996c; Pa´lfy et al. 2003) were available but not in agreement with a previous study (Hellmann & Lippolt 1981) and also conflicting with timescales derived from purported orbital cycles recorded in platform carbonates of the Dolomites (Hinnov & Goldhammer 1991; Goldhammer et al. 1993; Preto et al. 2001). Although the existence of the Latemar cycles is not disputed by research groups opposing this hypothesis, there is strong evidence that the periodicities are very different and much shorter than originally suggested (Mundil et al. 2003; Zu¨hlke et al. 2003). This conclusion is not based on radio-isotopic ages alone but also on strong evidence from magnetostratigraphic and biostratigraphic correlation (Brack & Rieber 1993; Brack et al. 1996; Kent et al. 2004; Hinnov 2006; Kent et al. 2006). The use of unrecognized inaccurate but precise data for the GTS 2004 (Gradstein et al. 2004) reveals many of the pitfalls associated with the use of high-resolution geochronology in that the uncertainties assigned to the ages of stage boundaries are up to an order of magnitude smaller than in previous compilations, but after a careful review all of the chosen ages fail to stand up to scrutiny for reasons given above. It is for this reason that we feel prompted to offer an updated review, but it is crucial to realize that this compilation is only a snapshot in time which will inevitably undergo further refinement. It is more important to realize that radio – isotopic ages with permil or now even sub-permil level uncertainties may often be inaccurate, and the uncertainties only reflect the mere analytical uncertainties (or in some cases, the results of subjective data-culling), which may not have any bearing on ‘geological’ uncertainties. For example, the time between the crystallization of a dated mineral (the time when its radio –isotopic clock was set to zero) and the time of its deposition may well exceed some of the uncertainties quoted above. This can either be due to time lag before the volcanic layer is deposited or a protracted time of crystallization of minerals (particularly zircons) in the magmatic system before they erupt in volcanic events. There is now abundant evidence that this magma residence time for zircons may be as long as 0.2 to 0.3 Ma, as demonstrated on zircons from young volcanic systems (e.g. Simon et al. 2008 and references therein). For the time being, quantifying the time lag for each individual eruption remains elusive, and it is therefore impossible to correct for this additional complication except by some ad hoc means as

suggested by Simon et al. (2008). It illustrates, however, how meaningful (or meaningless) uncertainties of 0.1 Ma are, which are now often assigned to individual radio–isotopic ages. The ages compiled in Table 3 and Figure 2 include uncertainties on tracer calibration that are typically not in excess of c. 0.1 permil and therefore negligible. Recent efforts to produce, calibrate and distribute tracer solutions that can then uniformly be used by all laboratories are underway and aimed at eliminating this bias (see e.g. Bowring et al. 2005; Condon & Members of the Earthtime U –Pb working group 2005). The ages chosen here have passed strict selection criteria: † They are from primary volcanic products (i.e. there is no evidence of re-deposition) within marine sedimentary sections containing age diagnostic fossils. Good indicators for primary products are the mineral inventory and the lack of features indicating redeposition (e.g. rounded crystals). Ash falls are the results of explosive volcanism from evolved volcanic systems characterized by a felsic mineral assemblage, which usually contain U –bearing (zircon) as well as K –bearing (sanidine, biotite) minerals. The presence of accretionary lapilli is rare but a good indicator of airborne ash falls. Rhyolithic flows containing accessory zircons are also prime targets for both U –Pb and K –Ar (40Ar/39Ar) analyses, as are basaltic flows for K –Ar (40Ar/39Ar) geochronology. Recently, the use of U– Pb analyses applied baddeleyite contained in basaltic volcanics has become more common. † Most of the analyses were performed on individual CA pre-treated zircon crystals resulting in coherent clusters of individual ages with no appreciable indication of open system behaviour (although in some cases analyses on air abraded crystals were included, which may bias the result towards a slightly younger age). Since the ages chosen for the compilation are exclusively based on the U –Pb system, uncertainties arising from the calibration of decay constants (here l238U) can be ignored (the effect is relatively small as the uranium decay constants are known at the 0.1% level: Jaffey et al. 1971). There is some ambiguity about the accuracy of l235U, however, all the ages presented here are from the 206 Pb/238U ages and therefore not affected by this problem (Mattinson 2000; Schoene et al. 2006). Uncertainties on decay constants are included in Table 3 for comparison with ages from different isotopic systems. In general, uncertainties on decay constants should be included since they significantly exceed the level of achievable analytical precision in

TRIASSIC RADIO –ISOTOPIC AGES

arguably most isotopic systems. The currently used decay constants are from a three-decade-old compilation (Steiger & Ja¨ger 1977), which was very useful and provided a common base, effectively ending an era where every group doing geochronological research used different decay constants for the same isotopic system. There is now solid evidence, however, that many of the constants given in Steiger & Ja¨ger (1977) no longer stand up to scrutiny. This is particularly true for the 40K decay constant as already discussed above (yielding 40Ar/39Ar ages, which are too young by c. 1% for Phanerozoic times), but also for example for l87Rb, which may be miscalibrated by an even greater amount (Davis et al. 1977; Rotenberg et al. 2005; Nebel et al. 2006). Two points from arising systematic and/or random biases stand out and must always be considered when any high-resolution radio – isotopic ages are used: † Inaccurate ages arising from systematic biases such as miscalibrated tracer solutions and decay constants can be corrected subsequently, once the corrected parameters are known with more confidence. † This in not true for inaccuracies arising from open system behaviour, that is, the loss or gain of the parent isotopic and/or its radiogenic daughter isotope, since the magnitude of the loss (or gain) is unknown. The effects from the latter are typically small but not insignificant and often exceed the quoted uncertainty. The inaccurate and younger U –Pb ages on zircons for sample MSG.09 from an older study (Mundil et al. 1996c) are a good example (Fig. 1a) to illustrate this complication. Pb loss is ubiquitous in zircons and typically correlates with magnitude of radiation damage within crystals. Numerous examples for both cases exist in the literature, and many are cited in this contribution. It is therefore always important to scrutinize the data from the original publication rather than taking the ages from timescale compilations at face value. For this task, a publicly available software toolkit (Isoplot) for the evaluation and calculation of geochronological data is of immense value (Ludwig 2003). If data sets are coherent and do not require the subjective rejection of individual data, the calculation of the weighted mean age from those individual data is straightforward. This is exemplified by zircon analyses on MSG.09 in Figure 1, where the individual analyses in Figure 1a are clearly affected by Pb loss and an age may only be extracted after about half of the data are excluded. The striking opposite is displayed in Figure 1b, where no rejection is necessary. Further careful consideration is necessary when it comes to using radio – isotopic ages as constraints

55

of geological processes, such as biotic change. There is increasing awareness that reliance on individual fossil markers (such as the definition for GSSPs) may not be as useful for reliable correlations since they may be diachronous, even if the magnitude of this diachronism is small. The increasing precision of radio–isotopic ages magnifies this problem so that alternative approaches for the evaluation of origination and extinction have to be considered. For example, more quantitative approaches such as developed by Guex (1991) or Sadler (2004), using computer algorithms that compensate for inherently incomplete biostratigraphies, can build high-resolution timescales by sequencing and calibrating numerous fossil ranges from many different localities with a more robust result than from a single fossil occurrence from a single site. Applying this approach is beyond the scope of this contribution, but future efforts may benefit from the improved timescale presented here. Finally it has to be mentioned that the application of promising new isotopic dating techniques, in addition to the ones discussed here, such as Re –Os dating of black shales (Hannah et al. 2004; Creaser et al. 2005; Selby & Creaser 2005) and U –Pb dating of paleosols and carbonates (Rasbury et al. 1998), may provide additional valuable data for timescale calibration where traditional dating techniques fail, for example due to the lack of volcanic products in the Upper Triassic. The manuscript benefitted from thoughtful reviews from Jim Ogg, Jean Guex and Hugo Bucher. Part of the research presented and reviewed here was supported by NSF grants 0125799 and 0617733 to RM and PRR. Analytical work at the Berkeley Geochronology Center was supported by the Ann and Gordon Getty Foundation.

References B ACHMANN , O., O BERLI , F., D UNGAN , M. A., M EIER , M., M UNDIL , R. & F ISCHER , H. 2007. 40Ar/39Ar and U–Pb dating of the Fish Canyon magmatic system, San Juan Volcanic field, Colorado: evidence for an extended crystallization history. Chemical Geology, 236, 134– 166. B AKSI , A. K. & F ARRAR , E. 1991. 40Ar/39Ar dating of whole-rock basalts (Siberian Traps) in the Tunguska and Noril’sk areas, USSR. Eos Transactions American Geophysical Union, 72, 570. B LACK , L. P., K AMO , S. L., W ILLIAMS , I. S., M UNDIL , R., D AVIS , D. W., K ORSCH , R. J. & F OUDOULIS , C. 2003. The application of SHRIMP to Phanerozoic geochronology: a critical appraisal of four zircon standards. Chemical Geology, 200, 171–188. B ORSI , S. & F ERRARA , G. 1968a. Determinazione dell’eta delle rocce intrusive di predazzo con i metodi del rb. Mineralogica et Petrographica Acta, 14, 171 –183.

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Y IN , H. F., Z HANG , K. X., T ONG , J. N., Y ANG , Z. Y. & W U , S. B. 2001. The Global Stratotype Section and Point (GSSP) of the Permian–Triassic Boundary. Episodes, 24, 102– 114. Z U¨ HLKE , R., B ECHSTADT , T. & M UNDIL , R. 2003. SubMilankovitch and Milankovitch forcing on a model Mesozoic carbonate platform – the Latemar (Middle Triassic, Italy). Terra Nova, 15, 69– 80.

The geomagnetic polarity timescale for the Triassic: linkage to stage boundary definitions MARK W. HOUNSLOW1* & GIOVANNI MUTTONI2 1

Centre for Environmental Magnetism and Palaeomagnetism, Lancaster Environment Centre, Lancaster University, Bailrigg, Lancaster, UK LA1 4YQ

2

Dipartimento di Scienze della Terra, Universita` di Milano, Via Magiagalli 34, I-20133 Milan, Italy (e-mail: [email protected]) *Corresponding author (e-mail: [email protected]) Abstract: Studies of Triassic magnetostratigraphy began in the 1960s, with focus on poorly fossilferous nonmarine red-beds. Construction of the Triassic geomagnetic polarity timescale was not consolidated until the 1990s, when access to magnetometers of sufficient sensitivity became widely available to measure specimens from marine successions. The biostratigraphicallycalibrated magnetostratigraphy for the Lower Triassic is currently largely based on ammonoid zonations from Boreal successions. Exceptions are the Permian–Triassic and Olenekian– Anisian boundaries, which have more extensive magnetostratigraphic studies calibrated by conodont zonations. Extensive magnetostratigraphic studies of nonmarine Lower Triassic successions allow a validation and cross-calibration of the marine-based ages into some nonmarine successions. The Middle Triassic magnetostratigraphic timescale is strongly age-constrained by conodont and ammonoid zonations from multiple Tethyan carbonate successions, the conclusions of which are supported by detailed work on several nonmarine Anisian successions. The mid Carnian is the only extensive interval in the Triassic in which biostratigraphic-based age calibration of the magnetostratigraphy is not well resolved. Problems remain with the Norian and early Rhaetian in properly constraining the magnetostratigraphic correlation between the well-validated nonmarine successions, such as the Newark Supergroup, and the marine-section-based polarity timescale. The highest time-resolution available from magnetozone correlations should be about 20–30 ka, with an average magnetozone duration of c. 240 ka, for the Lower and Middle Triassic, and about twice this for the Upper Triassic.

The early pioneering work of Brunhes (1906) and Matuyama (1929) recognized that volcanic rocks recorded magnetization directions similar to the orientation of the present day Earth’s magnetic field (i.e. of normal polarity), but also that some volcanic rocks recorded older magnetization directions that were in the opposite direction (i.e. of reverse polarity). Motonori Matuyama was the first to suggest that these directions recorded the reversal in the main (i.e. dipole) component of the Earth’s magnetic field (see discussion of early developments in Jacobs 1963). The first studies on the natural remanent magnetization and magnetic properties of sedimentary rocks were conducted in the late 1930s and 1940s, often with the focus on Pleistocene continental sediments (e.g. McNish & Johnson 1938; Ising 1942; Nagata 1945; Graham 1949; Torreson et al. 1949). In the 1950s, more comprehensive work on Neogene volcanic rocks showed a consistent stratigraphic pattern in the recorded polarity of magnetizations, that is, a magnetostratigraphy (see Irving 1964; Hailwood 1989; and McElhinney & McFadden 2000 for a review of these early developments). Palaeomagnetic data from Triassic red-bed sediments were first published by Clegg et al. (1954) and

Creer et al. (1954). The later authors also undertook the first published magnetostratigraphic study, focussing on the Late Proterozoic from the UK. Palaeomagnetic work on other Triassic successions, from the USA, quickly followed (Graham 1955; Runcorn 1955; Du Bois 1957), demonstrating that other sediments also recorded magnetizations of both reverse and normal polarity. Radiometric evidence, providing convincing support that the Earth’s magnetic field polarity changes were synchronous on a global scale, was firmly established in the early 1960s (Irving 1964; see review in McElhinney & McFadden 2000), which Vine & Matthews (1963) used in their sea-floor spreading model, linking Earth’s magnetic field polarity changes with sea-floor magnetic anomaly lineations.

Roots of a Triassic Geomagnetic Polarity Timescale (GPTS) Most of the palaeomagnetic work in the 1950s and 1960s was directed to providing data to support the concepts of continental drift (Irving 1964). Work by Creer (1958, 1959) on part of the UK Triassic

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 61– 102. DOI: 10.1144/SP334.4 0305-8719/10/$15.00 # The Geological Society of London 2010.

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was the earliest Triassic palaeomagnetic study that placed a set of palaeomagnetic samples into stratigraphic order to produce a simple magnetostratigraphy. The early pioneer in the development of magnetostratigraphy for stratigraphic correlation was A. N. Khramov, who published the seminal summary of on-going Russian work in 1958 (Khramov 1958). This was primarily focused on extensive studies of Neogene and Quaternary successions in western Turkmenistan (Cheleken Peninsula), but was also significant (Irving 1964; Glen 1982) in that it discussed fundamental magnetostratigraphic concepts, such as the use of multiple sections, minimum sampling requirements to define magnetozones, and palaeomagnetic data quality. It also anticipated the construction of a geomagnetic polarity timescale (GPTS) for dating and correlation. In addition, Khramov (1958) outlined a rudimentary working knowledge (without details) of the magnetostratigraphy from Upper Permian and Lower Triassic sections in the Vyatka River region of the Moscow Basin. Details of this multiplesection magnetostratigraphic study appeared subsequently (Khramov 1963), and it was quickly followed by studies on the Chugwater Formation in the western USA (Picard 1964) and the German Upper Buntsandstein (Burek 1967, 1970). The focus of Triassic magnetostratigraphic studies in the 1950s to early 1980s was on terrestrial red-bed successions, since these provided natural remanent magnetizations that could be easily measured on the early astatic magnetometers and the later fluxgate spinner magnetometers then available (Collinson et al. 1957; Gough 1964). During this period, the development of routine magnetic cleaning techniques, referred to as demagnetization (As & Zijderveld 1958; Creer 1959; Wilson 1961) and more rigorous analysis (i.e. using least-square best fitting methods: Kirschvink 1980) of palaeomagnetic data, were developed into methodologies that are routinely used today. The widespread use of full demagnetization techniques, now accepted as standard for extracting primary magnetizations, was only fully embraced in the 1970s. As such there is some scepticism about the validity of palaeomagnetic data generated prior to the 1970s, which is in part expressed in the quality criteria suggested by Opdyke and Channell (1996) for classifying magnetostratigraphic data. It was not until the development of superconducting quantum interference device (SQUID) magnetometers in the 1970s (Goree & Fuller 1976), and their widespread use since the 1980s and 1990s, that the weakly magnetic specimens found in many marine Triassic successions could be suitably measured, demagnetized, and primary magnetization components extracted. This development

finally heralded the expansion of detailed studies on the construction of a Triassic GPTS, after earlier attempts to apply the new instrumentation to limestone successions of other ages (e.g. Martin 1975; Heller 1977). The seminal magnetostratigraphic works of Lowrie & Alvarez (1977) and Channell et al. (1979) on Cretaceous limestones of the Apennines and the Southern Alps, respectively, were influential and were followed by work on the Triassic by Heller et al. (1988), McFadden et al. (1988) and Steiner et al. (1989), who provided the first detailed magnetostratigraphic studies of Triassic carbonates, in predominantly marine successions.

Early developments of the Triassic GPTS The first attempts at the construction of a Triassic GPTS through the 1960s and 1970s were inevitably fragmentary, being based around nonmarine successions, which were often imprecisely dated by vertebrates and palynomorphs. Khramov (1963) was the first to attempt the construction of a Lower and Middle Triassic GPTS, based on the Vetluga successions from the Moscow Basin and existing studies from the western literature. This was later followed by attempts at a complete Triassic GPTS by McElhinney & Burek (1971), Pergament et al. (1971), Pechersky & Khramov (1973), and Molostovsky et al. (1976). In spite of the rapid development of the GPTS for the latest Jurassic to Pleistocene, mainly through study of sea-floor linear magnetic anomalies, the absence of Triassic sea-floor largely impeded the development of a detailed Triassic GPTS until the widespread availability of SQUID magnetometers in the late 1980s. A feature that also characterizes most of the Triassic magnetostratigraphic studies and GPTS construction prior to the 1990s is the common lack of true integration with detailed biostratigraphies provided by the co-study of, for example, ammonoids and conodonts. It is the initial expansion of such integration in the early 1990s with studies such as Ogg & Steiner (1991) using ammonoids, and Gallet et al. (1992, 1993) using conodonts, that the Triassic GPTS has now been developed into such detail. In the last two decades there has been much progress, particularly in calibrating the pattern of reverse and normal magnetic field polarity changes against conodont biostratigraphies (Muttoni et al. 1996a, 2000, 2004; Gallet et al. 1998, 2000a, 2007; Channell et al. 2003).

The time resolution of magnetostratigraphic correlation Correlation using magnetostratigraphic principles is at two scales. Firstly, the changing pattern of

GEOMAGNETIC POLARITY TIMESCALE

magnetic polarity (i.e. magnetozones) over a stratigraphic interval can provide a distinctive ‘bar-code’ pattern for correlation. This is because magnetic field reversal is essentially a stochastic process, giving random length-durations of magnetozones (McElhinney & McFadden 2000; Lowrie & Kent 2004). The longer the fragment of the polarity bar-code, and the more constraints from other stratigraphic tools, the greater is the confidence in intersection correlation. For the Cenozoic, the maximum resolution of the GPTS is about 20 –30 ka, with reversals on average every c. 0.22 Ma (McElhinney & McFadden 2000; Lowrie & Kent 2004). The Late Triassic appears to have a reversal rate somewhat similar to the Cenozoic, with a maximum magnetozone resolution of about 30 ka (Kent et al. 1995; Kent & Olsen 1999). Secondly, correlation of the boundaries (transitions) of magnetozones provides the highest resolution of correlation. Studies on the Cenozoic suggest that time durations of magnetic field polarity transitions from reverse to normal (or vice-versa) are between 1000 to 8000 years, probably varying depending upon location and the actual magnetozone transition (McElhinney & McFadden 2000). In the Brunhes magnetochron (i.e. the normal polarity interval since 0.78 Ma: Cande & Kent 1995), the briefest evidence of pre-emptive polarity changes are geomagnetic excursions, which have a duration of less than 10,000 years (Langereis et al. 1997). It is probable that similar excursions existed in the Triassic, but without cm-scale studies in successions with high sedimentation rates, the prospect of using such excursions for correlation in the Triassic is remote. The time resolution provided by magnetostratigraphic correlation is also bound up with sampling density issues, sedimentation rates and the presence of disconformities. The highest resolution is achievable from continuously deposited and expanded successions with high sedimentation rates, sampled at the smallest stratigraphic interval. For this reason, magnetostratigraphic studies can have site-selection priorities contrary to those of biostratigraphic studies, which may focus on condensed successions with high fossil recovery rates.

Stratigraphic principles of magnetostratigraphic correlation Correlation using magnetostratigraphic normal/ reverse polarity bar-code patterns relies on a number of factors related to the preserved stratigraphy and its sampling: (a) A reasonable within-section consistency of sedimentation rate is advantageous to maintain the relative stratal (and time duration) thicknesses of magnetozones through the section.

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Given a detailed basin-wide sequence stratigraphy, it may be possible to estimate sedimentation rate distortions using sequence stratigraphic principles. (b) Stratigraphic gaps can distort the magnetostratigraphic pattern, unless properly identified using biostratigraphic data, supported by appropriate sedimentology and sequence stratigraphic studies. (c) Sections representing longer periods of time stand a better chance of providing convincing bar-code matches. (d) Sampling resolution should be matched with sedimentation rate, and the expected number of polarity changes in a section. If sampling resolution is low, polarity changes are frequent, and sedimentation rate is also low and variable, then the recovered pattern may be a poor match to the ‘real’ polarity pattern. For this reason, it is not good practice to count magnetozones for correlation purposes, whereas it is much more reliable to use the dominance of polarity as a means for correlation (or use other correlation constraints), since this is not as strongly affected by sampling density and changes in sedimentation rate. Comparison between sections with large betweensection sedimentation rate differences is best accommodated by stretching or shrinking the entire magnetostratigraphic height scale linearly, using as a guide, biostratigraphic or radiometric correlation constraints in addition to the magnetostratigraphy. Using such scaling, composite timescales can be constructed in a ‘pseudo-height’ scale, based on principles of graphic correlation (Shaw 1964; Pa¨like et al. 2005). We here use these principles to develop the Triassic GPTS using marine sections, because with such sections there is greater chance of proving continuity using other correlation constraints. The major ‘anchor points’ for these composites are indicated on the diagrams (e.g. Fig. 1). We then examine the higher detail sometimes available from nonmarine sections (e.g. Newark Supergroup in the Upper Triassic). We focus primarily on the stage boundaries, because these are the only real fixed points in Triassic time, and this approach is compatible with other contributions in this book. The magnetochron couplets (i.e. each successive N –R pair) in the GPTS are labelled LT, MT and UT for the Lower, Middle and Upper Triassic, respectively.

The magnetostratigraphy of the Permian – Triassic transition The Permian–Triassic boundary (PTB) is located at the first occurrence (FO) of the conodont Hindeodus parvus, in the global stratotype section and point

Normal polarity Reverse polarity

PTB extinction event (t=terrestrial, m=marine marker)

Uncertain polarity

δ C Isotope positive peak

Sampling gap

δ13C Isotope negative peak (o=organic matter, c=carbonate)

Magneto correlations

13

@ = Correlation and scaling anchor

Tetrapod range

Vikinghøgda (Spitsbergen, arctic Norway)

Miospore assemblage

Hiatus

Tetrapod assemblage

Condensed

Range extends beyond displayed section

Olenekian

Amm. Zones

mune

Shangsi

LT1

O

.b

or

ea

le

n.1r

Dienerian

Induan

O. com

O. concavum m

P5

Dorashamian

A. = Arctoceras E. = Euflemingites H. = Hedenstroemia O.= Otoceras V. = Vavilovites

senk ran striga tzi/ tus

Griesbachian

Bulla & Suisi

Ammonoid genera

Radiometric age

Magnetochrons

Scaling section Hechuan

A.blomstrandi Ns. svalbardensis

Konickites

Vh1r

O. boreale B. rosenkrantzi Ng. carinata

Vh3

Deltadalen Mb

Ng. meishanensis, Ng. taylorae

Vh6 Vh5 Vh4

Lusitaniadalen Mb

E. cirratus

H. hedenstroemi

B. strigatus

Smith Creek Mb

V. sverdrupi

Claraia

H. praeparvus H. typicalis

O. boreale

I. isarcica I. staeschei

H. aequabilis Zone

Confederation Pt. Mb

50 m

Palynomorph range Bivalve range

H. parvus

Mazzin Mb

Ammonoid horizon/ range Conodont horizon / range

B. ro

dus

Permian

ha

Conodont genera H. = Hindeodus I. = Isarcicella Ng. = Neogondolella Ns. = Neospathodus P. = Pachycladina

t

i-

ndi

Changhsingian

ng

hs

25 m

in

Bed 11

O. concavum

rup

P. c a

Dzhulfian

15

11 10

Vh2

erd

‘Chang-N’

16

t,m

o

c O.

V. s v

LT2

Steiner (2006)

18

c

ne

u mm

LT3

20 m

e

t

1

?

@

Bulla Mbr

Zon T-S

@ Dzhulfian Dorashamian

c m

29 28

25 m

P. obliqua Zone

Seis Mb

31

Bellerophon Fm T

LT1n.1r

21

Mag.zone (GC)

Bulla & Siusi (Italy)

Abadeh (Iran) 5m

@

33

P - P Zone

2

Andrez Mb

Ophiceras

Ophiceras

I. turgida

35

13

gi

Proptychites

Ns. dieneri

H. parvus

H. parvus

Claraia

an

37

36

3

H. anceps Zone

C

Conodont zones

38

I. staeschei

H. parvus

F1

Changxing Fm

15

42

40

17

16

Ellisonia sp.

Feixianguan Fm

F2

46 44

Hambast Fm

NC

20

252.5 Ma

LM

22

52

49

25 23

Bed

Bed 52

Campil Mb

& H. typicalis

100 m

Shuijiang

20 m

Bed

Dalong Fm Feixianguan Fm

c

CS

Talus breccia -no sampling

34

Changxing Fm

Shale

m

H. parvus

10 m Yingkeng Fm

Beds

Ns. dieneri H. parvus

Ng. changxingensis

m

LT1n.1r

F3

Shangsi composite

4

GPTS

fault Major provenance change

Fig. 1. Summary of the bio-magnetostratigraphy across the Permian –Triassic boundary. Section data from left to right: Lower Guandao (Payne et al. 2004; Lehrmann et al. 2006); Meishan composite (Li & Wang 1989; Yin et al. 2001; Zhao et al. 2007); Shuijiang (Chen et al. 1994; Heller et al. 1995); Hechuan (Steiner et al. 1989); Shangsi composite compiled from Heller et al. (1988), Steiner et al. (1989), Glen et al. (2009), Lai et al. (1996), Nicoll et al. (2002) and Wignall et al. (1995); Abadeh (Gallet et al. 2000b); Bulla & Siusi (Perri & Spalletta 1998; Scholger et al. 2000; Perri & Farabegoli 2003; Horacek et al. 2007); Griesbach Creek (Ogg & Steiner 1991; Henderson & Baud 1997; Hounslow et al. 2008a); Vikinghøgda (Hounslow et al. 2008a). South China conodont zones: CS, Neogondolella (Clarkina) subcarinata; NC, Ng. changxingensis yini – Ng. changxingensis; LM, Hindeodus latidentatus – Ng. meishanensis. Shangsi ammonoid zones: T– S, Tapashanites– Shevyrevites assemblage Zone; P–P, Pseudotirolites – Pleuronodoceras assemblage Zone. Conodont genus abbreviations: Ns, Neospathodus; Ng, Neogondolella. Thickness scales different for each section. Magnetozone width in the GPTS and section columns corresponds to data robustness and degree of confirmation from stratigraphically adjacent magnetostratigraphic sampling.

M. W. HOUNSLOW & G. MUTTONI

c

36

H. typicalis Ng. meishanensis I. staeschei Ng. tulongensis Ng. discreta

Meishan composite

Feixianguan Fm

F4

30 m

Ns. cristagalli; Ns. pakistanensis

Lower Guandao

South China Hechuan

64

Griesbach Creek, Axel Heiburg Is. (Canada)

GEOMAGNETIC POLARITY TIMESCALE

(GSSP) at Meishan, China (Yin et al. 2001) (Fig. 1). This FO follows two earlier key events in the latest Permian (i.e. in the latest Changhsingian), firstly, the marine extinction event, then slightly younger, a negative peak in d13Ccarb (Yin et al. 2001; Mundil et al. 2004). At all other sites, the location of the PTB is based on correlation to the Meishan GSSP, using conodont, carbon isotopic, sequence stratigraphic, palynological or magnetostratigraphic data, etc. At the Shangsi section in China, the marine extinction event is at the boundary of the Dalong and Feixianguan formations (Wignall et al. 1995), the negative d13Ccarb peak is between c. 1 to 5 m higher (see discussion in Mundil et al. 2004), and the first H. parvus is 4.5 m above the extinction event (Nicoll et al. 2002). At Meishan, the marine extinction event, the associated d13Ccarb negative peak and the FO of H. parvus are all within a stratigraphic range of about 0.3 m (Jin et al. 2000; Yin et al. 2001, 2005). Similarly, sections in Greenland show that the d13Corg negative excursion is slightly younger than an initial palynofloral turnover (which is a proxy for the extinction event) to assemblages that contain miospores typical of the Triassic (Looy et al. 2001). In the Karoo Basin, the peak of tetrapod extinctions is synchronous with the d13Corg negative excursion (Ward et al. 2005). Studies of d13Ccarb over this transition in the Alps are contradictory with respect to the position of the initial isotopic decline, which is either in the upper part of the Bellerophon Fm. (Magaritz et al. 1988; Sephton et al. 2005), or in the basal Werfen Fm. (Holser et al. 1989), and reaches a peak at either 10– 15 m or c. 25 m above the base of the Werfen Fm. In both cases, the most negative d13Ccarb is younger than the FO of H. parvus, which is at odds with data from other marine sections. The situation is more problematic when locating the polarity boundaries with respect to these events. In most sections where there is evidence of the latest Permian (such as at Shangsi and Guandao in China; Steiner et al. 1989; Lehrmann et al. 2006), and Abadeh in Iran (Gallet et al. 2000b), it is characterized by reverse polarity, which extends to include the late Permian extinction event itself (Glen et al. 2009). In the southern Italian Alps, the terrestrial extinction horizon is also located in the reverse polarity Bulla Member (Mb) (uppermost part of the Bellerophon Fm.; Scholger et al. 2000; Perri & Farabegoli 2003) about 0.5 m below the top (Cirilli et al. 1998), where there is a loss of typical late Permian miospores (e.g. Klausipollenites schaubergeri, Jugasporites delsaucei, Nuskoisporites dulhuntyi, Paravesicaspora splendens) and the introduction of forms such as Densoisporites playfordi, D. nejburgii, Convolutispora sp., and Rewanispora vermiculata. This same level also shows evidence of massive soil erosion (Sephton et al. 2005).

65

Sections such as Shuijang and Hechuan in China show a proxy for the extinction event (base of Feixanguan Fm.), as very near to the base of a normal polarity magnetozone. The magnetostratigraphy for the Meishan GSSP is anomalous with respect to other sections, in that it shows the extinction event in the middle parts of a normal magnetozone, which begins in the upper part of the Changhsingian. A reverse magnetozone in Meishan bed 27, spanning the PTB (Yin et al. 2001), has not been confirmed by further sampling (Yin et al. 2005). The most detailed and comprehensive marine-based magnetostratigraphic studies at this level are of the Shangsi section (Heller et al. 1988; Steiner et al. 1989; Glen et al. 2009), and place the base of a normal magnetozone (here called LT1n) 0.5 m above the base of the Feixianguan Fm. (Glen et al. in press), just above the extinction event and below the d13Ccarb negative peak. In all currently studied sections, the FO of H. parvus is within a normal polarity magnetozone (Fig. 1), but it also ranges into the lower parts of LT1r in the Shangsi, Abadeh and Lower Guandao sections (Fig. 1). However, the Shangsi data are different in detail from other well-dated marine sections, in that the FO of H. parvus (in bed 30; Nicoll et al. 2002) occurs below a well-defined reverse magnetozone (here equivalent of LT1n.1r, in beds 32 and 33; Fig. 1). Evidence for this magnetozone is strong at Shangsi, were it has been identified by Heller et al. (1988), Steiner et al. (1989) and Glen et al. (2009). There is also good evidence for a reverse magnetozone at about this level in the Deltadalen section on Svalbard (Hounslow et al. 2008a), and in S. China in the Shuijiang (c. 15 – 17 m above the base of the Feixanguan Fm.; Heller et al. 1995), and Hechuan sections (c. 3 m above the base of the Feixanguan Fm.; Steiner et al. 1989), although in all cases without evidence of H. parvus. In the Meishan GSSP, Li & Wang (1989) also detected a single reverse polarity level, some 1.5 m above the PTB (Fig. 1), which is here interpreted as probably LT1n.1r. This reverse magnetozone may also be present in the Guandao section (Fig. 1). Yin et al. (2001) recognized that the range of H. parvus and Otoceras boreale overlapped, but at that time placed the boreal Otoceras concavum Zone in the Permian. The new detailed data of Bjerager et al. (2006) has demonstrated that the PTB occurs within the range of O. concavum in the East Greenland successions, which indicates the PTB occurs within the lowest part of magnetozone LT1n (Fig. 1). The same conclusion can be inferred in the Canadian Arctic successions (Henderson & Baud 1997), using the maximum flooding surface to infer correlation between Otto Fiord and Griesbach Creek.

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M. W. HOUNSLOW & G. MUTTONI

In conclusion, the following succession of events and markers is associated with the marine PTB sections: 1. Initiation of a strong palynofloral turnover, corresponding to a major floral extinction event (in the late Changhsingian), located within a reverse magnetozone. This turnover appears to be coincident with extinction events in marine biota, as described by Jin et al. (2000a), Looy et al. (2001), Sephton et al. (2005) and others. 2. The base of normal magnetozone LT1n, within the latest Changhsingian. 3. A minimum in d13C representing the climax of the Late Permian extinctions and its effect on Earth systems. This level seems to approximate the major tetrapod extinction event (Ward et al. 2005). 4. The FO of H. parvus in Chinese sections, indicating the base of the Triassic, within the occurrence range of O. concavum. 5. The base of sub-magnetozone LT1n.1r. 6. The base of LT1r, which occurs within the occurrence ranges of H. parvus and the Boreal ammonoid O. boreale (Hounslow et al. 2008a). The composite GPTS in Figure 1 differs from the solution of Steiner (2006), in that the equivalent of LT1n.1r is a magnetozone of half-bar width, and we specifically tie the magnetostratigraphy to other events. Steiner’s (2006) magnetostratigraphic solution also has a Changhsingian part of the Meishan section in the Griesbachian and a clearly Griesbachian part of the Shangsi section in the Changhsingian, both clearly erroneous correlation solutions.

The magnetostratigraphy of the Induan– Olenekian boundary The base of the Olenekian is provisionally defined in the Mud M04 section in India (Spiti) by the FO of Neospathodus waageni s.l. (Fig. 2), corresponding also to the initial increase of a positive peak in d13Ccarb and an associated FO of the ammonoid Rohillites rohilla (Krystyn et al. 2007b). This isotopic peak has been dated at Guangxi, China at 251.2 (+0.2) Ma (Galfetti et al. 2007a; Fig. 2). Currently, the only section in which this boundary can be closely related to a magnetostratigraphy is West Pingdingshan, at Chaohu in China (Krystyn et al. 2007b; Sun et al. 2007, 2009), just below the base of normal magnetozone WP4n (Fig. 2). The magnetostratigraphy of the upper part of the West Pingdingshan section appears to bear a close correspondence to that from the Hechuan section, which however lacks significant biostratigraphy near the Induan–Olenekian boundary (Fig. 2). The Dienerian interval at Guandao is characterized by

Neospathodus dieneri and Ns. pakistanensis conodont faunas and is dominated by reverse polarity, although there are sampling gaps and intervals of breccia. This probably correlates to the interval at Hechuan that includes the upper parts of the Feixianguan Fm. (Fig. 2). The Ns. pakistanensis conodont fauna underlying the positive peak in d13Ccarb can be closely related to similar events at the Induan– Olenekian boundary in the proposed GSSP at Mud (Krystyn et al. 2007b). The sections from the Sverdrup Basin and Spitsbergen appear to provide the most continuous magnetostratigraphy across the Induan– Olenekian boundary, but have a somewhat spotty ammonoid and conodont biostratigraphy and are not easily related to the proposed GSSP or sections in China. New conodont data from the Creek of Embry section on Ellesmere Island in Canada (Baud et al. 2008; Beatty et al. 2008; T. Beatty pers. comm. 2008) show Ns. krystyni and Ns. kummeli overlain by Ns. dieneri, suggesting that the base of magnetozone CE1r is equivalent to the base of WP3r (Fig. 2), which suggests the base of the Olenekian is within the topmost part of CE1r (i.e. LT2r). The apparent equivalent to magnetozone CE1r at the Griesbach Creek section on Axel Heiberg Island (Arctic Canada) is GC2r, in which Vavilovites sverdrupi occurs some 15 m below its top (Fig. 1). The Creek of Embry section displays a further c. 300 m before the first Euflemingites, indicating a rapid sedimentation rate in this section during the earliest Olenekian. Unfortunately, in this section the magnetostratigraphy over this interval is not particu-larly well defined, with many uncertain levels, but nevertheless appears to display two reverse sub-magnetozones within a normal-polarity dominated interval that probably correlates to magnetozone Vh6 at Vikinghøgda and LT4n in the composite (Fig. 2). The interval LT3n to LT4r appears to correlate to a reverse-polaritydominated interval at Vikinghøgda, Hechuan and Gaundao (Fig. 2). Unfortunately, conodonts from strata containing Hedenstroemia hedenstroemi are not known from Arctic Canada, and Ns. waageni occurs commonly with ammonoids from the E. romunderi Zone (Orchard 2008). Likewise, Ns. cristagalli ranges from the B. strigatus Zone into the V. sverdrupi Zone (Orchard 2008), which, together with the range of this conodont at Chaohu, suggests that the Olenekian boundary is much higher in the Creek of Embry section than suggested by the magnetostratigraphy (Sun et al. 2009). These inconsistencies may be resolved with a more detailed bio-magnetostratigraphy from the lowest Olenekian. The R. rohilla ammonoid zone at the Mud section can be correlated to the Kashmirites densistriatus beds at Guangxi, which lie above beds in

LT4

H.

he

de

LT3

ns

tro

em

erd

LT2

rup

i-

P. c a

ndi

dus

Dienerian

i

V. s v

@ Konickites

O. com

mune

O From Fig. 1

LT1

.b or ea le

n.1r

O. concav

um

Scaling section

Induan

B. ro senk ran striga tzi/ tus Griesbachian

Ns. svalbardensis

@

LT5

Smithian

@

Olenekian

Lower Spathian

MC

Vikinghøgda

LT6

Ng. meishanensis, Ng. taylorae

E. romunderi

Vh8 Vh7 Vh6 Vh5 Vh4

B. rosenkrantzi

Vh3 Vh1r

O. boreale

Vh2

20 m c

A. blomstrandi

Lusitaniadalen Mb

Ns. cristagalli Ns. krystyni Ns. kummeli

Deltadalen Mb

Creek of Embry

1r

Ns. waageni

Smith Creek Mb E. romunderi

2

Ns. svalbardensis

3

An. tard us

An. tardus

100 m I. staeschei

H. aequabilis Zone

H. parvus

H. praeparvus H. typicalis

Ns. pakistanensis

Mag. Zone (CE)

Confederation Pt. Mb

Confederation Pt. Mb

4

Ns. dieneri

E. romunderi, M. gracilitatus

Smith Creek Mb

P. obliqua Zone

6

H. anceps Zone

I. isarcica

T

Mazzin Mb

Andrez Mb

7

5

Approx. first Euflemingites at Greisbach Creek

Smith Creek

~40-45 m Ns. cristagalii

Ar. nodosus

Mag. Zone (SC)

50 m

~ 251.2 Ma (CHIN 40)

Bulla & Siusi (Italy)

Ns. waageni

Ns. pakistanensis Ns. discreta

R. rohilla M. ver.

Gyronites fr.

Ns. dieneri

Ammonoid genera A. = Arctoceras An. = Anawaschites Ar. = Arctoprionites B. = Bukkenites. E. = Euflemingites H. = Hedenstroemia M. = Meekoceras O.= Otoceras R. = Rohillites V. = Vavilovites

2 1

Campil Mb

Seis Mb

Bed No.

9

Ophiceras sp.

Dalong Fm

?

10

Ammonoid biozones

Konickites sp. Claraia spp.

17

Flemingites Euflemingites

δC13carb

Induan Olenekian

Ns. dieneri Ns. discretus Ns. aff cristagalli Ns. waageni

18

? Talus breccia -no sampling

Ns. kummeli

19

11

?

3

c

25 m

F1

H. parvus

0

10 (m) 20

WP1r

(m)

Shale

21

16 15 14 13 12 11

13 12

22

WP2

100

24 23

Ng. planata

F2

14

c

Ng. krystyni

0

17 16 15

25

H. typicalis

F3

(m)

Feixianguan Fm

Ns. dieneri

30

Bed No

Jialingjiang Fm

c

WP4 Magnetozone

Shale

J1

WP3

c

Mud/ Spiti (India)

1m

Ns. gr. bransoni Ns. waageni

West Pingdingshan

F4

H. parvus Ng. changxingensis

Ns. cristagalli; Ns. pakistanensis c

Hechuan

Boreal Ammonoid Zone

GEOMAGNETIC POLARITY TIMESCALE

Ns. discretus; Ns. conservativus

Lower Guandao

GPTS

Vikinghøgda (Spitsbergen, arctic Norway)

E. romunderi

Ellesmere Is. (arctic Canada)

South China

Permian

67

Fig. 2. Summary of the bio-magnetostratigraphy across the Induan–Olenekian boundary. Section data from left to right: Lower Guandao (Payne et al. 2004; Lehrmann et al. 2006); Hechuan (Steiner et al. 1989), West Pingdingshan (Sun et al. 2007, 2009); Mud, Spiti (Krystyn et al. 2007b); Bulla & Siusi (Scholger et al. 2000; Perri & Farabegoli 2003; Horacek et al. 2007); Ellesmere Island (Ogg & Steiner 1991; Orchard 2008; Beatty et al. 2008; Baud et al. 2008; T. Beatty pers. comm. 2008); Vikinghøgda (Hounslow et al. 2008a; Nakrem et al. 2008). See Figure 1 for key. Ammonoid genus abbreviations in key, others on Figure 1. Thickness scales different for each section. MC ¼ magnetochrons.

68

M. W. HOUNSLOW & G. MUTTONI

which H. hedenstroemi is found (Galfetti et al. 2007a; Krystyn et al. 2007b), suggesting that the FO of H. hedenstroemi lies below the base of the Olenekian. This cannot be demonstrated in the Boreal sections with magnetostratigraphy, where H. hedenstroemi at Griesbach Creek (Fig. 1) occurs some 12 m above the top of the magnetozone GC3n (Ogg & Steiner 1991; Tozer 1994; Hounslow et al. 2008a), which is the probable equivalent of LT3n (Fig. 1). At Griesbach Creek, Euflemingites cirratus occurs a few metres above the top of the section measured by Ogg & Steiner (1991), an ammonoid species which does not co-occur with the zonal index E. romunderi in Canada (Tozer 1994), adding some support to our interpretation that magnetozone GC4n is the equivalent of LT4n, and that the first Euflemingites (at Griesbach Creek) occurs within LT4n (Figs 1 & 2). In the Bulla/Siusi sections in northern Italy, the correlation (Fig. 2) of the upper-most normal magnetozone (within the Seis Mb) to West Pingdingshan is consistent with the d13C positive isotope peak, which characterizes the Induan–Olenekian boundary interval (Tong et al. 2007; Richoz et al. 2007). This positive isotopic peak occurs some 40–45 m above the top of the section sampled for magnetostratigraphy (Horacek et al. 2007; Posenato 2009). However, according to Kozur and Bachmann (2005), the base of the Olenekian in the Italian Bulla/Suisi sections is at the base of the Pachycladina obliqua conodont Zone, at odds with both the carbon isotopic data, interpretation by others of the conodont data (Posenato 2009) and the magnetostratigraphy, which suggests instead that the P. obliqua Zone in these sections begins in the base of LT2r in the Dienerian (Figs 1 & 2). Within the normal and reverse polarity parts of magnetozone LT2, a number of sections show tentative magnetozones. Particularly significant may be those within the lower part of CE1r (Creek of Embry section), West Pingdingshan (WP3r), Gaundao and the F3 member of the Hechuan section (Fig. 2). These may indicate brief submagnetozones in LT2n and LT2r. The base of magnetozone LT2n appears to be a useful approximation of the Griesbachian –Dienerian boundary, as evident by Ns. dieneri and Proptychites sp. at this level in the Lower Guandao and Abadeh sections (Fig. 1). Our synthesis of the Induan and Lower Olenekian magnetostratigraphy is similar to that of Steiner (2006), but differs in detail, because she tried to integrate both marine and nonmarine studies in a composite ‘pattern matching’ GPTS. The age assignments of Steiner (2006) are also strongly influenced by the intersection correlations and biostratigraphy presented by Ogg & Steiner (1991), which are flawed (Hounslow et al. 2008a).

Magnetostratigraphy of the Olenekian– Anisian boundary The base of the Anisian is likely to be defined within the De˛sli Caira section in Romania, although the exact boundary is not yet decided (Gra˘dinaru et al. 2007; Hounslow et al. 2007a) (Fig. 3). The boundary is here informally placed at the FO of the conodont Chiosella timorensis. There are many magnetostratigraphic studies through the Lower –Middle Triassic boundary interval, from both low- and high-palaeolatitude marine sections (Muttoni et al. 2000 and references therein; Lehrmann et al. 2006; Hounslow et al. 2007a, b), and nonmarine sections (Steiner et al. 1993; Nawrocki & Szulc 2000; Huang & Opdyke 2000; Hounslow & McIntosh 2003; Szurlies et al. 2003; Szurlies 2007; Dinare`s-Turell et al. 2005). These studies provide independent assessment of the sequence of polarity reversals across the Olenekian–Anisian boundary and are supplemented by bio-magnetostratigraphies through the remainder of the Middle Triassic (Muttoni et al. 2000, 2004a; Szurlies 2007; Hounslow et al. 2008b). The Spitsbergen sections have a Spathian magnetostratigraphy that is the best constrained by an ammonoid biostratigraphy (Hounslow et al. 2008b). These data suggest that the lower Spathian is dominated by normal polarity, with a single reverse magnetozone (LT6r) detected at Milne Edwardsfjellet (MF1n.2r), Vikinghøgda (Hounslow et al. 2008b), and the Creek of Embry section on Ellesmere Island (i.e. CE7r). The data from Ellesmere Island suggest that the overlying reverse magnetozone (CE8r, equivalent to LT7r) has a more substantial thickness, which Hounslow et al. (2008b) correlated with a magnetozone within the lower part of the Keyserlingites subrobustus Zone in the Milne Edwardsfjellet and Vikinghøgda sections. It is not possible to confirm this pattern with data from the Hechuan and Guandao sections, because age dating of the Hechuan section is poor, and the lower Guandao magnetostratigraphy has many sampling gaps over this interval. The other reliable magnetostratigraphy at about this level appears to be that from the Moenkopi Group in northern Arizona, with a Tirolites ammonoid fauna (indicating Spathian), that succeeds beds with Anasibirites and Wasatchites faunas indicating the late Smithian (see discussion later; Fig. 4). The uppermost Olenekian is characterized by a reverse magnetozone LT9r, which can be correlated between sections at Kc¸ira (Muttoni et al. 1996a), Chios (Muttoni et al. 1995), De˛sli Caira (Gra˘dinaru et al. 2007), Guandao (Lehrmann et al. 2006; Orchard et al. 2007), and Milne Edwardsfjellet (Hounslow et al. 2008b). Equivalents of magnetozone LT9r also appear to be present in the Hechuan section, if the conodonts Neospathodus

Illyrian

MT6

Pelsonian

MT4r

Scaling section Upper Silesia

MF7

Aegean Anisian

Bith.

MT4n

.1r .1r

MT3n

Spitsbergen Desli Caira

Olenekian

.1n

LT9r

LT9n

LT8n

LT7n

low.

o

LT6

Smithian

E. romunderi

LT8r

upper Spathian

MF2

An. tardus

MT1

1n.2r

Spitsbergen

E.r, M.g, A.b

Ellesmere Is. & Spitsbergen

MF6 MF5 MF4 MF3

S.s

K. s A.n

Vendomdalen Mbr Botneheia Fm (part)

A. varium Grambergia sp.

Cs. timorensis, Ng. ex. gr. regalis

@

MT2

LT7r

MF1n

0

MT5

MT3r

MF1r

B. euomphala

Conodonts

MF8

Ng. cornuta

Ni. kockeli Ni. germanica

S5 S4 S3 S1

20 m Svartfjeld Mb

CE9 100

(m)

Cs. = Chiosella Gd. = Gladigondolella Ni. = Nicoraella Ng. = Neogondolella Ns. = Neospathodus Pa. = Paragondolella Pr. = Pridaella Tr. = Triassospathodus

(m)

Cs. gondolelloides Cs. timorensis Gd. tethydis

S6

Mid. Musch.

Lower Muschelkalk Gogolin Beds

Tr. ex. gr. homeri

Röt Fm

Pr.b.bifurcata Pa.bulgarica

LT9n

Cs. gondolelloides

Cs. timorensis

?

5

0

@

P. grambergi

S.s = Svalbardiceras spitzbergense K.s = Keyserlingites subrobustus E.r = Euflemingites romunderi M.g = Meekoceras gracilitatus A.b = Arctoceras blomstrandi A.n = Arctoprionites nodosus A. = Anagymnotoceras K. = Karangatites B. = Bajarunia P. = Parasibirites

m

20

10

K.s

Ns. waageni

(m)

0 0

Ns. homeri

Ammonoids

0

30

10

GPTS

@

Ellesmere Is. (arctic Canada)

CE8

J1 100

Talus breccia (no samples)

Cs. timorensis

0

@

K. evolutus

CE7

50

m

Spitsbergen Composite (arctic Norway)

@

(Mag. Zone numbering from Creek of Embry)

?

(m)

Kç1n

Ns. symmetricus

0

5

10

Ns. homeri

?

Kç1r

Ns. abruptus

J2

c

(m)

10

LT9n

(m)

c

Ns. triangularis

Jialingjiang Fm

?

LT9n

30

Ns. gr. bransoni

Tr. ex gr. homeri Ns triangularis Ns. crassatus

.1n

Ns. homeri

247.2 Ma

GU1r

J3

Ni. kockeli Gd. budurovi Pa.hanbulogi Pa.p.praezaboi Gd. tethydis Ng. cornuta Pa. liebermani

(Albania )

Kç2r

Kç2n

c

.2n

Ni. germanica

Desli Caira (Romania) A. ugra beds

Chios A+B (Greece)

N1r N1n

Upper Silesia composite (Poland)

(Albania)

0

Kç3n

Japonites sp. Pa. regale

Aegeiceras ugra

20 10

Deslicairites Japonites beds beds

MT3r- MT4n

Hechuan

Cs. timorensis

GU4 GU3r

GU3

30

Kç3r

0

GU2

40

Pa. bulgarica

10

N2n

Ni. kockeli

G1r

(m)

246.8 Ma

(Italy)

Granitovo

G2n

Kçira composite

Ni. kockeli Ni. germanica

~12m

G3r G3n G2r

Meandrospira deformata Pilamminadensa

Lower

DR1r

Nderlysaj

GEOMAGNETIC POLARITY TIMESCALE

Ni. germanica Cs. gondolelloides Cs. timorensis

Ng. regalis Ni. kockeli

Upper

Guandao

Pr.cornuta

(Italy)

S. China

Balatonicus Z. Trinodosus Z. s.l.

DontMonte Rite

69

Fig. 3. Summary of the bio-magnetostratigraphy across the Olenekian–Anisian boundary. Section data from left to right: Guandao (Lehrmann et al. 2006; Orchard et al. 2007); Hechuan (Steiner et al. 1989); Chios (Muttoni et al. 1995); Dont-Monte Rite and Nderlysaj (Muttoni et al. 1998); Granitovo (Muttoni et al. 2000); Kc¸ira (Muttoni et al. 1996a); De˛sli Caira (Gra˘dinaru et al. 2007). Ellesmere Island composite is from the Smith Creek and Creek of Embry sections of Ogg & Steiner (1991) and Orchard (2008). Upper Silesia composite (Nawrocki 1997; Nawrocki & Szulc 2000); Spitsbergen composite (Hounslow et al. 2008a, b; Galfetti et al. 2007b). The base Anisian used is the first occurrence of Cs. timorensis in the De˛sli Caira section (Gra˘dinaru et al. 2007). Thickness scales different for each section. See Figure 1 for key. Ammonoid genus abbreviations from Figure 2 and key.

V.h = Voltziaceaesporites heteromorpha C.p = Cycloverrutriletes presselensis

S-4

Lower Anisian L. Spath

Smithian

O l e n e k i a n

S-3 S-2 N

Pomorska Fm

Upper Spathian

MC

PAZ S-5

Lower M. Upper B.

Röt Fm Polczyn Fm

A.k I.c V.h D.n

Z4

S-L

2

PZr1

O.e

L-T

n.1r

TTS

Diener. Changh- Griesbachian I n d u a n singian

O

LT2

LT1

Tbn1

F. v

LT3

P(S-1)

Baltic Fm

Tbr1

Lower Buntsandstein

D.p

Tbn3

Tbn2

V.s L.v

Middle Buntsandstein

Ni. k Parotosuchus

C.p Quickborn Sandstone

5

LT5

LT4

4

Bernburg [10]

LT6

0

Tbr2

Z3

D.n = Densiosporites nejburgii D.p = Densiosporites playfordi

Ophiceras V.s = Vittatina spp. L.v = Lueckisporites virkkiae I.c = Illinites chitonoides A.k = Angustisulcites klausii

Zechstein

Miospores L.o = Lundbladispora obsoleta

LT7

Tbr3

C

252.6

MT1

LT8

Zechstein

11

100 m

O, C

Dinwoody Fm

Tbr4

Tbn4

C

MT2

Tbr6

Tbr5

Q

MT3

LT9

Tbn5

L.o

0

o

H ?

Tbn6

3

251

50

B.b

Eo.

Mag. Z. (CG) Solling [4]

10 9

0

Tbn7 ?

8

Hardegsen [12]

B.t

100 m

7

Detfurth [4]

Volpriehausen [12]

250

B.t

6

ATS (Ma) 247 249

Middle Buntsandstein

Pariot Fm

Tbr7

V

252

(m) 0

Chugwater Fm

Lystrosaurus sp. C Dicynodon

Katberg Fm

Palingkloof Mb

Lystrosaurus sp. A Lystrosaurus sp. B

(m) R1

50

D

(m)

r

4 LT

Alcova Lmst Mb

Sewemup Fm

5n LT

0

Dolores River region (Colorado)

Moenkopi Group

lower red

Ali Baba Fm

Virgin Fm

Lower Buntsandstein

A.k; T.s Wasatchites

LT7r

Karoo Basin (S. Africa)

N1

130 m

LT8

middle red

Calvörde [10]

0

248

100

Sinbad Fm

t

S

z4-z6 z7

Shnabkaib Fm

Moenkopi Group

Wellesaurus pebbodyi

Tenderfoot Fm

Tupilakosaurus

Benthosuchus Wetlugasaurus

R2

Parotosuchus

~750 m

N2

Kraznokamenskaya/ Petropavlovskaya

~330 m

Kzyl-Sal 280 m

Eo.

60

Vetluga

N3

Ramashkinskaya

N2

Spasskiy R1

KrasnobackovskiyShilihinskiy

Ryabinskiy N1

N2

Spasskiy Shilihinskiy

Krasnobackovskiy

R1

N1

Ryabinskiy

F.v, O.e

Holbrook Mb upper red/ Moqui Mb

Polish Buntsandstein Composite

TTS = Top Terrigenous Series [ ]= number of fining-up cycles in Buntsandstein units

Fig. 4. Correlation between the magnetostratigraphy of non-marine Lower Triassic sections and the marine GPTS. Section data from left to right: Russia–columns left to right from Figures 16, 29 and 14 of Molostovsky (1983), vertebrate data from Figure 35 of Molostovsky (1983), two left columns have no vertical scale; Karoo Basin (Ward et al. 2005; Steiner et al. 2003); Virgin River– Gray Mountain (Bissell 1973; Steiner et al. 1993; Lucas et al. 2007a, b); Dolores River (Helsley 1969; Helsley & Steiner 1974); Chugwater composite (Boyd & Maughan 1973; Shive et al. 1984; Steiner 2006); Central German composite (Szurlies 2007; Hounslow et al. 2007b; Heite et al. 2005, 2006; Heunisch 1999; C. Huenisch pers. comm.); Polish Buntsandstein composite (Nawrocki 1997; Szurlies 2007). Thickness scales different for each section except for Moscow Basin and Obslicliey Syrt which have no scale. See Figure 1 for key. PAZ ¼ Boreal miospore zonation (Hounslow et al. 2008b). MC ¼ magnetochrons.

M. W. HOUNSLOW & G. MUTTONI

Moscow Basin

Obslicliey Syrt

Urals region composite

Crow Mountain Sandstone Mb

Russia

[Blue Holes/ Red Grade sections]

Upper B. [8]

Virgin RiverGray Mountain (N. Arizona)

Eo.= Eocyclotosaurus

GPTS

Central Germany Composite

70

B.b = Beneckeia buchi

Chugwater Composite (Wyoming)

Red Peak Member

A.k = Anasibrites kingianus Ni.k = Nicoraella kockeli T.s = Tirolites spinosus F.v = Falsisca verchojanica B.t = Beneckeia tenuis O.e = Otonisporites eotriassicus

GEOMAGNETIC POLARITY TIMESCALE

triangularis and N. homeri are used to constrain the correlation with the Guandao sections (Fig. 3). Reverse magnetozone LT9r has at least one normal submagnetozone (LT9r.1n), found at Milne Edwardsfjellet, De˛sli Caira and upper Guandao. There is some evidence of a second normal polarity submagnetozone within LT9r in the upper Guandao (i.e. GU1r.1n) and Hechuan sections, although the magnetostratigraphic data from both Guandao sections are ‘noisy’ and contain many ‘half-bar’ tentative submagnetozones. Based only on the match of magnetostratigraphic polarity pattern, Steiner (2006) suggested this interval (i.e. J3 interval of the Jialingjiang Fm.; Fig. 3) at Hechuan to be early Spathian, which is not supported by the presence of late Spathian conodont faunas. Magnetochrons MT1 and MT2 characterise the Olenekian –Anisian transition and can be confidently correlated between Kc¸ira, De˛sli Caira, and the Spitsbergen composite section (Fig. 3). The reverse and normal parts of these magnetochrons appear to vary somewhat in relative thickness, probably due to sedimentation rate and/or sampling density differences in the sections near the Olenekian –Anisian boundary (Fig. 3). Magnetochrons MT1 and MT2 were not detected at Chios due to faulting and the presence of a hiatus at the top of LT9r (Muttoni et al. 1996a). A predominantly normal polarity interval (MT3n to MT4n; Fig. 3) succeeded by a predominantly reverse polarity interval (MT4r to MT6r) are present in the Anisian, up to the Pelsonian and Illyrian substages. This pattern is observed at Kc¸ira (Muttoni et al. 1996a), the Upper Silesia sections in Poland (Nawrocki & Szulc 2000; Nawrocki 1997), the Albanian Nderlysaj section, the Dont-Monte Rite section from the southern Italian Alps (Muttoni et al. 1998), as well as the Granitovo section from Bulgaria (Muttoni et al. 2000). The sections at Guandao and Hechuan are difficult to relate to the Kc¸ira and De˛sli Caira sections at the base of the Anisian, because MT1 and MT2 appear to be missing in the Chinese sections where MT3n (of Bithynian age, indicated by Nicoraella germanica) rests directly on LT9r (Fig. 3). The FO of Ch. timorensis also appears to be diachronous relative to the magnetostratigraphy (Hounslow et al. 2007b). The original correlation at Guandao proposed by Lerhmann et al. (2006) suggested that GU2n is the equivalent of MT4n (Fig. 3), a correlation that is driven by the conodont biostratigraphy but that largely ignores the magnetostratigraphy. They also suggested that submagnetozones GU1r.1n and GU1r.2n are the equivalent of MT1n and MT2n; this is only likely if order of magnitude fluctuations (on a metre to 10 m scale) in the sedimentation rate occurred, which is not evident in the lithology of these sections (cf. Lehrmann et al.

71

2006; Orchard et al. 2007). The problems of correlating Guandao to other sections also relate to the fact that Nicoraella kockeli and Ni. germanica appear very low in the section compared to the biomagnetostratigraphy in other sections. Both these conodonts appear to first occur high up in MT4n at Kc¸ira and in the Upper Silesia composite section (Muttoni et al. 1996a; Nawrocki & Szulc 2000; Szurlies 2007).

Magnetostratigraphy of the nonmarine Lower Triassic There is a close correspondence between the magnetostratigraphy of Lower Triassic marine successions and that from nonmarine successions, although correlation details are often debatable without other constraining stratigraphic data (Steiner 2006; Szurlies et al. 2003; Szurlies 2007; Fig. 4). The extinction events in the latest Permian are well constrained in the Karoo Basin (South Africa) and probably the German Buntsandstein, by a negative d13Corg peak within what appears to be the lower part of the equivalent of LT1n (Fig. 4). In the Karoo Basin, constraint is also provided by the vertebrate extinction event indicated by the last occurrence of Dicynodon (Ward et al. 2005). In contrast to the marine extinction event, the tetrapod turnover is in the lowest part of LT1n rather than in the underlying reverse magnetozone. In the Buntsandstein successions of Germany and Poland, the terrestrial extinction event is not well marked; typically late Permian palynomorphs such as Lueckisporites sp. and Vittatina sp. are separated from typically Triassic forms such as Lundbladispora and Densoisporites by a barren interval, covering the Z4 –Z7 part of the Zechstein and lowest part of the Lower Buntsandstein (Fijalkowska 1995; Heunisch 1999; Yaroshenko & Lozovsky 2004). The correlated proxy for the PTB in the Buntsandstein is the base of the Falsisca verchojanica conchostracan zone, coincident with a negative carbon isotopic peak (Kozur & Bachmann 2005). It may be that the major sediment provenance change in the Z7 cycle of the Zechstein (Hiete et al. 2005, 2006) is a proxy for the latest Permian extinction turnover (Fig. 4), in that the terrestrial extinction event may have radically affected sediment transport systems, through a decline in sediment trapping by vegetation loss (cf. Sephton et al. 2005). Magnetozone LT1n is also apparent in the multiple-section data from the Moscow Basin and the Urals region (Molostovsky 1983, 1996; Taylor et al. 2009; Fig. 4). In the Wordie Creek Formation in Greenland, the tetrapods Tupilakosaurus, Wetlugasaurus and Luzocephalus co-occur with ammonoids that range in age from Griesbachian to early

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Dienerian (Lozovsky 1998; Lucas 1999). Hence, the occurrence of these tetrapods in the Russian sections (shown under Urals composite in Fig. 4; Molostovsky 1983), allows an approximate correlation to the marine substages. The lowest parts of the Russian Vetluga successions also appear to preserve the Permian–Triassic transitional palynoflora that occurs below the PTB at other locations (Krassilov et al. 1999; Yaroshenko 2005), which appears to particularly characterise the lower part of magnetozone LT1n (Hounslow et al. 2008a; Metcalfe et al. 2009). In addition, Russian sections at Blyumental/ Kou-Su, Chesnokovka and Buzuluk–Grchevka (three of the six sections contributing to the Urals composite; Molostovsky 1983), Boyevaya Gora, Tuyembetka (Taylor et al. 2009) show the equivalent of LT1n.1r within the LT1n magnetozone (Fig. 4). Overlying magnetozone LT1n in marine sections is a mid Griesbachian to mid Smithian reversepolarity dominated interval (LT1r to LT4r; Figs 1 & 2), which has clear parallels in the sections from Russia (Molostovsky 1983, 1996), the lower threequarters of the Moenkopi Group in Colorado and the Chugwater Formation in Wyoming (Fig. 4). For many years the Russian composite magnetostratigraphy through this dominantly reverse polarity interval excluded several tentative short-duration normal polarity intervals (Molostovsky 1983; Lozovsky & Molostovsky 1993) detected in sections such as the Sosnovyiy/Vetlyanovskiy and Blyumental/Koya-Su gorges (Molostovsky 1983) and Sarysu (Khramov 1987), but these are now included (Molostovsky 1996). This uncertainty in part reflects the low palaeomagnetic sampling density used in some of these Russian sections and the common lack of proper demagnetisation. It is debateable where the LT1r –LT4r reversepolarity-dominated interval begins in the German and Polish Buntsandstein sections. Szurlies (2007) has suggested that CG3r/Tbr1 is the equivalent of LT1r (and CG5n/Tbn3 ¼ LT3n), a correlation that is strongly influenced by the debateable age assignments of the Buntsandstein conchostracan fauna outlined by Kozur (1999) and Kozur and Bachmann (2005); particularly, the placement of the base of the Olenekian (Posenato 2009). The primary means of locating the Olenekian in the Buntsandstein is the stratigraphically ‘close’ occurrence of M. truempyi with the Olenekian ammonoid Flemingites flemingianus in Madagascar (Kozur & Bachmann 2005), suggesting that M. truempyi is indicative of the Olenekian. Stratigraphic ‘closeness’ is here not a strong case for dating in illdocumented successions fluctuating from marine to non-marine conditions. However, support for the correlations of Szurlies (2007) comes from Galfetti et al. (2007a), who suggest a duration of c. 1.4 Ma

for the Induan, which is similar to the c. 1.2 Ma duration based on cyclostratigraphy, from the FO of F. verchojanica to the base of magnetozone CG5n (correlated to the base of LT3n by Szurlies 2007). However, a number of problems remain with the correlations of Szurlies (2007). Firstly the cyclostratigraphy from the Calvo¨rde Fm. suggests that CG3r is exceptionally brief at some c. 50 ka when other sections show a longer relative duration, and an interval occupied by some c. 1.5 ammonoid zones (Figs 1 & 4). Secondly the dominance of normal polarity (i.e. interval CG3n/Tbn1 to CG4n/Tbn2) in the Calvo¨rde Fm. is more compatible with this interval being equivalent to LT1n (Fig. 4). Thirdly, others have determined different numbers of cycles in the units of the Buntsandstein (Geluk & Ro¨hling 1999), which may reflect the non-basin centre focus of some of the sections of Szurlies (2007). For these reasons an alternative, more likely correlation, is suggested in Figure 4. The upper boundary of the reverse-polarity dominated interval from LT1r to LT4r probably represents the top of GC7r/Tbr5 in the Buntsandstein and an equivalent level in other sections (Fig. 4). Alternatively, Szurlies (2007) has correlated LT5n with GC7n which seems less likely, considering that normal polarity dominates from LT5n in Boreal sections (e.g. Fig. 2), and LT4n is most likely correlated to the Tbn4–Tbn5 interval in the Polish Middle Buntsandstein, which Szurlies (2007) has correlated to CG7n. Unfortunately, there appears to be no useful palynology from the Lower Buntsandstein, Volpriehausen or Detfurth formations to confirm or deny these two sets of proposed correlations (C. Heunisch, pers. comm. 2008; Fig. 4). The normal-polarity-dominated interval from LT5n to LT9n has clear parallels in the normalpolarity-dominated upper parts of the Middle Buntsandstein (Fig. 4). However, a confident match in the relative thicknesses of the three reverse magnetozones between the GPTS and the central German Buntsandstein composite is not visually convincing. The thickest reverse magnetozone in the marine composite is LT7r from the Boreal sections (Figs 3 & 4), which either represents magnetozone GC8r or GC9r. Similarly, magnetozone LT7r appears equivalent to the reverse magnetozone spanning most of the ‘middle red’ and Virgin Formation in the northern Arizona composite. The ammonoids Anasibrites kingianus and Wasatchites sp. from the Sinbad Fm. of the Moenkopi Group in Utah indicate the upper Smithian (Lucas et al. 2007b). However, the magnetostratigraphy for the Sinbad Formation in the Virgin River-Gray Mountain composite (Fig. 4) is derived from a thinner succession of the Sinbad Formation at Lees Ferry, farther south in Arizona (Steiner et al. 1993), so it

GEOMAGNETIC POLARITY TIMESCALE

is not clear that the ammonoids and magnetostratigraphy represent exactly the same levels. The ammonoid Tirolites spinosus collected from low in the Virgin Formation near the Utah –Arizona border (Bissell 1973) is good evidence of the Spathian. Like the Moenkopi Group in northern Arizona, sections from the Urals appear to display a particularly thick representation of LT7r. In the Russian sections, the correlation is supported by the cooccurrence of the tetrapod Parotosuchus with the Spathian ammonoid Tirolites cassianus, within the Bogdinskaya Member at Bolshoye Bogdo Mountain, in the Cis–Caspian depression (Molostovsky 1983, 1996; Molostovsky et al. 1998). Evidently, the Spathian displays particularly dramatic interregional changes in sedimentation rates that are probably the reason for such large variations in the relative thickness of these reverse magnetozones. Szurlies (2007) has outlined the reasoning for correlating the German Buntsandstein magnetozone GC10r to the equivalent of LT9r in the Kc¸ira section. The Anisian ammonoids Beneckeia tenuis, B. buchi, and Balatonites ottonis and the conodont Nicoraella kockeli occur in the Upper Buntsandstein and provide reliable biostratigraphic ties to the marine Middle Triassic (Fig. 4). This is supported by the presence of Stellapollenites thiergartii in a miospore assemblage from the Upper Buntsandstein (Visscher et al. 1993; Heunisch 1999). In addition, the FO of Illinites chitonoides, within the upper part of the Hardegsen Fm., has an equivalent FO in the Svalis-4 palynostratigraphic assemblage zone of Vigran et al. (1998). In the Vikinghøgda section, this FO is within the lower part of LT9n (Mørk et al. 1999; Hounslow et al. 2008a; Figs 3 & 4).

Anisian – Ladinian and Ladinian –Carnian boundaries The magnetostratigraphy and biostratigraphy of the Anisian –Ladinian boundary successions from the Tethyan marine realm have received considerable attention in recent years (Fig. 5). The first attempts at magnetostratigraphy were carried out in the latest Anisian Prezzo Limestone and Buchenstein Beds in the Southern Italian Alps, which are largely overprinted by excessive heating, caused by the Late Eocene –Early Oligocene Adamello batholith (Muttoni & Kent 1994). Hints of primary magnetic components that survived this overprinting were tentatively isolated and assigned to a latest Anisian interval of normal polarity with a duration of perhaps one million years and spanning the Trinodosus Zone and most of the overlying Lardaroceras-bearing beds (Muttoni & Kent 1994 and references therein). Subsequent analyses of coeval sections from the Southern Alps (Muttoni,

73

unpublished data) revealed, however, that these magnetic components may in fact be pre-folding remagnetizations of about Late Cretaceous age. Attention was therefore paid to the island of Hydra (Greece) where Angiolini et al. (1992) identified the Anisian –Ladinian boundary in the 24-m-thick Aghia Triada section, from the nodular, reddish Han-Bulog Limestone. The first results (Muttoni et al. 1994) yielded a consistent magnetostratigraphic pattern relating to a late Anisian –Ladinian conodont biostratigraphy, in which the FO of Pridaella trammeri (¼ Gondolella trammeri, ¼ Paragondolella trammen) was used as a proxy of the Anisian –Ladinian boundary, results which were refined by Muttoni et al. (1997, 1998) (Fig. 5). Confirmation of the Aghia Triada results were sought in the coeval 60-m-thick Vlichos section on Hydra (Muttoni et al. 1997, 1998, 2000), but there tectonic complexities disrupted the original stratigraphic continuity, confirmed later by field inspection (GM), and the Vlichos data are therefore excluded from this compilation. In the late 1990s, attention moved back to the Alps, and specifically to the Dolomites in Italy and the Northern Calcareous Alps in Austria. Magnetostratigraphic investigations on biostratigraphicallydated limestones and radiometrically-dated tuffs in the Buchenstein Beds from the Dolomites started with the Anisian –Ladinian boundary interval in the Fro¨tschbach section (Muttoni et al. 1996, 1997), and were followed by parallel studies on the nearby coeval Pedraces and Belvedere sections (Brack & Muttoni 2000). A satisfactory magnetostratigraphic correlation was obtained on laterally traceable limestone and volcaniclastic intervals, showing the high degree of reproducibility of the magnetic polarity fingerprint through the Anisian – Ladinian boundary interval, throughout much of the Buchenstein Basin of the Dolomites (Fig. 5). In the Northern Calcareous Alps, Austria, Gallet et al. (1998) produced a magnetostratigraphy from four coeval sections (Mendlingbach 1 and 2, and Gamsstein East and West), which again showed a consistent polarity pattern across the Anisian – Ladinian boundary interval constrained by conodont biostratigraphy (Fig. 5). A breakthrough study was carried out on the c. 110 m-long Seceda core, drilled by the Geological Survey of Bolzano in 1998 at Mount Seceda in the northwestern Dolomites (Brack et al. 2000). With over 90% recovery, this core offered a unique opportunity to reconstruct, in stratigraphic continuity, a portion of the Middle Triassic pattern of polarity reversals (Muttoni et al. 2004a). The Seceda core spans a complete succession of Buchenstein Beds with limestone and associated ‘Pietra Verde’ volcaniclastic layers, which were correlated to the nearby Seceda outcrop section with associated radiometric

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Fig. 5. Summary of the bio-magnetostratigraphy across the Anisian– Ladinian and the Ladinian– Carnian boundary. Section data from left to right: Margon–Val Gola (Gialanella et al. 2001 reinterpreted by Brack et al. 2001), Fro¨tschbach (Muttoni et al. 1996b, 1997), Seceda (Muttoni et al. 2004a), Felso¨o¨rs (Ma´rton et al. 1997; Vo¨ro¨s et al.

GEOMAGNETIC POLARITY TIMESCALE

75

Fig. 5. (Continued) 2003), Pedraces and Belvedere (Brack & Muttoni 2000), Aghia Triada (Muttoni et al. 1998 and references therein), Mendlingbach West, Gamsstein 1, Mayerling (Gallet et al. 1998), and Stuores (Broglio Loriga et al. 1999). Stratigraphic depth of sections expressed in metres. In left panel: (1) Middle Triassic stages, (2) Middle Triassic sub-stages, (3) ammonoid zonation and (4) composite magnetostratigraphic sequence arranged in magnetochrons MT6 to UT3. Thickness scales different for each section. See Figure 1 for key. Conodont genera: B., Budurovignathus; Gl., Gladigondolella; M., Metapolygnathus; N., Neogondolella; P., Paragondolella; Pr., Pridaella; S., Sephardiella. [for Fro¨tschbach, Seceda, Pedraces, and Belvedere, see compilation in Muttoni et al. (2004a); for Aghia Triada, Mendlingbach West, Gamsstein 1, and Mayerling, see compilation and notes in Muttoni et al. (2000); for Stuores, see Broglio Loriga et al. (1999)].

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and biostratigraphic age data. Two ash layers located in the ‘Lower Pietra Verde’ and ‘Upper Pietra Verde’ intervals in the Seceda outcrop section yielded U– Pb ages of 241.2 (þ0.8–0.6) Ma (SEC.22) and 238.0 (þ0.4–0.7) Ma (SEC.21), respectively (Fig. 5; Mundil et al. 1996; Brack et al. 1996; updated by Brack et al. 2007), indicating an average sediment accumulation rate of c. 10 m/ma. More recently, Brack et al. (2007) obtained a new radiometric age estimate (BIV-1) from the Upper Anisian Bivera Fm. at Monte Bivera (Trinodosus Zone) based on 12 individual zircon ages ranging from 243.3 Ma to 241.6 Ma. Because of Pb loss and the mild leaching step, Brack et al. (2007) interpreted the age of 243.3 Ma to be the minimum age for sample BIV-1, which should fall just below the base of the Seceda magnetostratigraphic sequence within the Trinodosus Zone (Fig. 5). Magnetostratigraphic data from Seceda were correlated with data from the coeval Fro¨tschbach, Pedraces, and Belvedere sections, as well as Margon-Val Gola from Trentino (Gialanella et al. 2001; reinterpreted by Brack et al. 2001), and a satisfactory correlation was obtained (Fig. 5). The magnetozone interval SC2n –SC3n at Seceda corresponds with F1n–F2n at Fro¨tschbach, P1n–P3n at Pedraces, SL1r–SL2n at Belvedere and M1n– M2n at Margon–Val Gola, and with similar patterns at Aghia Triada, Mendlingbach West, and Gamsstein 1. The magnetostratigraphic data from the Felso¨o¨rs section, Hungary (Ma´rton et al. 1997; Vo¨ro¨s et al. 2003) show normal polarity through the F1n magnetozone ranging through the Reitzi Zone, whereas at Seceda the same ammonoid zone is found reverse in magnetozone Sc1r (correlated to Seceda following Brack et al. 2005; Fig. 5). However, the Felso¨o¨rs section contains major sampling gaps due to thick tuff layers that mostly did not yield samples. Also, in the sampled layers some palaeomagnetic data yielded dubious directions, with in two cases normal and reverse polarity reported from the same level (limestone bed 99B and tuff layer between limestone beds 101 and 102; Fig. 5; Ma´rton et al. 1997). These dual-polarity problems presumably indicate the existence at Felso¨o¨rs of unresolved normal polarity overprints; therefore, we maintain SC2r as the main late Anisian reverse polarity zone corresponding to MT6r in the GPTS (Fig. 5). Support for this stance comes from the nonmarine studies in the Anisian, which suggest additional detail in MT6r, which is not seen in any marine section over this interval (see below; Fig. 6). The magnetostratigraphic correlations between Seceda, Fro¨tschbach, Pedraces, Belvedere, MargonVal Gola, Aghia Triada, Mendlingbach West, and Gamsstein 1 allow the generation of a reference magnetostratigraphy with the U–Pb dates from Seceda

and Monte Bivera and the ancestor– descendant faunal associations of paragondolellids, neogondolellids (conodonts) and ammonoids, present in different degrees of preservation, in all correlated sections from the Dolomites. This augments the numerical and biostratigraphic definition of the Anisian –Ladinian GSSP, at the FO of Eoprotrachyceras curionii (base of the E. curionii Zone), 5 m above the base of the Buchenstein Beds in the Bagolino section of northern Italy (Brack et al. 2005). This level can been traced to metre level 83.7 in the Seceda core corresponding to metre level 14.7 in the Seceda outcrop section, very close to the base of reversal SC2r.2r and c. 5 m above the level with a U –Pb age of 241.2 (þ0.8–0.6) Ma (Muttoni et al. 2004a; Brack et al. 2005; Fig. 5). The magnetostratigraphy of the Ladinian – Carnian boundary interval has not been as extensively studied as the Anisian –Ladinian boundary. Gallet et al. (1998) presented the magnetobiostratigraphy of the Mayerling pelagic limestone section from the Northern Calcareous Alps, which contains a rich conodont fauna encompassing the uppermost Anisian to Lower Carnian. The .60 mthick Mayerling section, with 14 well-defined polarity intervals spanning a succession of agediagnostic conodont events, is currently the most continuous marine section studied through the Ladinian –Carnian boundary interval (Fig. 5). The 160 m-thick Stuores section in the Dolomites, with higher sediment accumulation rates, covers a shorter time interval compared to Mayerling and has been extensively studied for biostratigraphy and magnetostratigraphy (Broglio Loriga et al. 1999). Stuores is the ratified GSSP for the base of the Carnian, with the FO of the ammonoid Daxatina canadensis as the basal Carnian marker (Mietto et al. 2007), with a succession of polarity reversals that can confidently be correlated to Mayerling (Broglio Loriga et al. 1999). The only other section studied for magnetostratigraphy across the Ladinian –Carnian boundary is from Spitsbergen (Arctic Norway), in which a large part of the late Ladinian is missing near the boundary (Hounslow et al. 2007a). The boundary interval is dated by sporadic occurrences of conodonts and ammonoids. The magnetostratigraphy confirms that the correlated base of the Carnian occurs within the stratigraphic gap, below the traditional Boreal Carnian (at the base of the Stolleyites tenuis ammonoid biozone).

Age calibration of the Middle Triassic GPTS According to Hinnov & Goldhammer (1991) and Preto et al. (2001) the Latemar carbonate platform

GEOMAGNETIC POLARITY TIMESCALE

77

GPTS

Illyrian Pelsonian Bith.

ANISIAN

MT9

Curionii

MT8 MT7

Reitzi 5 MT4

Trino.

MT6

SS = substage

Aegean

MT3

??

OLENEKIAN

CG10

BSPB = Budleigh Salterton Pebble Beds MZ = magnetozone

P.g = Protodiploxypinus gracilis

LT9

LT8

CG9

M.d = Microcachryidites doubingeri

A.z = ammonoid zone

Solling Formation

V.h = Voltziaceaesporites heteromorpha

50 m

D.n = Densiosporites nejburgii M.f = Microcachryidites fastidioides

Paracrochordiceras

MT2

M.f S.t M.d

1

Solling disc.

S.t = Stellapollenites thiergartii (H. muelleri)

Balatonit.

MZ

V.h

CG12 CG11

Röt Formation

D.n

Jena Fm

Muschelkalk

S3

0

Costatoria goldfussi mansuyi

Kocaelia

S6 S5 Eocyclotosaurus

S1

??

MZ

Röt Fm 20 m

40 m

Central Germany composite

S4

Lower Muschelkalk Gogolin Beds

BS2

BSPB

BS1

MZ

20

Upper Silesia composite (Poland) Mid. Musch.

Unit C

BS3n

Unit A

S. thiergartii, T. falcata

Caldes SC Fm

Unit B

BS3r

Rillo Fm S. thiergartii, Triadispora falcata

Lundbladispora sp. Cycadophytes sp.

El Figaró SM Fm

Ns. kockeli

MT4r

Thust

Mb 1

?

RSJ Units

Mb 2

Badong Fm

Mb 3

100 m

Otter Sandstone Formation

BS6

BS5 BS4

R5

Eocyclotosaurus, Kapes

BS8

Pennington Point Mb

Seced.

Mercia Mudstone Group

BS7

P. g, T. falcata, M. doubingeri Triadispora aurea, M. fastidioides, Ovalipollis sp.

Mb 4

UESC complex

S. thiergartii V. heteromorpha, Triadispora sp.

M1

Ng. constricta

Torete Fm

Badong, Sangzhi & (NE Spain) Nanzhang (S. China)

LADINIAN

MC A.z SS

Molina de Aragón composite Iberian Basin Budleigh Salterton(Spain) Riera de Sidmouth Tramascastilla Sant Jaume Fm (M3) (Southern UK) Catalan Basin

Hardegsen disc.

Fig. 6. Summary of the bio-magnetostratigraphy from non-marine Middle Triassic sections, and possible correlations to the marine bio-magnetostratigraphy. Section data from left to right: Badong, Sangzhi and Nanzhang (Huang & Opdyke 2000); Riera de Sant Jaume (Arche et al. 2004; Dinare`s-Turrell et al. 2005); Molina de Arago`n (Rey et al. 1996; Arche et al. 2004; Dinare`s-Turrell et al. 2005); Budleigh Salterton-Sidmouth (Benton 1997; Hounslow & McIntosh 2003); Upper Silesia (Nawrocki & Szulc 2000; Nawrocki 1997); Central Germany (Heunisch 1999; Szurlies 2007; Hounslow et al. 2007b, C. Heunisch, pers. comm. 2008). Thickness scales different for each section. See Figure 1 for key. GPTS column from Figures 3 and 5.

in the Dolomites, which has a platform interior characterized by a c. 470 m-thick lagoonal succession consisting of c. 600 shallowing-upward cycles can be attributed to a 9– 12 Ma record of precessional forcing of sea level change. However, U –Pb dating of zircons from volcaniclastic layers within the Latemar succession (from top to bottom: LAT –32, 241.7 þ1.5/20.7, Mundil et al. 2003; LAT-30: 241.2 þ0.7/20.6, Mundil et al. 2003; re-dated to 242.8 + 0.2 Ma, Brack et al. 2007; LAT-31: 242.6 + 0.7 Ma, Mundil et al. 2003),

and the correlative basinal Buchenstein Beds (SEC.22, SEC.21; see above) suggests that the Latemar cycles only span at most a few million years (Brack et al. 1996; Mundil et al. 1996, 2003). Similarly, Kent et al. (2004) showed that most of the Latemar succession is of normal magnetic polarity, which together with biostratigraphic and lithostratigraphic correlations between beds in the adjacent Buchenstein Basin, suggested that the bulk of the Latemar platform deposition was coeval with magnetozone SC2n at Seceda

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(MT7n in GPTS; Fig. 5). Kent and coworkers therefore concluded that magnetozone SC2n (at Seceda) and the time-equivalent Latemar deposition, in fact, had a duration an order of magnitude less (c. 1 Ma). Therefore, straightforward interpretation of the U –Pb age model for Buchenstein deposition, and the magnetostratigraphy through the Secedensis Zone are internally consistent, and opposed to the cyclostratigraphic analyses of the Latemar succession by Hinnov & Goldhammer (1991) and Preto et al. (2001). We, therefore, like Kent et al. (2006), abandon the long duration hypothesis for the Latemar succession based on cycle counting, accepting U –Pb zircon dates as the main constraints on the duration of the Middle Triassic. Like the problematic magnetostratigraphy from the Felso¨o¨rs section, the multigrain U –Pb-zircon ages of Pa´lfy et al. (2003) are distinctly too young compared to the radiometric age data from the Dolomites. As reported by Brack et al. (2005), these ages overlap within error with respect to SEC22 from Seceda. The younger mean values of the seemingly stratigraphically older layers at Felso¨o¨rs may be due to the (unresolved) contribution from grains affected by lead loss. Robust age constraints on the composite Middle Triassic GPTS are based on the U –Pb age data of BIV-1 (minimum age of 243.3 Ma) from the Trinodosus Zone at Monte Bivera, and SEC.22 (241.2 þ0.8–0.6 Ma) and SEC.21 (238.0 þ0.4–0.7 Ma) from the Buchenstein Beds as discussed above, with the 241.2 Ma age closely associated with the Anisian –Ladinian boundary at the base of the Curionii Zone (Brack et al. 2005). A Ladinian– Carnian boundary at c. 235 Ma (Brack et al. 2005) to c. 236 Ma (Fig. 5) is derived from U –Pb data and field observation from the upper Ladinian granites at Predazzo in the Southern Alps. This intrusion, dated at 237.3 (þ0.4/21.0) Ma (sample PRE-26, Fig. 5; Brack et al. 1996), post-dates the Ladinian Buchenstein Beds in which the youngest U-Pb age is 238.0 (þ0.4/20.7) (Mundil et al. 1996), and pre-dates sediments of the Wengen Volcanosedimentary Group (S. Cassiano Formation), which toward its top contains the Ladinian– Carnian boundary at Stuores (Broglio Loriga et al. 1999). Therefore, the Ladinian–Carnian boundary should be just a few Ma younger than c. 237 Ma (i.e. c. 235 –236 Ma; Fig. 5). With a

Ladinian –Carnian boundary at c. 235.5 Ma and an Anisian –Ladinian boundary at c. 241 Ma, the Ladinian Stage is c. 5.5 my duration. The composite sequence of magnetic polarity reversals outlined here encompasses 34 magnetozones spanning the late Anisian– early Carnian that are arranged in a sequence of magnetochrons from MT6 to UT3 (Fig. 5).

Magnetostratigraphy of the nonmarine Middle Triassic As with the Lower Triassic, a substantial number of studies of the Anisian have been made in nonmarine successions (Fig. 6). The generally distinctive character of the Anisian, with a lower part dominated by normal polarity followed by dominantly reverse polarity in the mid Anisian, generally allows for a confident correlation with the Middle Triassic GPTS (Fig. 6); correlations largely follow those used by the original authors. Magnetochrons MT1 and MT2 are probably present in the Riera de Sant Jaume section (Catalan Basin), and the Middle–Upper Buntsandstein boundary interval (uppermost Solling Fm. to lowermost Ro¨t Fm.) from the Germanic Basin (Fig. 6; Dinare´sTurell et al. 2005; Szurlies 2007). One of the features at this level, which supports the central German Basin magnetostratigraphic correlation in Figure 6, is the appearance of Triadispora sp., which in the Milne Edwardsfjellet section is consistently present from magnetochron MT2 (Hounslow et al. 2008b), the approximate correlative level at which this pollen becomes common to abundant in the German Upper Buntsandstein. Magnetochron MT2 also appears to have been detected by Huang & Opdyke (2000) in the Badong Fm. in South China (Fig. 6). Correlations to the GPTS from the Upper Silesia (Poland) and the central Germany composites, over the magnetochron interval MT3 and MT4, are constrained by a variety of ammonoid, conodont, and basin-wide borehole geophysical log data (Kozur 1999; Szurlies 2007). Studies in the Catalan and Iberian basins can be correlated to the Triassic GPTS, and are largely constrained by a fragmentary palynostratigraphy at the formation level, a regional lithostratigraphy, and overlying marine sediments (Muschelkalk facies) that have generally better

Fig. 7. Summary of the bio-magnetostratigraphy across the Carnian– Norian boundary. Section data from left to right: Guri Zi (Muttoni et al. 2005); Pizzo Mondello (Muttoni et al. 2004b); Silicka´ Brezova´ (Channell et al. 2003); Bolu¨cektasi Tepe (Gallet et al. 1992, 2000a); Kavaalani (Gallet et al. 2000a). Stratigraphic thickness of sections in meters. In right panel, composite GPTS arranged in magnetochrons UT19 to UT14 with indication of the position of the conodont-based Carnian–Norian boundary (provisionally placed at the FOs of Metapolygnathus echinatus and M. parvus at Pizzo Mondello; Nicora et al. 2007) and U –Pb radiometric age estimate of magnetochron UT10n from Furin et al. (2006). Thickness scales different for each section. See Figures 1 and 5 for key.

GEOMAGNETIC POLARITY TIMESCALE

Fig. 7. Continued.

79

80

M. W. HOUNSLOW & G. MUTTONI

age control in the Ladinian (Arche et al. 2004; Dinare´s-Turell et al. 2005). Those for the Otter Sandstone Fm. in the Budleigh Salterton to Sidmouth section (UK) are age constrained by tetrapods, including Kapes and Eocyclotosaurus, which indicate the Anisian Perovkan land-vertebrate faunachron (Benton 1997; Hounslow & McIntosh 2003). There are indications from the sections at Molina de Aragon and Budleigh SaltertonSidmouth that the magnetostratigraphic pattern may contain additional short normal magnetozones over the MT4r to MT6r interval that have not been well characterised in the marine sections shown in Figure 5. The character of the polarity pattern during magnetochrons MT5n–MT6r may therefore be better represented by data from these nonmarine sections.

Carnian – Norian boundary Candidate sections for defining the base Norian GSSP are Black Bear Ridge (Williston Lake, British Columbia, Canada; e.g. Orchard et al. 2001, 2007) and Pizzo Mondello (Sicily, Italy; Muttoni et al. 2004b; Nicora et al. 2007). Proposed guide forms for the base of the Norian are favoured as the FO of Metapolygnathus echinatus at Black Bear Ridge (Orchard 2007) or the FO of M. echinatus and Metapolygnathus parvus at Pizzo Mondello (Nicora et al. 2007). Integrated magnetostratigraphic and biostratigraphic studies across the Carnian –Norian boundary in the marine realm started in the early 1990s with the pioneering work of Gallet et al. (1992) on the 73 m thick, upper Carnian –upper Norian Bolu¨cektasi Tepe section in Turkey. The basal bedded limestone member, 10 m thick in this section, yielded a sequence of polarity reversals covering the stratigraphic range of Metapolygnathus communisti, which was used by these authors as a proxy for the base Norian (Fig. 7). Nearly 10 years lapsed until the broadly coeval, c. 55 m-thick, Kavaalani section, in Turkey, was published (Gallet et al. 2000a). In its basal 10 m, this section extends through the Carnian– Norian boundary interval and has age diagnostic conodonts that can be correlated to the Bolu¨cektasi Tepe section (Fig. 7). A magnetostratigraphic investigation of Middle and Upper Triassic fossiliferous limestones cropping out around Williston Lake in British Columbia (Canada), particularly at Black Bear Ridge, was attempted but the strata were found to be remagnetized during the early stages of Laramide folding in the Cretaceous (Muttoni et al. 2001a). The first magnetostratigraphic and biostratigraphic study of an expanded (c. 150 m-thick) Carnian –Norian boundary section at Pizzo Mondello in Sicily was published by

Muttoni et al. (2001b). Muttoni et al. (2004b) refined these initial findings, focusing primarily on the distribution of Metapolygnathus communisti and additional key species for the definition of the boundary (Fig. 7), and extended the analysis upwards through an additional c. 280 m of strata previously attributed to the late Norian by Gullo (1996). In a recent re-analysis of the Pizzo Mondello section (after Muttoni et al. 2004b), Nicora et al. (2007) established the FOs of Metapolygnathus echinatus and M. parvus in sample NA36, 8 m above Gonionotites maurolicoi and 7 m below a Norian radiolarian assemblage, with the conodont marker considered as suitable to define the base of the Norian Stage. The Silicka´ Brezova´ section, Slovakia is a composite compiled from seven separate but partially overlapping sections correlated by means of lithostratigraphic marker beds, biostratigraphy and magnetostratigraphy, and straddling c. 120 m of upper Carnian –upper Norian strata (Channell et al. 2003). Magnetostratigraphic and biostratigraphic data through the Carnian– Norian boundary were obtained from the Lower Trench, and Massiger Hellkalk Quarry sections (boundary used by these authors at FO of Norigondolella navicula: Fig. 7). Similarly, a magnetostratigraphy and conodont biostratigraphy through the Carnian– Norian boundary interval was obtained from the 70 m-thick limestone section at Guri Zi in northern Albania (Muttoni et al. 2005), with the boundary placed between the LO of Metapolygnathus nodosus (¼ Epigondolella nodosa) and the FO of Epigondolella abneptis. All these sections through the Carnian –Norian boundary interval contain a similar assemblage of conodonts within a framework of broadly correlative magnetozones (Fig. 7), although Guri Zi is affected by variations in sediment accumulation rates associated with turbiditic deposition of calcarenites (Muttoni et al. 2005). These biomagnetostratigraphies define a sequence of 16 magnetozones (organized in magnetochrons UT9 to UT14) of late Carnian (Tuvalian) to early Norian (Lacian) age, with the Carnian –Norian boundary provisionally placed at the FOs of Metapolygnathus echinatus and M. parvus at Pizzo Mondello (Nicora et al. 2007) (Fig. 7). A U –Pb zircon date of 230.91 + 0.33 Ma from late Carnian limestones in southern Italy (Furin et al. 2006; Fig. 7), are tentatively correlated with the lower part of the Silicka´ Brezova´ and Pizzo Mondello sections using a constraining conodont biostratigraphy (Fig. 7). A consensus (detailed below) has been reached for correlating the conodont-based Carnian –Norian boundary interval into the Newark Supergroup astrochronological polarity timescale (APTS) (Kent et al. 1995; Kent

NORIAN

GPTS UT6 UT5

N A I C

UT4

N

UT8

R

UT9

A

Paulckei Spinosus Subbullatus Dilleri

UT7

(I)

(II)

UT3 Aonoides

Scaling section

Substage & divisions Ammonoid Zones (II)

Julian 1

UT10

Austriacum

(I)

Lacian 2

Lac. 1 (II)

Tuvalian 3

(II)

(I)

Pizzo Mondello

Julian 2

Mayerling

Julian 1

Bolücektasi Tepe

Julian 2

Julian 2

UT11

UT2

UT1

MT13

MT12

20

LADINIAN

50 m

UT12

Daxatina

S1r

?

UT13

Regoled.

MA3n

@

UT1n

I

Jandianus

UT14

Arch.

S2n MA3r

I

From Fig. 7

S2r

MA4n

30

II

S4n S3r

40

1m

Julian 1

Trachyceras aon Subzone

@

II

Long 2

MA4r

I

I 1m

D. cf canadensis Subzone

MA5n

50

Julian 1

II

Regoledanus Subzone

Auriformis TadMostleri pole

60

I

Stuores (Italy)

Tuvalian 2

Tuvalian 3

Tuvalian 3

Stockton Fm

SB2n

Tuvalian 2

20 m

Bolücektasi Tepe Erenkolu Mezarlik II @

Diebeli

80 m

Ng. liardensis, Metapolygnathus ex. gr. polygnathiformis and M. cf. lobatus.

Stolleyites plana

1m

SB1n

Longobardicus

DL3

A.a = Aulisporites astigmosus, I.c = Illinites chitonoides K.m = Kuglerina meieri, K.r = Kraeuselisporites reissingeri S.r = Sellaspora rugoverrucata

Lacian 1

Erenkolu Mezarlik (Turkey)

5m

II

II

1m

Mayerling (Austria)

2

DL1

De Geerdalen Formation Tschermakfjellet Fm

DL1n.1r

Bolücektasi Tepe (Turkey)

SB5r

10 m

SB4r SB4n SB3r

I

I-

E3

?

II

10 m

@

Massinger

SB3n.1r

SB2r

Tuvalian 3

M-

E2

K.m A.a

PM3n

100 m

?

DL3

4

SB3r

PM4n

SB1r

40 m

DL5

Isfjorden Member

100 m I.c

Lunde Fm (Middle Mb)

S.r

E5 E4

SB4n

40 m

?

Sn1

@ 1 PM5n

PM2r

Tuvalian 2

Stockton Fm

Sn2

E6

Lower Trench

Pizzo Mondello (Sicily)

Newark Supergroup (USA)

Vendomdalen (Spitsbergen)

Sn3

2

@

E7

Sn4

Lacian

K.r

Div. D Division E

UT13r

Division F

Lunde Fm (Upper Mb)

Sn6 Sn5

Tuvalian 1

Silickà Brezovà Kaavalani (Slovakia) (Turkey)

Tampen Spur Snorre (North Sea, Norway)

Sn7

81

Guri Zi (Albania)

GEOMAGNETIC POLARITY TIMESCALE

Fig. 8. Summary magnetic polarity pattern for the Carnian. Section data from left to right: Tampen Spur, Snorre (Nystuen et al. 1989; Eide 1989; Bayer & Lundschien 1998); Vendomdalen (Hounslow et al. 2008b); Newark Supergroup (Kent et al. 1995; Le Tourneau 1999); Pizzo Mondello (Muttoni et al. 2004b); Mayerling (Gallet et al. 1998), Silicka´ Brezova´ (Channell et al. 2003); Stuores (Broglio Loriga et al. 1999); Kavaalani (Gallet et al. 2000a); Erenkolu Mezarlik and Bolu¨cektasi Tepe (Gallet et al. 1992, 2000a); Guri Zi (Muttoni et al. 2005). Julian, Tuvalian Lacian are substages and their subdivisions detailed in Gallet et al. (1992) and Krystyn et al. (2002). Thickness scales different for each section. See Figures 1 and 3 for key.

& Olsen 1999; Olsen & Kent 1999), suggesting that the Carnian– Norian boundary falls within the interval c. 227 –228 Ma (Muttoni et al. 2004b).

The Carnian GPTS The Carnian GPTS is in part built by extending those sections which cover the Ladinian –Carnian and Carnian –Norian boundary intervals, and using additional sections which are less well-dated (Fig. 8). The upwards extension from the Ladinian– Carnian boundary is largely achieved

using the Bolu¨cektasi Tepe section (Gallet et al. 1992; Gallet et al. 2002a), which appears to correlate approximately to the palynologically-dated lower Carnian section from the De Geerdalen Fm. in Spitsbergen (Hounslow et al. 2007a). These Spitsbergen sections may extend upwards into the middle or upper Carnian in the overlying Isfjorden Member (Fig. 8), but the palynological data is not sufficiently clear to confidently distinguish Carnian from Norian (Hounslow et al. 2007a). Likewise, the polarity pattern from the Carnian – Norian boundary can be extended down into UT9 (i.e. equivalent to SB1n) using the Lower Trench

82

M. W. HOUNSLOW & G. MUTTONI

section at Silicka´ Brezova´ (Fig. 8). The nonmarine sections from the Stockton Fm. (Newark Supergroup: Kent et al. 1995), and the Upper Lunde Fm. E/F divisions at Snorre (Bayer & Lundschien 1998) can be confidently matched to those marine sections from the boundary interval (Muttoni et al. 2004b). The Snorre composite through the Upper Lunde Fm. E and F divisions is placed in the upper Carnian, since the FO of Kraeuselisporites reissingeri (associated with Ovalipollis pseudoalatus in the Snorre assemblages; Eide 1989) is in the Tuvalian (Roghi 2004). In the Snorre area the K. reissingeri assemblage overlies more diverse assemblages in the lowest part of the Upper Lunde and Middle Lunde (Eide 1989; Nystuen et al. 1989) containing Enzonalasporites vigens, Granuloperculatipollis rudis, Triadispora obscura and O. pseudoalatus, amongst others. Comparison to ammonoid-dated sections from the Alps (Roghi 2004) suggests these assemblages are approximately Tuvalian in age, supporting the suggested magnetostratigraphic correlation. The magnetostratigraphic match of the Stockton Fm. to the marine sections suggests that magnetochrons E2 to E5 probably extend into the Dilleri Zone (Tuvalian– 1; Fig. 8). These correlations are weakly supported by the polarity pattern from the Spitsbergen Isfjorden Mb, in which magnetozone interval DL4 to DL5 appears to approximately match the Stockton Fm. E3 to E5 magnetozones (Fig. 8). Supporting data to confirm these correlations (or any other) in the middle Carnian are absent, and for this reason, the magnetochron UT5 to UT9 interval is displayed as half grey to reflect this.

The Norian GPTS The correlation of marine sections to the Newark APTS suggests that the Norian is the longest stage in the Triassic and as such merits special attention in terms of its magnetic polarity pattern. Key marine sections for the construction of a Norian GPTS are Bolu¨cektasi Tepe, Kavaalani and Kavur Tepe from Turkey (Gallet et al. 1992, 1993, 2000a), Scheiblkogel, Austria (Gallet et al. 1996), Silicka´ Brezova´, Slovakia (Channell et al. 2003), and Pizzo Mondello, Sicily (Muttoni et al. 2004b).

The GPTS and biostratigraphically-constrained polarity succession through the Norian has been much debated, and different solutions proposed by Krystyn et al. (2002); Channell et al. (2003), Hounslow et al. (2004), Muttoni et al. (2004b) and Gallet et al. (2007). Bolu¨cektasi Tepe, Kavur Tepe, Scheiblkogel, and Kavaalani were correlated by Krystyn et al. (2002) by fitting magnetic polarity zones into a correlation scheme based on conodont zonations. These conodont zonations (shown in Fig. 9) are related to age, largely through co-linked studies of ammonoids and conodonts in Timor (Indonesia) and some other Tethyan locations (Krystyn et al. 2002). This biozone framework was used to construct a composite sequence of magnetic polarity reversals, scaled to equal conodont biozone duration, tied to the numerical calibrations of the Upper Triassic stages of Gradstein et al. (1994). Subsequently, Gallet et al. (2003) used the same composite succession with updated (and more appropriate) numerical constraints, to construct a tentative uppermost Carnian –Norian biozone-scaled GPTS. Muttoni et al. (2004b) attempted to construct the uppermost Carnian– Norian magnetic polarity reversal pattern by adopting the 430 m-thick Pizzo Mondello section as a reference section that was correlated to the c. 140 m composite stratigraphy of the Silicka´ Brezova´ section. The Muttoni et al. (2004b) solution resulted in a somewhat lower number of Norian normal and reverse polarity zones (20) compared to Krystyn et al. (2002) (25), or the even higher number (31) that arises if the new Sevatian to Rhaetian data of Gallet et al. (2007) are taken into account. This discrepancy arose from the different procedures adopted for correlation, and the differing recognition of missing time in the sections. For example, Muttoni et al. (2004b) attempted to establish statistical correlations by using magnetostratigraphic fingerprints in a one-to-one magnetozone matching approach, assuming, as a first order approximation, that stratigraphic thickness is a linear function of time. Priority was given to expanded and lithologically homogeneous sections such as Pizzo Mondello that tend to minimize the problematic occurrence of stratigraphic or fault gaps (although high

Fig. 9. Summary magnetic polarity pattern for the Norian, based on inter-section correlation. Relative thickness of magnetozones determined by the correlation grid (horizontal lines), with the scale provided by the appropriate section indicated on the right (see text for details). Section data from left to right: Scheiblkogel (Gallet et al. 1996); Kavur Tepe (Gallet et al. 1993, 2000a); Pizzo Mondello (Muttoni et al. 2004b); Kavaalani (Gallet et al. 2000a); Silicka´ Brezova´ (Channell et al. 2003); Bolu¨cektasi (Gallet et al. 1992, 2000a). Thickness scales different for each section. See Figure 1 for key. Conodont zonation and sub-stage divisions of Gallet et al. (1992, 2000a) and Krystyn et al. (2002). Substage divisions: Tu–3 ¼ Tuvalian–3; La –1, La –2, La –3 ¼ Lacian 1, 2 and 3; Al–1, Al– 2, Al– 3 ¼ Alaunian 1, 2 and 3; Sev–1, Sev– 2 ¼ Sevatian 1 and 2. C.bZ ¼ conodont biozone. Conodont biozonation on the Silicka´ Brezova´ column from Channell et al. (2003). Sections segmented according to the hiatus and faulting information given by authors.

Fig. 9. Continued.

E. = Epigondolella M. = Misikella Me. = Metapolygnathus 20 m N. =Norigondollela P. = Parvigondolella

= Key section for vertical scaling

La-1

@

K

E. nodosa

E primitia

E.

ept

abn

. co Me

unis mm

ti B

. co Me i Me. oerti

ti unis mm

2

(II)

(I) 1

E. spatulata 3

Pizzo Mondello

N

E. n. sp. D

UT20

@

Pizzo Mondello

A l a u n i a n R I A N

@

Scheiblkogel

O

UT19

a ter 2

pos

Silická Brezová

E. 3

L a c i a n

E. multidentata

@

E. triangularis E. triangularis n. subsp

UT18

Al-2

E. bidentata

UT21

20 m

Sev-2

Sev-2

he rn ste ini

From Fig. 10

Sevatian

M.

UT22

P. andrusovi (Sev-2)

24

RHAETIAN

Division & substage

E. M. biden pos t.t.

UT23

M. bidentata M. hernsteini-

1

From Fig. 7

@

UT17

Sev-1

5m

Sev-1

Kavaalani

C. bZ

C A R N I A N

5m

UT16

5m

Sev-1

Sev-2

GPTS

Tu v a l i a n - 3

5m 15

5m

UT13 14

Silická Brezová

Al-3

Al-3

M. posteraM. zapfei

Me. nodosus

La-1

Bolücektasi

@

UT12

La-1

5m

E. triangularisN. halstattensis

Al-2?

E abneptis

Al-3 Al-3

Al-3

Al-2

5m

0

UT11

?

Al-1

5m

K

20 m

K

La-2

Al-1

Al-1

K

Silická Brezová

UT10

K

La-2

Al-1 La-3

La-3

20 m

Kavaalani

10

La-2

La-2

(m)

30

upper

Breccia

La-2

Alaunian

K

5m

La-2 20

Kavur Scheiblkogel Tepe

P. carpathica

Tu-3(II)

Pizzo Mondello

lower

Tu-3(I)

Pizzo Mondello

GEOMAGNETIC POLARITY TIMESCALE 83

2

is A

1

K

Section scaling

84

M. W. HOUNSLOW & G. MUTTONI

accumulation rate is no guarantee of completeness). However, these sections, with high accumulation rates, tend to be less fossil-rich than condensed sections with lower accumulation rates, so biostratigraphic completeness cannot be demonstrated. Fossiliferous sections were preferred by Krystyn et al. (2002) using first-order correlation by matching key conodont ranges from individual sections, allowing a multiple-section composite to be constructed. However, these fossiliferous sections are frequently condensed with stratigraphic gaps (indicated by hardgrounds) and faults, which need to be included in the correlation modelling. For example, the middle Norian (Alaunian) part of the composite bio-magnetobiostratigraphic sequence of Krystyn et al. (2002) has been constructed by piecing together magnetozones from individual sections segmented by hiatus and fault gaps. The GPTS equal-duration biozone concept (Krystyn et al. 2002; Gallet et al. 2003) has restricted value for marine to nonmarine correlations, which has in part lead to the continuing debate about the Norian. Instead of utilising the equal-duration biozone concept, we use an approach that consists of constructing a composite bio-magnetostratigraphy (Fig. 9), scaled to section thickness, using the following principles: † Data from the thickest sections is used to scale the composite, because such sections are likely to provide the best magnetostratigraphic detail. † A magnetostratigraphic correlation grid is constructed, based on the correlation between section magnetozones, guided by the biostratigraphic framework. † The correlation grid constrains the amount of vertical stretching which can be applied to the section data, and provides limits on the relative thickness of the magnetozones in the composite, by minimising the amount of missing section at stratigraphic or fault gaps. † Section repetition across faults is only detectable by the conodont biozone framework (some details in Krystyn et al. 2002), which for most of these Norian sections, has only been published in summary form. If sedimentation rates were constant in each section, but differed between sections, vertical stretching of scales would produce a perfect horizontal correlation grid, but since sedimentation rates tend to vary within sections, the correlation grid lines can be inclined (e.g. see magnetozone UT17r in Fig. 9). In addition, a limited number of anchor correlation lines exist, which constrain the amount of vertical stretch (shown as ‘@’ in Fig. 9). This process is similar to that used in constructing splicedcore composite sections, typically used on IODP

cores (Pa¨like et al. 2005), which are based on the principles of graphic correlation (Shaw 1964). The least constrained part of this procedure centres on the disconformities and faultedboundaries that fragment the section magnetostratigraphy (Fig. 9). The most parsimonious solution is one that minimises the likely missing intervals (Shaw 1964; Edwards 1989). The best constrained parts of the GPTS composite are in the lower Norian, where sections, without apparent breaks, provide good inter-section relative thickness constraints (Fig. 9). The Norian– Rhaetian boundary interval is also well constrained, although some problems remain (see Fig. 10 and later). The Pizzo Mondello section is particularly important for scaling the entire Norian composite. The least constrained part of the GPTS is in the middle Norian (Alaunian), where the composite magnetozones UT18 and UT19 are not well constrained in relative thickness. For example, UT19n is only constrained in the incomplete data at Pizzo Mondello, and UT19r is only constrained in the lower part of the fragmented Scheiblkogel section (Fig. 9). Similarly, the Alaunian-1 part of the Bolu¨cektasi Tepe section may represent a younger sub-magnetozone in UT18r, rather than being equivalent to UT18n. Likewise, the relative thickness of UT20n depends much on the amount missing and unsampled at Pizzo Mondello. The solution presented is similar to the ‘equal biozone’ GPTS of Gallet et al. (2007) but scaled to section thickness. This difference is particularly noticeable for the Lacian-3 interval whose only thickness constraint is c. 2 m of strata at the Kavur Tepe section (Fig. 9).

The Norian – Rhaetian boundary The GSSP section for the base of the Rhaetian is likely to be in Austria in the Steinbergkogel section A, at the FO of Misikella posthernsteini (Krystyn et al. 2007a). The conodont biostratigraphy, palynology and magnetostratigraphy of this and nearby sections have been studied in detail (Krystyn et al. 2007a). These demonstrate the similarity in conodont ranges and magnetostratigraphy of the Steinbergkogel sections, and previously published magnetostratigraphic data from Scheiblkogel (Austria), Pizzo Mondello (Sicily), Silicka´ Brezova´ (Slovakia) and the Bolu¨cektasi Tepe, Kavaalani, Kavur Tepe and Oyuklu sections in Turkey (Fig. 10). Overall, there is good agreement between the various magnetostratigraphies and biostratigraphies from the sections covering the Norian–Rhaetian boundary interval (Fig. 10). The Pizzo Mondello section has the largest sedimentation rate and most detailed magnetostratigraphy in the latest Norian, Steinbergkogel provides

Al-2?

D+ 15

5 m

5m

5m @

N R O

A l a u n i a n

3 @

N

Al-3

I

A

M. rhaetica

S e v a t i a n

R H A E T I A N

UT27 UT26

Brumano

post. hern. bident.post.

UT25 UT23 24 UT22

1 UT21

Pizzo Mondello

Sev-1 Al-3

Sev-2

Steinbergkogel

Kavur Tepe (Turkey)

O. paucidentata

P. andrusovi M. bidentata

Kavaalani (Turkey)

E-

Sev-1

G-

2

UT20

20 m

18

Al-3 2n

@

2 B7

Al-2 Scaling section

Zones & Substage

Al-3

2r

19

9n

I-

UT19

5m

10

GPTS

Scheiblkogel

Al-3

(m)

30 20

0

0

M. hernsteini

10n

9n

3n

M. posthernsteini

11n

9r

5

Al-2 = Alaunian 2 AL-3= Alaunian 3 Sev-1= Sevatian 1 Sev-2= Sevatian 2

11n

Bolücektasi (Turkey)

3r

Silická Brezová (Slovakia)

12n

10n

85

M. postera group

4n

Portello Gebbia Fm

Scheiblkogel (Austria) Sev-2

M. hernsteini M. posthernsteini

5m

4r

M. hernsteini

STK B+C

ST4/STK-A

M. posthernsteini Paracochloceras

E. bidentata 0

5n

P. andrusovi E. bidentata

Conodont genus E. = Epigondolella M. = Misikella O. = Oncodella P. = Parvigondolella

5

(m)

d

M. hernsteini

ate pe Re

Thrust

5m

50 m

117 m to FO of M. posthernsteini

B-

Metasibirites

BA+

M. hernsteini

BT1

UT25 5r

E. bidentata

D-

Sev-2

F-

C+

Sev-1

Thrust

?

E. bidentata

H-

BT2

M. rhaetica

BT3

M. posthernsteini O. paucidentata

ZU 3a

J-

BT4

ZU 2

Zu Limestone Fm (part)

ZU 3b

BT5n

Pizzo Mondello (Italy)

Steinbergkogel (Austria)

(m)

Oyuklu (Turkey)

Sev-1

Brumano (N. Italy)

M. hernsteini

GEOMAGNETIC POLARITY TIMESCALE

Fig. 10. Summary magnetostratigraphy of the Norian– Rhaetian boundary. Section data from left to right: Brumano (Muttoni et al. 2010); Oyuklu (Gallet et al. 2007); Steinberkogel (Krystyn et al. 2007a); Scheiblkogel (Gallet et al. 1996); Pizzo Mondello (Muttoni et al. 2004b); Silicka´ Brezova´ (Channell et al. 2003); Bolu¨cektasi (Gallet et al. 1992, 2000); Kavaalani (Gallet et al. 2000a); Kavur Tepe (Gallet et al. 1993, 2000a). Sections segmented according to the hiatus and faulting information given by authors. Thickness scales different for each section. See Figure 1 for key.

the most detail in the earliest Rhaetian, and Brumano (Italy) in the remainder of the Rhaetian (Muttoni et al. 2010). An important constraint in these correlations near the Norian –Rhaetian boundary is the FO of Misikella hernsteini (i.e. base of Sevatian 2) in normal magnetozone UT22n in most sections. At Pizzo Modello, this is in PM11n, whereas at Steinbergkogel this event appears to be slightly higher (Fig. 10). A second constraint is the relatively thick reverse magnetozone in the uppermost Alaunian (i.e. in Epigondolella n. sp D Zone, Alaunian 3) at Kavur Tepe (E2), Kavaalani (K18), and Pizzo Mondello (PM9r), which confidently defines the magnetozone UT20r in the GPTS (Figs 9 & 10). The solution for the Silicka´ Brezova´ to Pizzo Mondello correlation (Fig. 10) is that used by Gallet et al. (2007) and Muttoni et al. (2004b). Channell et al. (2003) originally suggested correlating magnetozones SB9n at Silicka´ Brezova´ with 3n at Scheiblkogel, B7n at Bolu¨celtasi Tepe,

KV15 at Kavaalani, and Dþ at Kavur Tepe (Fig. 10). These correlations seem less likely, because they would require: (a) substantial withinsection changes in the sedimentation rate at the Silicka´ Brezova´ section; (b) absence of reverse magnetozones within SB10n in the Silicka´ Brezova´ upper trench section (which was densely sampled); and (c) because it violates the biostratigraphic age constraints at Bolu¨cektasi Tepe. Signs of condensation are however seen in the lower part of the Silicka´ Brezova´ upper trench section, where it is most reddened and clasts (possibly reworked) occur in the lowest 5 m (Channell et al. 2003). An alternative correlation solution for Scheiblkogel is to correlate magnetozones 4n with PM11n at Pizzo Mondello and the underlying 3r with PM10r, because there is no direct evidence of the Alaunian 3 zone at Scheiblkogel (Fig. 10; Gallet et al. 1996). However, the correlation in Figure 10 is compatible with the data from Kavur Tepe in which Epigondolella (¼ Mockina) bidentata

86

M. W. HOUNSLOW & G. MUTTONI

ranges into the E2 magnetozone, suggesting that the base of the Sevatian 1 zone (E. bidentata biozone) extends into the top of UT20r (Figs 9–10). The Oyuklu section is complicated by a thrust near the base of the section, which repeats the overlying B2 magnetozone within the occurrence range of E. bidentata and M. hernsteini. Gallet et al. (2007) correlated magnetozone PM12n at Pizzo Mondello with magnetozone Eþ in the Oyuklu section (Fig. 10), but a more likely correlation is with Aþ at Oyuklu (Fig. 10). This solution is more compatible with the FO of M. posthernsteini that occurs within the B2 magnetozone at Oyuklu and within the Portella Gebbia Fm. (above the Cherty Limestone and magnetostratigraphy) at Pizzo Mondello (Gullo 1996; Fig. 10). The solution of Gallet et al. (2007) would have placed the correlated FO of M. posthernsteini within PM11r at Pizzo Mondello, which is incompatible with the biostratigraphy. Magnetozone UT25n in the composite GPTS is poorly represented in the section data (Fig. 10); its base appears to be represented in the Steinbergkogel and Oyuklu sections (i.e. Eþ at Oyuklu) and the upper part at Brumano (BT1 to BT2n) and Oyuklu (Gþ). The mid parts of UT25n appear to be represented at Brumano and Oyuklu, but the most continuous record in the Oyuklu section is disrupted by a thrust that forms the upper boundary to the apparent reverse sub-magnetozone (equivalent to BT1r? in the Brumano section) within UT25n. A section with a more complete magnetostratigraphy over this interval would improve the polarity pattern within magnetochron UT25n.

Norian – Rhaetian nonmarine studies and correlations to Tethyan sections The most important nonmarine magnetozone succession through the Norian is from the Newark Supergroup (Kent et al. 1995; Fig. 11). The magnetostratigraphy through the Lockatong Fm. of the Newark Supergroup is confirmed by studies from the Dan River –Danville basins in the eastern USA (Kent & Olsen 1997; Fig. 11). The magnetostratigraphy through the Chinle Group of the southwestern USA (Molina-Garza et al. 1996) is of insufficient sampling resolution to confirm the Newark Supergroup magnetostratigraphy, but appears to confirm the general reverse polarity character of the E8r to E12r and E17r to E20r intervals (Fig. 11). The polarity pattern through the lower and middle parts of the Passaic Fm. is confirmed by data from the Fundy Basin (Kent & Olsen 2000), St Audrie’s Bay (Hounslow et al. 2004) and the Lunde and Lewis formations in the northern North Sea (Hounslow et al. 1995; Bayer & Lundschien 1998; Bergan 2005; Fig. 11). The

dominantly reverse polarity interval E18r to E20r in the Newark Supergroup is confirmed by data from the Fundy Basin, the Chinle Group, St Audries Bay and the Upper Lunde Fm. (Kent & Olsen 2000; Molina-Garza et al. 1996; Hounslow et al. 2004; Bergan 2005; M. Bergan pers. comm. 2008). Likewise, the collection of poorly-dated UK and Norwegian non-marine sections and cored intervals appears to confirm the general character of the E14r to E18n interval from the Newark Supergroup (Fig. 11). The polarity character of the E13 to E14 interval is so far not strongly supported by data from any area outside the eastern USA. The correlation of the Newark Supergroup magnetostratigraphy to the numerous Norian–Rhaetian marine sections has attracted considerable debate, with alternative solutions offered by Gallet et al. (1993, 2000a, 2007), Krystyn et al. (2002), Channell et al. (2003), Muttoni et al. (2004b, 2010) and Hounslow et al. (2004). The fundamental reason for such continued debate is the absence of any strong supporting biostratigraphic information that allows detailed correlation of the Newark Supergroup with successions outside the eastern USA. Palynofloral zonations and land-vertebrate faunachrons provide mostly ambiguous, low resolution correlations (Cornet 1993; Fowell & Olsen 1993; Lucas & Tanner 2007). Correlations are also hampered by the somewhat ambiguous middle Norian magnetostratigraphy from the marine sections, which are fragmented by disconformities and faulting (e.g. Fig. 9). These debates have focused on two issues, firstly the location of the Carnian –Norian boundary in the Newark Supergroup and secondly the position and extent of the Rhaetian in the Newark Supergroup APTS.

The Carnian –Norian boundary in the Newark Supergroup Muttoni et al. (2001b) sought a match of the magnetozone patterns across the conodont-based Carnian –Norian boundary at Pizzo Mondello and the palynology-based Carnian –Norian boundary in the Newark Supergroup, which suggested a correlative Carnian– Norian boundary within Newark magnetozones E14n–E16n and hence above the palynological Carnian– Norian boundary as originally placed, within Newark magnetozone c. E13 (Kent et al. 1995). Krystyn et al. (2002) used data from Bolu¨cektasi Tepe, Kavur Tepe, Scheiblkogel and Kavaalani to construct an equal biozone upper Carnian –upper Norian composite which was then correlated to Newark Supergroup magnetozones E3 –E22 with the Carnian –Norian boundary placed in magnetozone E7 (Fig. 8), using only pattern matching criteria. Channell et al. (2003)

E23r

E23

t

20

0

Lower Passaic- Heidlersburg

c i a s P

a

s

E15

E14

E12 E12

E11

E10

New Oxford- Lockatong

D3

Lockatong Formation

E13

UND

Trujillo Fm

E16

CSM D5

4m

Stockton Fm

D1

500 m

Division E

E17

F

o

r

m

E18

D2

Sn5

E19

a

10

D4

UND = unnamed division LCM = lower Cow Branch Member

Upper Balls Bluff- Upper Passaic

500 m

n o i

(m)

F o r m a t i o n B l o m i d o n

Wingate Sandstone

Bl1

PAZ = Miospore assemblage Zone C.m =Corollina meyeriana,

30

Upper Cow Branch Mb

?

40

LCM

UT1

50

Dan River-Danville Basin (USA)

SA1r

Red Head Mb

E15n

Bl2

Revueltian

SA2n

Bull Canyon Fm

E16n

Chinle Group San De Cristo Mts (E. New Mexico)

SA3n.1r

SA3n

Bull Canyon Fm

20 m

Bl4

Bl3

Redonda Fm

SA3r.1n

Fault

E20

Bl5

SA3r

SA2r

4n

20 m

E22

E21

Bl6

Rock Point Fm

Apachean

10 m

SA4n

Redonda Fm

Branscombe Mudstone Fm

Lewis Formation 50 m

Division D

Redonda Fm

SA4r

III

Sn7

Sn6

C.m PAZ

Kayenta Fm

Newark Supergroup (USA)

North Mt. Basalt

SA4r.1n

? II

Apachean

Granuloperculatipollis rudis Ovalipollis pseudoalatus Rydon Member

Blue Anchor Fm

Beryl field, N. Sea (UK)

A Division B Division C

Sn12 Sn11 10

IV

Sn9

Upper Lunde Fm (34/7-9) 50 m

Chinle Group, Glen Canyon Group (Utah)

Bl7

Sn8

U p p e r L u n d e F m

Land vertebrate fauna-chron

St Audrie’s Bay (U.K.)

13

Sn14

Magnetozone Snorre, Tampen Spur, N. Sea (Norway)

Chinle Group (E. New Mexico)

87

GAV 77/3 Fundy Basin (Canada)

GEOMAGNETIC POLARITY TIMESCALE

E9

E8

Fig. 11. Summary of non-marine magnetostratigraphic studies of the Norian–Rhaetian. Section data from left to right: Snorre, Tampen Spur (Nystuen et al. 1989; Bayer & Lundschien 1998; Bergan 2005; M. Bergan, pers. comm. 2008); Lewis Fm., Beryl area (Hounslow et al. 1995; Bond 1997); St Audrie’s Bay (Hounslow et al. 2004); Chinle group (Reeve & Helsley 1972; Molina-Garza et al. 1996; Lucas 1999); Fundy Basin (Kent & Olsen 2000); Dan River-Danville basins (Kent & Olsen 1997); Newark Supergroup, Newark Basin (Kent et al. 1995). Thickness scales different for each section. See Figure 1 for key.

reached similar conclusions regarding the magnetostratigraphic position of the Carnian– Norian boundary in the Newark Supergroup, using a supporting argument based around the vertebrate Paleorhinus. Muttoni et al. (2004b) attempted a correlation using a correlation-coefficient-based statistical approach, relating the thickness of Pizzo Mondello magnetozones to the duration of the correlative Newark Supergroup magnetozones, for each of the 16 possible relations. From a statistical standpoint their option #2 was the most robust and indicated the position of the Carnian –Norian

boundary in magnetozone E7 (Fig. 8), similar to that proposal by Krystyn et al. (2002) and Channell et al. (2003), a correlation solution for the Lockatong Fm. and upper parts of the Stockton Fm. that is now generally accepted (Figs 8 and 12).

Location of the Norian – Rhaetian boundary in the Newark Supergroup Broadly there have been three proposed correlation options, which place the conodont-defined

88

M. W. HOUNSLOW & G. MUTTONI

Fig. 12. The main options for correlating the Newark Supergroup magnetostratigraphy to the GPTS from Figures 9 and 10. Double-headed dashed line emphasises the key interval of E20r and its suggested correlation to key intervals in the Newark Supergroup for these three options. See text for discussion. G2003 – alternative correlations for option B proposed by Gallet et al. (2003). See Figure 1 for key.

Norian –Rhaetian boundary in the Passaic Fm., correlation option A suggesting the lowest and option C the highest (Fig. 12). Option A has been proposed by Muttoni et al. (2004b, 2010), working from the Carnian –Norian boundary upwards (based on the above statistical

matching approach), suggesting the base Rhaetian correlates approximately with Newark Supergroup magnetozone E17 (Fig. 12), or somewhat above, at a level similar to the original palynological and astrochronological estimate of Olsen & Kent (1999). Channell et al. (2003) had reached much

Conodont Zones/ Substage 201.6 Ma 2

1

R H A E T I A N NORIAN

bident.post.

post. hern.

hae

tica

ult

im M.

M. r

a

UT28 UT27 .3r UT26 UT25 24

H E T T A N G I A N

Magnetochrons

Scaling section

Montcornet St Audrie’s Brumano ?

Sevatian

E18

Passaic Palynofloral event

UT23

M. hernsteini

E19

Talcott Basalt New Haven Fm

From Fig. 9

BT2 50 m

E21

Chatham Group

ZU 3a ZU 2

?

Agawan Group

Portland Fm

Newark Basin (USA)

LO Bulbilimnadia

K. reissingeri acme BT5n BT4

ZU 3b

M. ultima

δC

13 carb

Italcementi section BT5n 4 3

Zu Limestone Fm

E22

fault

E20

M. posthernsteini

R. germanicus

O. pseudoalatus

Conodonts

E23

BT1

Brumano section

D-

Holyoke Basalt Shuttle Meadow Fm

E24n

BT3

5r

?

H. Basalt East Berlin Fm

E23r

UT2

thrust

R. germanicus

Malanotte Fm

M. posthernsteini M. rhaetica

40 m

ZU 3b

M. hernsteini M. bidentata

H-

5m

SA4r

J-

Conodont bearing white Lmst

R. tuberculatus

Conodonts absent M. ultima

Phylloceras

ZU 3c .1r

K. reissingeri acme

Oyuklu (Turkey)

Ri. tuberculatus

10 m

Argille de Levallois

0 H24r

Lombardian Basin (N. Italy)

B-

SA4n

(m)

UT22

Classopollis > 20%

H25r

1060m 1070m

C.thiergartii (approx)

UT26r

Rh. rhaetica

SA5

5n.1r

Last conodonts

Westbury Fm

.3r .2r

10 m

Williton Mbr

Blue Anchor Formation

Palynomorphs C. = Cerebropollenites K. = Kraeuselisporites O. = Ovalipollis Ri. = Ricciisporites R. = Rhaetopollis Rh. = Rhaetogonyaulax

I

Rydon Member

Last conodonts

O. pseudoalatus R. germanicus, Rh. rhaetica

5m

13 org

Lilstock Fm

Tpo TH-zone

UT27r

RPo RL

δC

“Initial negative excursion”

Kössen Fm

Schattwald beds

M1

500

1050m

Miospore zone equivalents

C1 TPo

K. reissingeri acme

C2

TH

( To Sinemurian 130 m above )

P. planorbis

org

SA5r

M2

13

SA6n

δC

Lias Group

Miospore Zones

K. reissingeri acme C.thiergartii

P. cf spelae P. ex gr. tilmanni P. pacificum

Tiefengraben Mb

St. Audrie’s Bay (UK)

GEOMAGNETIC POLARITY TIMESCALE

Rhaetian Jurassic

Kuhjoch (Austria)

GPTS H26r

500m

Isotope correlations Magneto correlations

Hartford Basin (USA)

Sinemurian?

1040m

Montcornet core Paris Basin (France)

89

Fig. 13. Summary magnetostratigraphy of the Triassic–Jurassic boundary. Section data from left to right: Kuhjoch (Kuerschner et al. 2007; von Hillebrandt et al. 2007); St Audrie’s Bay (Hesselbo et al. 2002; Hounslow et al. 2004); Montcornet core (Yang et al. 1996); Oyuklu (Gallet et al. 2007); Lombardian basin (Muttoni et al. 2010), Newark and Hartford basins (Kent et al. 1995; Kent & Olsen 2008). Thickness scales different for each section. See Figure 1 for additional keys. See Figures 9 and 10 for conodont genus abbreviations, and Sevatian sub-divisions. Ammonoid genus abbreviations, P. ¼ Psiloceras. I, M1, M2 negative isotopic excursions at Kujoch. C1, C2 possible correlation options between the Kujoch and St Audrie’s Bay sections (see text).

90

M. W. HOUNSLOW & G. MUTTONI

the same conclusion using the fragmented stratigraphy at Silicka´ Brezova´. Option A implies that the Newark interval E13–E14 interval is incomplete in comparison to the marine sections, and that the interval UT23r –UT24r in the GPTS is equivalent to E17r to E20r in the Newark Supergroup (Fig. 12). Option B (Fig. 12) and variations on this option have been proposed by Krystyn et al. (2002), Gallet et al. (2003) and option 1 of Gallet et al. (2007). The strength of this option is the good polarity pattern match between the lower part of the Passaic Fm. and the UT17 to UT20r interval in the GPTS. A weakness is the absence of a clear match to the reverse polarity dominated interval in the E18r – E20r interval in the Newark Supergroup (Fig. 12). Gallet et al. (2003) proposed correlations between the base of E21n and UT22n, whereas a variation shown in Figure 12 is a higher correlation between E21n and UT25n that is more consistent with the underlying reverse polarity dominated interval UT23r to UT24r (Fig. 12). This option tends to imply large sedimentation rates changes in the marine successions around the Sevatian– Rhaetian boundary interval. A dramatic reduction in sedimentation rate in the studied marine sections in the lowest Rhaetian, and large within-section sedimentation rate changes in the Sevatian, is a characteristic which is shown in the ST4/STK-A sections at Steinbergkogel (Fig. 10). In the nonmarine European sections the interval correlated to E18r to E20r (Fig. 11) certainly witnessed dramatic environmental changes related to much wetter environments and initiation of marine transgressions. This may be reflected in the Tethyan pelagic sections around this interval by large reductions in sedimentation rates. Magnetozone UT20r has been an attractive target for correlation to the E18r –E20r interval (i.e. option C in Fig. 12) because it has relatively the thickest reverse magnetozone in the Pizzo Mondello, Kavur Tepe and Kavaalani sections. This option has been proposed by Hounslow et al. (2004) and option 2 of Gallet et al. (2007). The correlation in option C implies that both the UT17 –UT20n interval in the marine sections and the E21 to E22 interval in the Newark Supergroup are incomplete (Gallet et al. 2007). Support for this option comes from: (a) the absence of typical European latest Triassic miospores from the Newark Supergroup (Van Veen 1995; Kuerschner et al. 2007), although this may be a reflection of the differing floral province of the Newark Basin; (b) the mid Norian GPTS through the interval UT17 – UT20n is the most fragmented, and therefore might be expected to be incomplete; and (c) the only substantial fault in the cored Passaic Fm. appears to be in magnetozone E22n (Olsen et al. 1996).

Weaknesses of correlation option C are: (a) it lacks the additional thin normal magnetozones characteristic of the E18r–E20r interval in the Newark Supergroup; (b) within the Blue Anchor Fm. in the UK the dinoflagellate cyst Rhaetogonyaulux rhaetica is known from lower, reverse polarity, levels (Orbell 1973) than those reported by Hounslow et al. (2004). In Tethyan sections in the Alps, R. rhaetica appears to characterise the middle and upper Rhaetian (Krystyn et al. 2007a), suggesting that the base of the Rhaetian probably lies within or below the level of SA4r and its correlative interval E19r– E20r (Figs 11 & 12). There is no current resolution to these correlation problems, so the cyclostratigraphy timescale from the Newark Supergroup cannot be easily applied to the GPTS (Fig. 14) through the Norian and Rhaetian. The fact that other nonmarine sections through this interval seem to confirm the polarity character from the Newark Supergroup below E21n (Fig. 11) suggests that the problems largely reside with the marine section data: either missing or duplicated intervals or large withinsection changes in sedimentation rates. A common assumption made for the Newark Supergroup magnetostratigraphy is that because of the very large sedimentation rate, it is the most complete record of the magnetic polarity in the Norian– Rhaetian (Gallet et al. 2007). This assumption is only valid if the coring obtained a complete succession. In the Newark Supergroup coring program, intercore-correlation was supported by ground mapping along with horizon correlation based on lithology and colour, supported by magnetostratigraphic correlation from core and limited outcrops (Olsen et al. 1996). Nevertheless, full succession recovery can be difficult to confirm in cyclically bedded red-beds like the Passaic Fm., because of small faults, unless very good well-log coverage and seismic surveys exist.

The Triassic – Jurassic boundary The Triassic– Jurassic boundary (TJB) is proposed to be defined in the Kuhjoch section in Austria (von Hillebrandt et al. 2007) at the FO of the ammonoid Psiloceras cf. spelae. Since this or nearby sections have no magnetostratigraphy and no sections containing this ammonoid have a magnetostratigraphy, the identification of this boundary in other sections with magnetostratigraphy, such as St Audrie’s Bay and Oyuklu, is based on other correlation criteria. The two best possibilities are the use of carbon-isotopic curves, and palynological changes near the boundary, which demonstrate correlation to the St Audrie’s Bay magnetostratigraphy. The organic carbon isotopic data at Kuhjoch and St Audrie’s Bay are quite similar, both display a

GEOMAGNETIC POLARITY TIMESCALE

91

Cs. timorensis Ns. triangularis

Ns. dieneri

H. typicalis H. parvus

Conodont ranges wrt chrons

Cs. gondolelloides

I. staeschei Ng. meishanenesis Ns. waageni

Tr. ex gr homeri

Polarity

S-5

H.

he

de

ns

tro

em

i

251

ANISIAN 250

249

247

248

~250.5 Ma

247.2 Ma

249

250

PAZ

Aegean

OLENEKIAN 251

251.2 Ma

252

S-4

S p a t h i a n

Chrons AZ

252

252.6 Ma

MT3

C. decipiens

INDUAN

K. subrobustus

S-3

S-2

S m i t h i a n

MT2

G. taimyrensis

Changhsingian

LT 9

LT 8

S. spitzbergense S.Z

N

Griesbachian Dien.

P. grambergi

E. romunderi

N. contrarium

L-T

A. tardus

O

P(S-1)

LT 7

LT 6

B. euomphala

V. s v e rdr u piP. ca ndidu s B. ro s e n krantz striga i/ O. communtus e

O. boreale

LT 5

LT 4

LT 3

L. caurus

LT 2

1

n.1r

O. concavum

S-L

LT 1

Substage Stage

Radiometric ties & age Buntsandstein 247 ATS

246.8 Ma

248

Polarity 1

MT3

2

Khara ulakh.

Rotelliforme

Olenekian

Kock.

Paracroch.

Aegean

Nevadanus

Balatonit. Trino. Pelsonian

Bith.

MT8

MT9

Oleshkoi

Constantis

MT7

MT6

5

Seced.

Reitzi

246

Mc

Kru

Mc

rni

Gredleri

Lin

Con

lea

gi

.

UT1

MT13

MT12

dst

nell

i

241

243.3 Ma

mi

Substage

CARN. 237

239

240

241.2 Ma

roe

Longobardia n

LADINIAN 242

243

Chrons

Regoled. Daxatina

Arch.

F a s s a n i a n

Illyrian

244

245

Ner aen

Curionii

ANISIAN 247

246.8 Ma

MT11

MT10

AZ

Ta im Cau Dec yr. rus ip.

MT4

Radiometric ties & age

236

238.0 237.3 Ma Ma

Stage

~235.5 Ma

Polarity

Aonoides

Austriacum

Tuvalian 1 Dilleri

(II)

Tuvalian 2

Tuvalian 3

Subbullatus

Spinosus

1 Jandianus

Arch. Regoled. Daxatina

(I)

(II)

(II)

Julian 2

237

235

238.0 Ma 237.3 Ma ~235.5 Ma

233

234

232

231

UT16

UT17

2

3

L a c i a n

UT18

Alaunian

Paulckei

CARNIAN

LADIN.

15

(I)

(II)

14

(I)

Julian 1

UT10 UT11 UT12 UT13

9

12

UT2 UT3 UT4 UT5 UT6 UT7 UT8

UT1

MT13

229

227 225 223

C.bZ & Substage AZ

NORIAN 230

Chrons

221

220

Stage 219 218

Radiometric ties & age

217

~ 230.9 Ma

Polarity

UT17

UT18

UT19

UT21

UT20

217

24 bi. - post.

UT25

UT26 UT27 post. - hern. M. rhaetica M. ultima

UT28 Chron C.bZ &

S e v a t i a n

Substage

RHAETIAN 204

203

HETT.

219

UT23 2

NORIAN

1

3

2

1

3

2

A l a u n i a n

Lac.

UT22

202

201.6 Ma

Stage

Radiometric ties & age

Fig. 14. Summary bio-magnetostratigraphic timescale for the Triassic, based on joint scaling of Figures 1, 2, 3, 5, 7, 8, 9, 10 and 12. No attempt is made to standardize the linked biostratigraphic zonations, which simply show the character of the supporting biostratigraphy (see text and previous figures for details). Radiometric ages from Brack et al. (2001), Mundil et al. (2004), Furin et al. (2006), Lehrmann et al. (2006), Galfetti et al. (2007a), Schaltegger et al. (2008), with additional discussion in text. Age in 1 Ma increments based on linear interpolations of radiometric ages for Induan to Carnian. Buntsandstein astronomical timescale (ATS) from Szurlies (2007), based on Figure 4 correlations. See Figure 1 for key. C.bZ, conodont biozones; AZ, Ammonoid zones; PAZ, Miospore assemblage zones (Lower Triassic from Hounslow et al. 2008a, b). See Figures 8, 9, 10 and 12 for details of Upper Triassic bio-zonations.

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dual-peaked initial negative isotopic excursion, prior to the first Psiloceras (marked ‘I’ in Fig. 13). At Kuhjoch, this initial dual peak is concentrated in the 10 cm at the base of the Tiefengraben Mb (von Hillebrandt et al. 2007), whereas at St Audrie’s Bay, it extends through the upper part of the Lilstock Fm. (Hesselbo et al. 2002; Fig. 13). In both sections, this initial negative peak includes the LO of conodonts. Above this level both sections show a peak in positive d13Corg values (at the position of magnetozone SA5r at St Audrie’s Bay and within the Schattwald beds at Kuhjoch; Fig. 13), followed above by a decline to more negative values, with Psiloceras cf. spelae at Kuhjoch about 0.5 m above a peak in negative d13Corg (marked M1 in Fig. 13), and coinciding with a second smaller positive peak in d13Corg (Fig. 13). Using only the isotope record for correlation suggests that the TJB proposed at Kuhjoch is approximately coincident with the first P. planorbis at the St Audrie’s Bay section (correlation C2 in Fig. 13). At Kuhjoch, above the TJB, there is a second negative excursion in d13Corg (marked M2 in Fig. 13), which appears to be shown at St Audrie’s Bay within the lower range of P. planorbis (Fig. 13), although this later peak is not shown in other TJB interval isotopic records (McRoberts et al. 2007). Kuerschner et al. (2007) suggested a correlation at a slightly lower level using the FO of Cerebropollenites thiergartii (correlation C1 in Fig. 13). The last occurrences of other significant miospore species (e.g. Ovalipollis pseudoalatus, Rhaetopollis germanicus and Ricciisporites tuberculatus) fall at different correlated levels, with only R. germanicus showing a LO within the main positive peak in the isotopic curve (Fig. 13). At Kuhjoch, the boundary between the TPo (Trachysporites–Porcellispora) and TH (Trachysporites –Heliosporites) zones is similar to that near the top of the Lilstock Fm., with both displaying an abundance peak in Kraeuselisporites (Heliosporites) reissingeri (Hounslow et al. 2004; Kuerschner et al. 2007), although the timing of these events appears to be different with respect to the carbon isotope data (Fig. 13). The Malanotte Fm. in the Lombardian Basin displays a similar association to that seen at St Audrie’s Bay (Fig. 13), with an initial peak in negative d13Ccarb and an acme of K. reissingeri a little above (Galli et al. 2007). Support for the lower correlation level (i.e. C1 in Fig. 13) is consistent with a wider set of much debated ammonite data. Whilst global correlation of species of Psiloceras is difficult, P. planorbis is commonly inferred to be age-equivalent or younger than the Psiloceras pacificum ammonite faunas (Guex et al. 2002; von Hillebrandt et al. 2007). Similarly, the ammonites of the genus Neophyllites, occurr prior to P. planorbis in NW Europe, but in

the New York Canyon area (Nevada, USA) it overlaps the range of Psiloceras tilmanni group ammonites (Guex et al. 2002). The lower correlation level (i.e. C1; Fig. 13) implies that the M1 isotopic negative excursion at St Audrie’s Bay occupies about 3 m of strata across the TJB at Kuhjoch. The magnetostratigraphy of the Montcornet core from the Paris Basin can be easily related to that from St Audrie’s Bay in that a detailed palynostratigraphy exists for both sections near the boundary (Yang et al. 1996; Hounslow et al. 2004). The upper boundary of the Argille de Levollois at 1075 m in the Montcornet core is closely coincident with a palynological change, very similar to that at the Lilstock Fm.–Lias Group boundary at St Audrie’s Bay. Both sections show abundance peaks of K. reissingeri followed closely above by low diversity miospore assemblages dominated by Classopollis. The LO of Ricciisporites tuberculatus also occurs slightly above the acme of K. reissingeri. For this reason, the reverse magnetozone at 1073.8 m in the Montcornet core is probably the equivalent of SA5r, some metres below the likely position of the TJB at St Audrie’s Bay (Fig. 13). Between 1075 m and 1067 m in the Montcornet core are acmes of K. reissingeri (1074.9– 1074.23 m), Deltoidospora (1074.9–1067.8 m), and Concavisporites (1074.9– 1070.1 m). As in the St Audrie’s Bay section, this acme interval appears to be the equivalent of the upper part of the TPo or lower part of the TH assemblage zones of Kuerschner et al. (2007). Other thin (,0.3 m) reverse magnetozones occur higher in the Hettangian part of the Montcornet core, but have not apparently been detected at St Audrie’s Bay, probably because of insufficient sampling density and the much thicker Hettangian succession (some 140 m) in the latter area. The best constraint on the correlation of the Oyuklu section magnetostratigraphy with that from St Audrie’s Bay, is provided by the LO of conodonts, which at St Audrie’s Bay have their highest occurrence close to the initial negative dC13 org peak. In some other sections, conodonts may range into younger strata closer to the positive peak in d13C (Lucas & Tanner 2007; McRoberts et al. 2007). At Oyuklu, the presence of the ammonite Phylloceras some 6 m above the last conodonts is the only solid evidence of truly Jurassic strata. The highest of the reverse magnetozones at Oyuklu (i.e. J2; Fig. 13) is probably the equivalent of SA5n.1r at St Audrie’s Bay, and BT4r at the Brumano section. The Rhaetian nature of the Williton Mb. and Westbury Fm. in St Audrie’s Bay is shown by the dinoflagellate cyst Rhaetogonyaulax rhaetica, whose first occurrence appears to be younger than Sevatian-1 (Krystyn et al. 2007a), although its first occurrence is often strongly influenced by

GEOMAGNETIC POLARITY TIMESCALE

environmental conditions in the Germanic facies rocks of the UK, France and Germany. There has been much discussion about correlating from marine successions within the TJB interval into the thick nonmarine successions of eastern North American (Kent & Olsen 1995; Muttoni et al. 2010; Hounslow et al. 2004; Whiteside et al. 2007; Lucas & Tanner 2007; Gallet et al. 2007). The most recent synthesis places the TJB interval somewhere within the succession of interbedded basalts and sedimentary units above the Passaic palynofloral event (Lucas & Tanner 2007). A magnetostratigraphic constraint is provided by the good match between the lower part of the St Audrie’s Bay succession and the Newark Supergroup-Hartford Basin magnetozones E14 to E20 (Fig. 11), which suggests Newark Supergroup magnetozone E23 is Rhaetian (Fig. 13). A further constraint is provided by the magnetostratigraphy of the Portland Fm. in the Hartford Basin, overlying the basalt succession, which based on magnetostratigraphic correlation, probably places the base of the Sinemurian around magnetozone H26r (Kent & Olsen 2008). Around the Passaic palynofloral event, Whiteside et al. (2003) have an associated negative peak in d13Corg which is similar to that from marine successions. An alternative, supported by the magnetostratigraphy, is that this event may be a negative excursion in d13Corg equivalent to that in the Westbury Fm. close to SA5n.1r (Fig. 13). This suggests that Newark Supergroup magnetozone E23r may be the equivalent of BT5n.1r in the Italcementi section and UT27n.3r in the composite GPTS, a correlation which is consistent with the conclusions of Lucas & Tanner (2007).

Conclusions The duration of the Triassic is some 51.1 Ma using the Changhsingian–Induan boundary at 252.6 Ma (Mundil et al. 2004) and the Rhaetian –Hettangian boundary at about 201.6 Ma (Schaltegger et al.

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2008). Using linear extrapolation in the pseudoheight composite (Fig. 14) the base of the MT magnetochrons are very close to 247.2 Ma, essentially those radiometric dates from the Guando sections (Lehrmann et al. 2006). The age for the base of the UT magnetochrons is approximately 235.5 Ma based on the arguments presented previously for the Middle Triassic (Fig. 5). Linear interpolations (using the pseudo-height composite) of the radiometric ages have been used to add 1 Ma increments for the Induan through to the early Norian GPTS (Fig. 14). Those for the upper part of the Rhaetian are based on the Newark Supergroup APTS, which are constrained by a similar upper tie-point at c. 201.6 Ma (Schaltegger et al. 2008). This cannot be usefully performed for the Norian, since its c. 25 Ma duration is not well constrained with radiometric ages, and there is no certainty in how to best correlate the Newark APTS to the marine-based GPTS. The GPTS for the Triassic has some 133 magnetozones that appear to be soundly validated by existing data, but with some 37 additional tentative submagnetozones (Fig. 14). We have divided these into 50 magnetochrons corresponding to major N –R couplets. The validated magnetozones give a reversal rate of 2.6 rev/Ma, and average magnetochron duration of 0.38 Ma (Table 1). This reversal rate is similar to that in the Cenozoic. The Cande & Kent (1995) timescale has 171 reversals from the base of magnetochron C29n at 64.745 Ma to the base of C1n at 0.78 Ma, yielding a mean reversal frequency of 2.64 rev/Ma with a mean magnetochron duration of 0.379 Ma. The Lower and Middle Triassic have similar reversal rates of c. 4 rev/Ma, but the Upper Triassic has a reversal rate which is approximately half of this (Table 1), indicating that the maximum magnetostratigraphic resolution available for dating and correlation is during the Lower and Middle Triassic. The proposed and ratified Triassic stage boundaries are for the most part reasonably well

Table 1. Statistical information about the Triassic magnetic field divided into intervals corresponding to the chron numbering and radiometric age scaling used in Figure 14. Statistics are shown for validated polarity boundaries, and additional tentative, often short duration magnetozones. [..] indicates the number of magnetozones Chron Interval

UT MT LT Triassic

Age range (Ma)

33.9 11.7 5.5 51.1

Validated magnetozones

Including tentative sub-magnetozones

Reversal rate (rev/Ma)

Mean chron duration (Ma)

Reversal rate (rev/Ma)

Mean chron duration (Ma)

1.9 [64] 3.8 [45] 4.4 [24] 2.6 [133]

0.53 0.26 0.23 0.38

3.0 [102] 4.4 [51] 9.8 [54] 4.1 [207]

0.33 0.23 0.10 0.25

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characterised by a known and validated magnetic polarity pattern. The Induan, Anisian and Ladinian are perhaps the best characterized with multiple studies, along with the lower Norian and upper Carnian. In the Lower Triassic the lower part of the Olenekian appears to be not strongly validated by data from multiple sections, which in part may relate to inadequate biostratigraphic constraints. The Middle Triassic GPTS is for the most part well characterized by conodonts and secondarily by ammonoids, with low to high palaeolatitude correlations of the biozones supported by magnetostratigraphy (Hounslow et al. 2008b). Conodont biozonations provide the primary means of age calibrating the Upper Triassic GPTS. Parts of the Upper Triassic GPTS are not strongly validated in multiple sections with the existing data. The GPTS in the middle parts of the Carnian is the least well documented (UT5 to UT9), with intervals in the middle Norian (UT17 to UT20n) and middle Rhaetian (UT24 to UT26) possessing somewhat lesser degrees of uncertainty. For the Lower Triassic, the ages based on the radiometric control points, and the Buntsandstein astronomical timescale (ATS), are in some parts more than 1 Ma divergent. Some of this, particularly in the Olenekian, relates to the uncertainty in how to best correlate the magnetostratigraphy between the marine and nonmarine successions. Part of the reason for the ‘bunching’ of the 0.2 Ma intervals in the Buntsandstein ATS presumably relates to the likely out-of-phase sedimentation rates, between these interior continental basins (i.e. Buntsandstein) and continental margin (i.e. Sverdrup –Barents Sea) records of magnetic polarity. The apparent confirmation of the Newark Supergroup magnetostratigraphy by data from other nonmarine sections (Fig. 11) indicates the generally robust nature of the Newark Supergroup APTS. Therefore the problems in relating this to the GPTS are either omission (or duplication) in the marine Norian –Rhaetian section data, or that Upper Triassic, nonmarine clastic and marine carbonate successions have strongly out of phase sedimentation rates: problems that will tax future research. Both of the enormously detailed magnetostratigraphic studies on the Buntsandstein and Newark Supergroup indicate that without additional correlation constraints there will often be additional uncertainty in using such polarity records for age control. This suggests that better integrated, multi-tool studies will be required to provide more detailed and better understanding of environmental and sedimentary systems in the nonmarine Triassic. Jonathan Glen kindly allowed use of his unpublished/in press manuscript. Carmen Heunisch allowed use of unpublished data on the palynology of the Buntsandstein.

Tyler Beatty and Morten Bergan allowed use of inpress and unpublished data. Leopold Krystyn, Mike Orchard, Steve Hesselbo, and Dennis Kent provided helpful discussion. Detailed comments from Geoff Warrington were provided on an earlier draft. Jim Ogg and Robert Scholger improved the text with thoughtful reviews. Russian translations by Vassil Karloukovski.

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The Triassic isotope record LAWRENCE H. TANNER Department of Biological Sciences, Le Moyne College, Syracuse, NY 13214, USA (e-mail: [email protected]) Abstract: A variety of stable isotope measurements have been found useful in studying processes of environmental change. Measurements of d13C, d18O, d34S and 87Sr/86Sr all can provide information about the conditions of the water column in which sediment deposition occurred, but the most widely applied of these is d13C. The carbon isotope record for the Triassic System is a complex one; a pronounced negative excursion begins below the base of the Triassic System and continues into the basal Triassic. The succeeding 4 to 6 Ma Lower Triassic interval is marked by isotopic instability, with positive and negative excursions, continuing through the basal Middle Triassic. In contrast to the Lower Triassic, most of the Middle and Upper Triassic display relative isotopic stability, with rising values of d13C likely reflecting environmental recovery and increasing storage of organic carbon in terrestrial environments. The uppermost Triassic is marked by a pronounced negative excursion near the system boundary that has been linked to significant biotic turnover. The causes of the various excursions remain under investigation, particularly those at the system boundaries, with outgassing during volcanic activity, changes in productivity, ocean anoxia, and seafloor methane releases all suggested as mechanisms both for perturbing the global carbon cycle and for forcing biotic extinction.

In recent decades, sedimentary geology has benefited substantially from the techniques of measurement of a variety of stable-isotope ratios, each of which has its own special applications, limitations and pitfalls. Among the more commonly utilized are the isotopic ratios of 13C to 12C (d13C), 18O to 16O (d18O), 34S to 32S (d34S) and 87Sr/86Sr, all of which can provide information about the conditions of the water column in which sediment deposition occurred, and thus potentially inform our understanding of global-scale environmental changes. Although these studies also may be conducted on terrestrial sediments, for example, the carbon and oxygen isotope composition of calcretes and lacustrine carbonates (e.g. Cerling 1999), or the carbon isotopes of terrestrial plant remains (e.g. McElwain et al. 1999), most analyses of stable isotopes in sediments are performed on marine strata and marine fossils. It is the potential for these measurements to provide information on ocean water chemistry, salinity, oxygenation, productivity and temperature that makes them such attractive tools in the modern arsenal of analytical techniques. Major palaeontological boundaries attract the greatest research scrutiny, and so the majority of studies have centered on narrow stratigraphic intervals containing these boundaries with the intent to discern underlying causes, or at least the conditions that prevailed during biotic turnover. Consequently, fewer studies document long-term (i.e. system-long) secular trends. Prominent amongst these are the remarkable databases compiled by Veizer et al. (1999) and Prokoph et al. (2008) that examine the

record (from fossils) of d13C, d18O and 87Sr/86Sr for the entire Phanerozoic (in the former) and the entire geologic record (in the latter). Unfortunately, the resolution of these data are mostly insufficient to resolve questions of specific events at the stage level. The first dataset spanning the Upper Permian through Upper Triassic at a useful resolution was the d13C record constructed by Atudorei (1999). Korte et al. (2003) compiled a robust 87 Sr/86Sr dataset from Tethyan and Muschelkalk brachiopods and conodonts for the entire Triassic (Fig. 1). Subsequently, Korte et al. (2005) produced a similarly useful record from previously published and new measurements of d13C and d18O from brachiopod shell, conodonts and whole rock samples from Iran, Sicily and the Southern Alps, also covering the entire Triassic (Fig. 2). While not at the level of resolution of most other stage-level studies, these data provide valuable information on baseline secular trends of the isotopic composition of the Triassic ocean.

Carbon isotopes The most frequently utilized of the isotopic tools is the composition of carbon, d13C, which can be measured both from inorganic carbon, as carbonate from either whole rock or fossil material, and organic carbon, generally from whole rock analyses. Changes in the ratio of 13C to 12C in the ocean may result from changes in the fluxes and partitioning of carbon between organic and inorganic carbon

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 103– 118. DOI: 10.1144/SP334.5 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Korte et al. (2003) produced this data set of Sr isotopes from Tethyan and Muschelkalk brachiopods. High-Sr samples ¼ .400 ppm (and or Mn , 250 ppm); low-Sr ¼ ,400 ppm (and/or Mn . 250 ppm).

reservoirs, from differences between surface and deep water chemistry, or they may result from disturbances of the rates of carbon cycling through the ocean-atmosphere system (see review in Kump & Arthur 1999). The utility of d13C in these studies is enhanced by the presumed resistance of the isotopic composition of carbonate carbon to change by low-temperature diagenetic processes, although organic matter is considered much more resistant to post-depositional change than carbonate (Banner & Hanson 1990). Therefore, diagenesis can be ruled out where d13C displays parallel trends for both carbonate and organic matter (Kump & Arthur 1999). The introduction of large volumes of isotopically light carbon into the global carbon reservoir will produce parallel trends in organic and carbonate carbon, but there are multiple sources of light carbon. The eruption of large igneous provinces is, perhaps not coincidentally, associated with the isotopic excursion events associated with some palaeontologic boundaries, including the endPermian (the Siberian traps eruption), the endTriassic (Central Atlantic Magmatic Province or CAMP), the end-Cretaceous (Deccan Traps), and the LPTM (the North Atlantic Igneous Province; Wignall 2001). Sudden decreases in primary productivity, potentially forced by prolonged

volcanism, may result in a rapid buildup of light carbon in the carbon reservoir, and this mechanism has been offered to explain the shifts at the K –T boundary (Kump 1991) and the end-Permian (Holser & Magaritz 1992; Magaritz et al. 1992). The dissociation of methane hydrates in ocean-floor sediments has been proposed as a mechanism for the rapid introduction of large volumes of very light carbon (d13C ¼ 260 to 265‰) to the ocean-atmosphere system (Dickens et al. 1995, 1997). In theory, once the release of sea-floor methane is triggered, ocean warming and dropping of the thermocline may result in continued dissociation and a ‘runaway greenhouse’ effect (Dickens et al. 1995). The isotopic shifts at the end-Palaeocene clearly coincide with a warming event of 5 –7 8C, the late Palaeocene thermal maximum (LPTM). This event has been attributed to a massive methane release, possibly triggered by a change in oceanic thermohaline circulation (Dickens et al. 1995).

Oxygen isotopes The oxygen isotope composition (d18O) of carbonate is not considered nearly so robust a measurement as the carbon isotope composition as it is more

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Fig. 2. Korte et al. (2005) compiled previously published and original data to produce this record of d13C and d18O variation for the entire Triassic.

strongly affected by diagenesis (Marshall 1992). This is due to: (1) the potential for recrystallization of the carbonate in equilibrium with formation waters that have an isotopic composition substantially different from the original waters in which crystallization occurred; and (2) the strong sensitivity of isotopic fractionation to temperature during carbonate crystallization. However, this temperature sensitivity allows calculation of the temperature of the fluids from which the carbonate is precipitated. If assumptions are possible regarding the d18O of the water from which carbonate crystallizes, the mean annual temperature of the water can be related through the empirical functions of O’Neill et al. (1969) and Hays & Grossman (1991). These relationships allow examination

of long-term trends in ocean temperature (e.g. Prokoph et al. 2008). They also serve as a means for identifying diagenetic processes affecting the isotopic composition of carbonate; covariance of d13C and d18O is often taken as evidence of diagenetic equilibration of the carbonate with subsurface fluids (Marshall 1992), although this is not always the case, and petrographic confirmation is important (Veizer et al. 1999). Notably, it has often been assumed that the Phanerozoic ocean has maintained d18O ¼ c. 20.3‰SMOW, although the modern ocean in an ice-free world would have a composition closer to 21.4‰ (Jaffre´s et al. 2007). In fact, analysis of the isotopic record of whole rock and brachiopod carbonate across the Phanerozoic (and before)

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reveals a secular trend of increasing d18O from ancient to modern (Veizer et al. 1999; Jaffre´s et al. 2007). These data broadly suggest that the d18O of the Triassic oceans rose through the period, but that for most of the Triassic, the d18O of the ocean was lower than the modern. The data of Korte et al. (2005), from Tethyan and Muschelkalk brachiopods and whole rock carbonate demonstrate a similar trend, with d18O rising to values greater than 21.0‰ by Norian time, but dropping by 1.0‰ or more during the Rhaetian. Korte et al. (2005) argued that a mean ocean d18O of c. 20‰SMOW is required to allow for Late Triassic ocean MATs in the range of tolerance of reefbuilding corals.

Sulphur isotopes At its simplest, the sulphur –isotope (d34S) composition of seawater reflects the relative fluxes of sulphur sequestered in marine sediments versus the input of sulphate from weathering on the continents (see review in Bottrell & Newton 2006). Fractionation occurs during bacterial sulphate reduction, so reduced sulphur (e.g. pyrite) is isotopically depleted, while the sulphate remaining in the water column becomes proportionally enriched. Sulphur is also withdrawn from the ocean reservoir as sulphates in evaporite deposits, although flux rates for this mechanism vary widely across geological time. Thus, in a steady-state ocean, reduced sulphur lost to deposition (in pyrite) is replaced by riverine input containing sulphur derived from weathering of pyrite and/or evaporite-bearing sediments. Traditionally, studies of changes in d34S relied on evaporite deposits as an archive of seawater-derived sulphate. The disadvantage of this technique is that the geologic record of evaporite deposits tends to be somewhat discontinuous. Hence, the newer technique of analysis of sulphate bound with carbonate minerals, called carbonate-associated sulphate (or CAS), is now considered to have greater potential for producing continuous records of the sulphur isotope composition of the ocean (Bottrell & Newton 2006). Marenco et al. (2008) have cautioned that the d34S of CAS may be altered by dolomitization, and thus that petrographic studies should accompany sulphur isotope analyses.

Strontium isotopes The Sr-isotope ratio trend for the Triassic is wellpresented within the data of Veizer et al. (1999). Although the resolution is low, the trend is identical to that presented in much greater detail by Korte et al. (2003). Both studies demonstrate a steep rise in the ratio, beginning in the Late Permian and

continuing into the Early Triassic, reaching a peak during the late Olenekian, falling through the Anisian and Ladinian, before rising again during the Carnian and Norian, and falling sharply from the late Norian through the Rhaetian. The most detailed Sr-isotope data record for the Early Triassic was produced by Huang et al. (2008); their results support the trend identified by Korte et al. (2003) of 87Sr/86Sr rising from 0.70714 at the Permian– Triassic boundary to 0.70823 at the top of the Spathian. As Sr-isotope ratios in seawater are controlled in large part by the rate of continental weathering, Korte et al. (2003) interpreted the Sr-isotope ratio peaks as the result of humidity enhanced weathering, in the Late Permian through Early Triassic, and a combination of humidity and orogenic uplift, in the Carnian to Norian.

Lower Triassic Permian –Triassic boundary All discussions of the isotope record of the Lower Triassic necessarily begin with the events at the Permian– Triassic boundary. The age of the system boundary at the GSSP at Meishan (southeastern China) is established by radio–isotopic dating of ash beds in the section; the U/Pb date of 252.6 + 0.2 Ma (Mundill et al. 2004) is indistinguishable from the 40Ar/39Ar date of 249.25 + 0.14 Ma (Reichow et al. 2009) when the latter is adjusted for 40K decay constant bias. The presence of a well-defined negative carbon isotope excursion (CIE) at the boundary in both d13Corg and d13Ccarb has been long-known from marine strata at localities in the Tethyan and Neotethyan realms and regions bordering Panthalassa, in addition to South China (Holser & Magaritz 1987; Baud et al. 1989, 1996; Holser et al. 1989; Magaritz et al. 1992; Xu & Yan 1993; Kajiwara et al. 1994; Wang et al. 1994a, b; Wignall et al. 1995, 1998; Atudorei 1999; Jin et al. 2000; Krull et al. 2000, 2004; Dolonec et al. 2001; Stemmerik 2001; Twitchett et al. 2001; Cao et al. 2002; Krystyn et al. 2003; Sarkar et al. 2003; Payne et al. 2004; Richoz 2006; Galfetti et al. 2007a; Horacek et al. 2007a). The negative CEI has been recognized also in terrestrial sediments (Morante 1996; Krull & Retallack 2000; MacLeod et al. 2000; Twitchett et al. 2001; de Wit et al. 2002). Exact stratigraphic placement of the excursion depends on the location of the system boundary in the section being measured, that is, the FAD of the conodont Hindeodus parvus (which varies), but many workers now also accept the excursion itself as a chemostratigraphic marker for the system boundary (Krull et al. 2004); the base of the

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Fig. 3. Comparison of d13Ccarb data from widespread localities demonstrates global nature of the disturbances of ocean d13C from the Chingxingian (latest Permian) through the earliest Anisian (Middle Triassic), and the relative stability that succeeded in the Middle to Late Triassic. The South China data are a composite curve presented by Payne et al. (2004). The curve from the Dolomites is adapted from Horacek et al. (2007b), and the Iranian data are from Horacek et al. (2007a).

excursion occurs in the latest Permian and extends into the earliest Triassic (Corsetti et al. 2005). Generally, the excursion is marked by a pronounced drop of between 2 and 4‰ from latest Permian values (Fig. 3), and anomalously large excursions initially reported for the GSSP section at Meishan (Chen et al. 1991; Xu & Yan 1993) are now considered a diagenetic artifact (Jin et al. 2000; Krull et al. 2004). Although there is now widespread agreement on the placement and magnitude of the excursion, the duration has been debated intensely, with ranges of 104 to 106 years suggested (Corsetti et al. 2005). As described by Krull et al. (2004), for example, data from two shallow marine platform sections in the Nanpanjiang Basin, south China, display distinct negative excursions of both d13Corg and d13Ccarb beginning below the top of the Changhsingian (uppermost Permian) and continuing into the Griesbachian (lower Induan), with most of the excursion in the basal Griesbachian (in the Hindeodus parvus Zone). This excursion is coincident with a facies change (calcimicrobial framestone) and a sharp drop in species abundance and diversity, although most of the stratigraphic range of the

excursion is above the actual level of the boundary extinction event. The claim also has been made that the negative CEI at the system boundary can be recognized in abrupt shifts in d13C of pedogenic carbonate in paleosols spanning the boundary (Ward et al. 2005; Retallack et al. 2007). Explanations for the negative CIE have been linked necessarily to the biotic events spanning the Permian –Triassic boundary, and have included: overturn of a deep, anoxic ocean, which would have released isotopically light deep-ocean CO2 and potentially caused anoxia (Wignall & Hallam 1992; Horacek et al. 2007a) or hypercapnia (Knoll et al. 1996); the cessation of bioproductivity (Rampino & Caldeira 2005); productivity collapse following a bolide impact (Retallack et al. 1998; Becker et al. 2001; Kaiho et al. 2001); the catastrophic release of methane hydrates from oceanfloor sediments (Erwin 1993, 1994; Bowring et al. 1998; Krull & Retallack 2000; Krull et al. 2004); and environmental stresses associated with the extrusion of the Siberian Traps (Knoll et al. 2007). This last factor has received considerable attention because of the temporal coincidence of the boundary extinction and the eruptions; Kamo et al.

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(2003), for example, reported a U/Pb zircon age of 251.7 + 0.4 Ma from lavas the Siberian Traps, and Reichow et al. (2009) have produced a 40Ar/39Ar age of 250.3 + 1.1 Ma from the upper part of the Siberian volcanics. Payne and Kump (2007) proposed that the negative CIE records a short-term release of isotopically light (d13C ¼ 225‰) carbon by thermal metamorphism of coals during the initial eruption of the Siberian Traps basalt. In their model, the release of 2  1018 mol of organic carbon (d13C ¼ 225‰) over a span of 20 ka would cause a negative CIE of about 3‰, and the same volume released over an interval of 100 ka would result in a CIE of nearly 4‰. Berner (2002) used a mass-balance approach to examine the problem and concluded that no single mechanism operating in isolation was likely to account for the observed isotopic record, but that multiple processes operating synergistically could. For example, an impact event potentially triggered both the release of isotopically depleted methane (13C ¼ 265‰) stored as clathrate, and Siberian Trap volcanism, which released volcanogenic CO2, which in turn triggered mass mortality and the release of depleted organic carbon. Payne et al. (2004) calculated that a negative CEI of this magnitude would require the release of much more than 104 Gt of methane. Notably, however, Berner’s calculations were based on the assumption of the high amplitude excursion (8‰ or greater) originally measured at Meishan, which is now considered partly an artifact of diagenesis. Knoll et al. (2007) argued that the pattern of selective survivorship across the boundary mandates the operation of multiple mechanisms, including global warming, ocean anoxia and hypercapnia, with Siberian Traps volcanism as the most likely common trigger. Environmental disturbance across the Permian– Triassic boundary is recorded also by dramatic changes in sulphur isotopes (d34S), reflecting disturbance of the global sulphur cycle. The most complete record of d34S across the Permian–Triassic boundary is that presented by Newton et al. (2004). These authors produced a d34S record from CAS across the Permian–Triassic boundary in northern Italy in which d34S rises in the Changhsingian from an exceptionally low value of 11.5‰ to a high of 26.9‰ just above the extinction horizon, before falling to 15‰ and fluctuating c. 18‰ at the system boundary, and rising again in the Griesbachian to 25.3‰. The authors suggested that the strong increases in d34S record episodes of ocean anoxia, during which increased bacterial sulphate reduction removed isotopically light sulphur from the water column, thus explaining extinction at the system boundary and the delayed recovery during the Early Triassic (Newton et al.

2004). This interpretation is supported by the geochemical modelling of Berner (2005).

Griesbachian and later Considerable attention has been devoted to the Early Triassic recovery from the Permian–Triassic extinction. Chen et al. (2007), for example, described the pronounced recovery of d13C values at the Meishan section (south China) from 22.0‰ at the boundary to þ2.0‰ by middle–late Griesbachian, and attributed this strong isotopic shift to increased ocean productivity. Furthermore, analysis of more extended Lower Triassic sections has revealed that the CIE at the Permian–Triassic boundary was not the result of an isolated event, but actually the first of a series of abrupt and profound variations in both d13Corg and d13Ccarb. As multiple authors (e.g. Atudorei 1999; Corsetti et al. 2005) have described, the CIE at the Permian– Triassic boundary is just the onset of a period of instability lasting 4 to 6 million years, characterized by strongly oscillating values of d13C that are correlative in sections from China to India to Oman. For example, Horacek et al. (2007a), in Iranian sections, described the following: (1) a Griesbachian increase in d13C, similar to that noted by Chen et al. (2007); (2) a short negative excursion at the Griesbachian–Dienerian boundary; (3) a rise during the Dienerian to a major positive peak near the Dienerian –Smithian (Induan–Olenekian) boundary; (4) a negative excursion during the late Smithian; (5) a subsequent rise and positive peak in the early Spathian, just above the Smithian–Spathian boundary; (6) decreasing values during the Spathian; and (7) a final rise and positive peak just above the Spathian –Anisian boundary (Fig. 3). The data of Payne et al. (2004) from south China, and Korte et al. (2005) from Iran, Sicily and the Southern Alps yield a nearly identical pattern of isotopic oscillations. Thus, significant positive excursions have been measured from widespread sections at the Dienerian –Smithian and Smithian– Spathian boundaries, and a positive peak occurs in the lower Anisian, with the sole prominent negative excursion occurring during the Smithian (see review in Payne & Kump, 2007). For example, Richoz et al. (2007a) observed the positive CIE at the Induan–Olenekian (Dienerian –Smithian) boundary in the Mud section at Spiti, in the Himalayas, as did Tong et al. (2002) in south China. Horacek et al. (2007b) found that d13C increased by as much as 6‰ at the Dienerian –Smithian boundary in the Tethyan realm (as measured in the Dolomites), and Horacek et al. (2007c) verified that in the Chaohu section (south China) d13C increases to the top of the Induan and produces a pronounced positive

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excursion at the stage boundary, but drops in the early Olenekian, with a pronounced negative excursion in the Smithian, before rising sharply at the Smithian –Spathian boundary. Galfetti et al. (2007a) attributed the drop (from þ2‰ to 21‰ in the Loulou Formation, South China) to the influence of isotopically light volcanogenic CO2. Galfetti et al. (2007a) noted that the prominent positive CIE that spans the Smithian –Spathian boundary appears to be recorded globally and may coincide with organic-rich shale deposition during a eustatic regression-transgression cycle. Galfetti et al. (2007b) emphasized the coincidence of this CEI with major ecosystem changes, marked in the marine realm by a major ammonoid turnover event and on land by the rise of woody gymnosperms and the decline of lycopods. The authors postulated that these changes were driven by a massive injection of CO2 from the Siberian Traps eruptions. The duration of this magmatism is not known precisely, but pulses potentially extended to c. 240 Ma (Ivanov et al. 2009). Galfetti et al. (2007b, c) speculated that high CO2 flux from magmatic episodes caused sudden warming, and reduced pH of marine waters, resulting in a calcification crisis. Korte et al. (2005) noted that the Smithian –Spathian boundary also coincides with the end of the radiolarite gap that initiated with the Permian–Triassic boundary. Hence, either increased primary production and/or ocean anoxia drove the increased burial of isotopically light carbon, resulting in a positive CIE (Galfetti et al. 2007c). The positive CIE that spans the Spathian– Anisian boundary is also global in extent, and as with the excursion at the Smithian –Spathian boundary, may be related to increased organic carbon burial (Galfetti et al. 2007a). Corsetti et al. (2005) examined the potential causes of the Lower Triassic isotopic oscillations and found that the massive release of clathrates from the ocean floor, proposed by some (e.g. Krull & Retallack 2000; Heydari et al. 2008) as the cause of the CIE at the Permian –Triassic boundary, was unlikely to have caused the repeated fluctuations recorded in the post-boundary Triassic. These authors based this conclusion on the assumption that extended time (estimated at 10 Ma) is needed to recharge the methane reservoir. Furthermore, the symmetry of the profiles of the Lower Triassic CIEs does not match the asymmetry predicted for this mechanism. Consequently, the authors explored the possibility that a strongly stratified ocean that underwent episodic overturn caused the strong fluctuations of the isotopic composition of the ocean. Corsetti et al. (2005) noted, however, that the pattern of oscillation, whatever the cause, is superimposed on a long-term negative shift of the isotopic composition of the ocean, probably driven by a

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fundamental reorganization of the carbon cycle; during the Early Triassic ‘coal gap’ (Berner & Canfield 1989) the high-biomass flora of the Late Palaeozoic was replaced by a low biomass flora, with the consequence that the proportion of isotopically light carbon buried in terrestrial environments decreased substantially (Corsetti et al. 2005). Similarly, Payne & Kump (2007) studied the composite 5 Ma record of fluctuating d13C from the Permian –Triassic boundary through the lower Anisian, and also ruled out clathrates as a forcing mechanism of isotopic fluctuation. Their simple carbon cycle box model suggested that repeated episodes of volcanism, by themselves, are capable of producing only very modest negative excursions in the composition of marine carbonate. However, the outgassed CO2 triggered a feedback cycle of warming, enhanced weathering and enhanced productivity (stimulated by phosphate recycling) that resulted in significant increases in the rate of ocean carbon burial, thereby producing a substantial positive excursion in the composition of marine carbon. The system then returned to equilibrium because carbonate burial rose above steady state conditions. Payne & Kump (2007) thus proposed that continuing eruptive episodes of the Siberian Traps triggered continuing fluctuations in the marine carbon reservoir during the Early Triassic.

Middle Triassic By comparison to the strongly oscillating isotopic record of the Lower Triassic, the Middle Triassic was relatively calm. The d13Ccarb data of Korte et al. (2005), derived from brachiopod and whole rock data, are not of very high resolution. Nonetheless, their data display a very pronounced positive CIE at the very base of the Anisian. d13Ccarb values are stable, but depleted through most of the Anisian and Ladinian, beginning to rise near the top of the Ladinian. In considering the cause of the light isotopic composition of Middle Triassic carbonate, the authors noted the contradictory evidence for rising sea level and increased carbon burial in shelf settings, which should result in heavier isotopic values, and low carbon burial on land (the so-called Triassic coal gap), which should allow d13C to remain low. The authors go on to speculate that enhanced volcanic activity during this interval released sufficient magmatic CO2 to depress d13C. Payne et al. (2004) presented a composite section of data from south China that, similar to Korte et al. (2005), displays a relatively stable d13Ccarb curve following the basal Anisian positive excursion up through the basal Carnian (Fig. 3). These data demonstrate isotopic enrichment that characterized the uppermost Ladinian

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through Carnian portion of the Korte et al. (2005) curve. The Middle Triassic d18O data of Korte et al. (2005), also primarily from brachiopod and whole rock carbonate, suggest a broad trend of rising d18O values, although significant scatter obscures the trend. The authors suggested that much of the scatter reflects local changes in salinity or temperature (e.g. upwelling). Indeed, Preto et al. (2005) examining nodular carbonate in correlative sections of Anisian age in the Dolomites, found that d18O fluctuated significantly about a relatively stable mean value of 23.0‰, while d13C rose c. 1.5‰ across the same interval.

Upper Triassic Carnian through the Rhaetian The data of Atudorei (1999), Payne et al. (2004) and Korte et al. (2005) indicate that d13C values began to rise in the uppermost Ladinian, or possibly slightly earlier. d13C continued to rise through the Carnian, as carbon isotope values rose by an average of almost 3.5‰ by the base of the Norian (Korte et al. 2005). Relatively high values persisted through the lower to middle Norian, before decreasing in the upper Norian and Rhaetian. Increased sequestration of organic carbon, as coal deposition recommenced in the Late Triassic, was cited by Korte et al. (2005) as the cause of the positive shift in the Carnian. In discussing the middle Norian decline in d13C, the authors noted facies changes to more red-bed dominated sedimentation in a number of basins, and suggested a decline in terrestrial carbon burial. In more detailed analysis, Richoz et al. (2007b) analysed a section across the Ladinian–Carnian boundary at Weissenbach (Austria) in the northern Calcareous Alps, and found that d13C values averaged nearly þ3.0‰ through the section, although with two short negative deflections of c. 0.5‰ on either side of the boundary. The data of Korte et al. (2005) indicate a substantial drop (c. 1.5‰) in d13Ccarb during the mid-Carnian, commencing during the late Cordevolian and continuing through the Julian, before rising at the start of the Tuvalian. Hornung et al. (2007) noted the coincidence of this isotopic shift with a Tethyan-wide ‘carbonate productivity crisis’. These authors attribute the carbonate decline to the well-known ‘Carnian pluvial episode’ (Simms & Ruffel 1990), which produced increased siliciclastic runoff due to enhanced weathering, and the isotopic decline to the dissolved inorganic carbon carried to the ocean by the increased runoff. Muttoni et al. (2004) found evidence of an abrupt and long-lasting positive shift in d13Ccarb (.1.0‰)

just below the Carnian –Norian boundary at the Pizzo Mondello section in Sicily. Nicora et al. (2007), analysing the same section, obtained similar results. Williford et al. (2007a), however, analysed a Carnian –Norian boundary section from Black Bear Ridge, British Columbia (Canada) and concluded that d13Corg values remained stable across this interval. Sephton et al. (2002) presented d13Corg data from the Black Bear Ridge section that suggested that a significant and extended positive excursion occurs at the Norian– Rhaetian boundary. Furthermore, these authors suggested that biotic extinctions at this stage boundary resulted from conditions of a stratified ocean, sluggish circulation, and a low latitudinal gradient that caused ocean anoxia. Ward et al. (2004) critiqued this work subsequently, noting that the Rhaetian stage section analysed by Sephton et al. (2002) amounts to only 10 m and thus is severely condensed. Nonetheless, Ward et al. (2004) also interpreted a positive CIE in d13Corg from data obtained from the section spanning the Norian –Rhaetian boundary at Kennecott Point, Queen Charlotte Islands (British Columbia), and noted that the interval in which they interpreted this excursion corresponds with the interval in which the pectinid Monotis is severely reduced. A lithologic change from bioturbated to thinly laminated facies occurs in this interval, suggesting again that increased ocean anoxia is the mechanism responsible for the isotope enrichment. Notably, the data of Ward et al. (2004) display a complex pattern of short-term isotopic fluctuations of significant magnitude that contrasts with the simple, prolonged excursion interpreted by Sephton et al. (2002). Williford et al. (2007b) re-analysed the section at Kennecott Point in greater detail. Their results again display great short-term variability, but do not indicate any pronounced excursion in one direction or the other. Tanner et al. (2006, 2007) examined the d13Ccarb record of strata spanning the Norian–Rhaetian boundary in the Lagronegro succession, southern Apennines. At the section near Pignola, Tanner et al. (2006) found that the d13Ccarb fluctuated strongly, with individual sample points deviating by more than 4.0‰ from baseline values. However, the authors also found that d18Ocarb values covaried with d13Ccarb. Consequently, the authors concluded that the apparent negative CIE was likely an artifact of diagenesis. Indeed, later analysis of a correlative basinal section from elsewhere in the Lagronegro basin (at Monte Sirino) found consistent values of d13Ccarb across the Norian–Rhaetian boundary (Tanner et al. 2007). The Carnian through Rhaetian is the portion of the Triassic System for which d13C of pedogenic carbonate is best documented, primarily because it

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was during these stages that semi-arid to arid conditions prevailed over large regions Pangaea, allowing the formation of calcic paleosols. Various data sets (Ekart et al. 1999; Ghosh et al. 2001; Tanner et al. 2001) have produced values of d13C that allow calculation of palaeo-pCO2 via the diffusionreaction model (Cerling 1991; Ekart et al. 1999); these calculations suggest pCO2 levels rose during the Middle to Late Triassic, approaching or exceeding 2000 ppm by Norian time. The data of Cleveland et al. (2008) are interpreted as indicating that pCO2 levels varied strongly, from ,100 ppm in the late Norian to .3000 ppm in the Rhaetian; these results, however, are inconsistent with other proxies or geochemical models for this interval that suggest relative stability of atmospheric composition (cf. Royer et al. 2004).

Latest Rhaetian A significant negative CIE has been identified in a number of marine sections in close proximity to the Rhaetian-Hettangian boundary. For example, the sections at St Audrie’s Bay, SW England (Hesselbo et al. 2002, 2004), Cso¨va´r, Hungary (Pa´lfy et al. 2001), and Tiefengraben, Austria (Kuerschner et al. 2007) display a negative d13Corg excursion of approximately 2.0 to 4.0‰ from a baseline that generally varies from 226 to 229‰. Typically, these excursions begin below the highest occurrence of conodonts and Triassic ammonites (e.g. choristocerids), supporting their correlatability, and also consistently below the lowest occurrence of Jurassic (psilocerid) ammonites, which demonstrates unequivocally that the base of the CIE lies below the biostratigraphic system boundary, as presently defined (the FAD of Psiloceras spelae). For example, in the Ferguson Hill section (New York Canyon) of Nevada, USA (Fig. 4), a negative d13C excursion of about 2.0‰ begins just below the highest occurrence of conodonts, Triassic ammonites (Choristoceras crickmayi and Arcestes spp.) and Triassic bivalves (Guex et al. 2004; Lucas et al. 2007). A negative CIE of similar magnitude spans the system boundary at the Kennecott Point section in the Queen Charlotte Islands, Canada; here the excursion begins immediately below the highest occurrence of Triassic ammonites and radiolarians, and continues above the lowest occurrence of Jurassic radiolarians (Ward et al. 2001, 2004; Williford et al. 2007b). Characteristically, the excursion continues upward into basal Hettangian strata and is succeeded by a strong positive excursion, as at Ferguson Hill (Guex et al. 2004; Ward et al. 2007), St. Audrie’s Bay (Hesselbo et al. 2002, 2004) and Kennecott Point (Ward et al. 2001, 2004). A similar negative excursion in the uppermost Rhaetian has been demonstrated in several

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other sections based on the d13Ccarb data (e.g. Pa´lfy et al. 2001; Galli et al. 2005, 2007). However, these data generally are of somewhat lower resolution than the d13Corg data because lithological considerations and/or diagenesis limit the sample density. Several authors have claimed a negative CIE recorded by organic carbon in terrestrial environments that is correlative with that in the marine realm (McElwain et al. 1999; Hesselbo et al. 2002). In so doing, the authors assume that the marine CIE resulted from a drastic alteration of the d13C of the global ocean/atmosphere carbon reservoir that was recorded similarly by vascular plants. Lucas & Tanner (2007), however, critiqued these data and found them unconvincing. Indeed, the isotope data from plant macrofossils for sections in Scania spanning the system boundary do not display a distinct CIE, and although the d13C data from Greenland display an apparent trend that appears to correlate with the marine data, this trend is based on very few samples and lacks the consistency that is displayed in the marine record (McElwain et al. 1999; Hesselbo et al. 2002). Moreover, the exact correlation of the system boundary between terrestrial sections and their marine counterparts has not been demonstrated unequivocally (Lucas & Tanner 2007). Additionally, it is important to note that variations in the isotopic composition of plants may result from environmental factors other than the d13C of the atmosphere (e.g. water stress). Furthermore, the isotopic composition of terrestrial organic matter varies by specific organic compound (Huang et al. 1999; Schefuß et al. 2003), and thus significant isotopic variation occurs between species (Jahren 2004), and even between locations on a single specimen (Dawson et al. 2002). These caveats apply to the analysis of organic carbon in marine sediments as well; van de Schootbrugge et al. (2008) have argued that in many instances negative CIEs identified in marine bulk organic carbon analyses reflect shifts in provenance of the organic matter (e.g. pollens v. spores). Consequently, future studies of the isotopic composition of organic matter from both terrestrial and marine realms are likely to be based increasingly on compound-specific analyses that can more reliably track variations in isotopic composition with age. The widespread nature of the negative isotope excursion draws comparisons to other major palaeontological boundaries in which negative carbon isotope excursions are prominent features, such as the end-Permian (Holser & Magaritz 1992; Magaritz et al. 1992), the end-Cretaceous (Kump 1991) and the late Palaeocene (Koch et al. 1992; Norris & Ro´´hl 1999), although these various isotopic events vary in their magnitude and duration.

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Fig. 4. Curves of d13Corg from the sections at Kennecott Point, British Columbia (adapted from Williford et al. 2007), Ferguson Hill, near New York Canyon, Nevada (from Guex et al. 2004) and St Audrie’s Bay (from Hesselbo et al. 2004) display parallel trends, including a negative excursion starting just below the Rhaetian–Hettangian boundary and a succeeding positive excursion in the Hettangian.

The significance and causes of most of these isotopic events remain incompletely understood, however, although diverse mechanisms have been proposed and various attempts have been made to explain them through mass-balance modelling (Kump & Arthur 1999; Beerling & Berner 2002; Berner 2002; Dickens 2003). Beerling & Berner (2002) and Berner (2002) have pointed out that the loss of primary productivity alone can account for no more than one-half of the observed isotopic shift in the latest Rhaetian. As has been suggested for the endPermian excursion, the rapid release of dissolved CO2 derived from organic decomposition during

ocean overturn has been proposed as a source of substantial volumes of light carbon (Wignall & Hallam 1992; Knoll et al. 1996). Such an overturn should be recorded by deposition of anoxic ocean sediments, but widespread anoxia is not recognized at the Triassic –Jurassic boundary (Pa´lfy et al. 2001). Methane release also has been proposed to explain the latest Rhaetian CEI (Wignall 2001; Berner 2002), as it has for the end-Permian (Krull & Retallack 2000). Pa´lfy et al. (2001) proposed that CAMP eruptive activity in some way triggered methane release that led to biotic extinction at the system boundary. However, mass balance

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calculation of the volume of methane required to effect a significant isotopic shift in the ocean carbon reservoir raises questions regarding the potential to store sufficient quantities of methane in the warm ocean of the Late Triassic. The modern clathrate reservoir, estimated at 3000 Gt of carbon (Buffet & Archer 2004), is less than the 5000 Gt invoked by Beerling & Berner (2002) to explain apparent warming at the system boundary. Moreover, methane storage decreases substantially with ocean warming, as much as 85% less with a 3 8C increase in deep ocean temperature (Buffet & Archer 2004). Given the high pCO2 of the Late Triassic, possibly 2000 ppm or higher, according to proxies and geochemical modelling (Ekart et al. 1999; Berner & Kothavala 2001; Tanner et al. 2001), the Earth was warmer and ice-free at this time (Sellwood & Valdes 2006). Hence, the clathrate reservoir in fact may have been substantially smaller than at present. Although McElwain et al. (1999) initially suggested a major role of CAMP degassing of isotopically light CO2 in driving the isotopic shift, subsequent authors have discounted this hypothesis on the basis that unreasonably large volumes of mantle-derived CO2 with d13C ¼ 25 to 26‰ are required to effect the observed shift. Notably, these calculations assume a deep mantle origin for CAMP magmas, which is not necessarily the case. Wignall (2001) pointed out that much lighter CO2 (d13C ¼ 220‰) may be derived from the carbon in recycled lithosphere at subduction boundaries. Current understanding of the source of the CAMP magmas does not dictate against such recycling. Consequently, degassing of isotopically light CO2 remains a viable, and potentially testable hypothesis for explaining the negative CIE in the latest Rhaetian. This hypothesis need not exclude other hypothetical mechanisms, as the environmental consequences of the CAMP eruptions, including atmospheric opacity and acid fall-out, potentially forced a significant decline in primary productivity. For example, Quan et al. (2008) identified a shift in the nitrogen isotope composition across the system boundary from a core in Germany. This shift from lower to higher d15N is accompanied by a rise in redox sensitive trace metal concentrations and coincides with a peak in abundance of trilete spores (i.e. a ‘fern spike’). The authors explain these concomitant trends as a consequence of water column denitrification and decreasing oxygenation driven by enhanced weathering following SO2 emissions of the CAMP eruptions. Together, these cumulative effects of the CAMP eruptions are likely to have produced the observed change in the isotopic composition of carbon in the ocean-atmosphere system and forced the biotic turnover that marks the system boundary. The

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subsequent positive CIE in the Hettangian (Hesselbo et al. 2002, 2004; Guex et al. 2004; Ward et al. 2007; Williford et al. 2007b) has been explained by van de Schootbrugge et al. (2008) as largely a consequence of enhanced organic carbon burial driven by higher sedimentation rates, forced by a higher pCO2 resulting from the CAMP eruptions.

Summary Among the various stable isotopes utilized in sedimentary geology, carbon isotope stratigraphy provides the most complete record of environmental change during the Triassic. A significant negative CIE that begins in the uppermost Permian and extends into the basal Triassic, has been explained as the consequence of ocean anoxia and overturn, productivity collapse, seafloor methane release, or release of volcanogenic CO2 during the Siberian Traps eruptions. A simultaneous positive sulphur isotope excursion is consistent with ocean anoxia at this time. The ensuing 4 Ma to 6 Ma of the Lower Triassic records substantial instability of the carbon cycle, with pronounced positive and negative excursions that may reflect some combination of volcanogenic CO2 induced warming that caused enhanced productivity and anoxia. Most of the succeeding Middle and Upper Triassic record is stable, lacking any significant positive or negative excursions, but rising d13C over much of this extended interval likely reflects increased burial of organic carbon in terrestrial environments. A pronounced negative CIE in the uppermost Rhaetian may have resulted from some combination of outgassing during the CAMP eruptions and primary productivity collapse. The warm seas of the Late Triassic greenhouse world were unlikely to have stored sufficient methane to cause the observed isotope excursion. I wish to express my gratitude to Spencer Lucas for providing the opportunity to contribute this chapter, and to James Ogg, Nereo Preto and Sylvain Richoz for their invaluable assistance in reviewing the manuscript.

References A TUDOREI , V. 1999. Constraints on the Upper Permian to Upper Triassic marine carbon isotope curve. Case studies from the Tethys. PhD Thesis, University of Lausanne, Switzerland. B ANNER , J. L. & H ANSON , G. N. 1990. Calculation of simultaneous isotopic and trace element variations during water-rock interaction with applications to carbonate diagenesis. Geochimica et Cosmochimica Acta, 54, 3123–3137.

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Cyclostratigraphic record of the Triassic: a critical examination LAWRENCE H. TANNER Department of Biological Sciences, Le Moyne College, Syracuse, NY 13214, USA (e-mail: [email protected]) Abstract: High frequency (fourth- and fifth-order) cyclicity is a common feature of sedimentary sequences in all depositional settings. While tectonism and autocyclic processes are clearly responsible for this cyclicity in some instances, many cases are interpreted as resulting from orbitally forced variations in solar insolation at the Milankovitch frequencies, that is, the precession and short and long eccentricity cycles at scales of tens of thousands to hundreds of thousands of years. This forcing is presumed to have controlled sedimentation through periodic changes in climate or sea-level. Examples of interpreted Milankovitch-frequency cyclicity occur throughout the Triassic record, and include much of the German Triassic, the Alpine Triassic and the Newark Supergroup of North America. The cyclostratigraphy of these sections has been used as a tool for intrabasinal and interbasinal correlation, and for chronostratigraphy. These interpretations are not always without controversy, however, as conceptual arguments and radio– isotopic age data have called some of these conclusions into question.

The stratigraphic record is inherently cyclical, but the scale of this cyclicity varies in scale from the outcrop to the continental. The largest scale, or firstorder cycles are those that are truly continental in scope, as they are thought to record changes in sealevel caused by the operation of the Wilson cycle at a scale of hundreds of millions of years (Prothero 1990; Jacquin & Graciansky 1998). Second-order cycles, corresponding to the sedimentary sequences of Sloss (1963), last from tens of millions up to one hundred million years, and are generally considered the result of long-term trends in eustasy, potentially in direct proportion to the activity of the mid-ocean ridges. The cause of the third-order sea-level cycles, of millions of years duration, is not well understood; these cycles are often apparent at the basin scale and may be related to short-term tectonic activity. Higher frequency fourth- and fifth-order cycles frequently are evident at the outcrop scale. Fourth-order cycles include the classic cyclothems of the Carboniferous (e.g. Heckel 1986) and have periods in the range of hundreds of thousands of years, while fifth-order cycles operate at frequencies of tens of thousands of years. These high-frequency (fourth- and fifth-order) cycles are often referred to as Milankovitch cycles due to the match of their interpreted periods with the calculated Milankovitch orbital frequences. It is these high-frequency cycles that are the focus of much modern cyclostratigraphy, and the record and application of these cycles in the Triassic system is the focus of this chapter.

Orbital forcing The Milankovitch hypothesis has become a conceptual paradigm in sedimentary geology because it

appears to not only explain the rhythmicity of the Pleistocene glaciations, but also provides a model of stratigraphic succession that can be recognized, seemingly ubiquitously, in the rock record. Since the 1980s, an enormous volume of literature has been generated that is dedicated to the hypothesis that the cyclicity of sedimentary sequences has been controlled by the orbital forcing of climate on scales of 104 to 105 years. The concept is certainly not a recent one. Herschel (1832) suggested that changes in the amount of solar radiation reaching the Earth, caused by the eccentricity and precession cycles, could significantly affect climate and geological processes; Croll (1864), for example, specifically hypothesized orbital control for the cyclicity of Pleistocene glaciation. G. K. Gilbert (1895) pioneered the use of presumed climatic cycles in pre-Pleistocene stratigraphy by interpreting sequences of Cretaceous limestone-shale couplets as resulting from the precession cycle, but it was Bradley’s (1929) analysis of varves in the Green River Formation that provided the first solid evidence linking orbital periods and sedimentary cycles. Milutin Milankovitch (1941) is justly famous for calculating the changes in insolation that resulted in the past from the variation of the three key orbital parameters, citing these variations as cause for the Pleistocene glaciations. Corroborating evidence for the hypothesis appeared in the stable-isotope record of deep-sea sediments (Emiliani 1955; Emiliani & Geiss 1958), but did not achieve rigor until the ages of the deep-sea sediments were constrained by magnetostratigraphy (Hays et al. 1976). Even though the evidence for regular periodicities of sedimentary cycles in the pre-Pleistocene

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 119– 137. DOI: 10.1144/SP334.6 0305-8719/10/$15.00 # The Geological Society of London 2010.

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generally lacks the strength demonstrated for the Pleistocene, convincing arguments have been made for orbitally forced climate control of many cyclic sedimentary sequences. The theoretical aspects of the Milankovitch hypothesis are well-summarized by Schwarzacher (1993, 2000) and Hinnov (2004). The various gravitational forces to which the Earth is subjected create a series of perturbations of the Earth’s orbit that affect both the Earth –Sun distance and the angle of incidence of the Sun’s rays and, consequently, the amount of solar insolation. Three primary orbital cycles are recognized: (1) The eccentricity of the Earth’s orbit varies from nearly 0 (almost circular) to 0.06 (slightly elliptical) and back with an average period of 100 ka (Berger 1984). Superimposed fluctuations on the degree of variation have been detected at intervals of 400 ka, 1300 ka, and 2 Ma (Berger et al. 1992); (2) The inclination of the Earth’s axis of rotation relative to the plane of the ecliptic, that is, obliquity, is presently about 23.58, but varies from 228 to 24.58 and back over an average period of 41 ka; and (3) The precession or wobble of the Earth’s axis of rotation, caused by the gravitational pull of the Sun and Moon on the equatorial bulge, results in a periodic change in the relationship of the seasons to Earth’s perihelion. The precession cycle is also modified by the gradual rotation of the elliptical orbit, resulting in average periods of the precession index at 19 and 23 ka (Berger 1988; Hinnov 2004). These three parameters combine to produce predictable fluctuations in the amount of solar energy reaching the atmosphere. The eccentricity cycle by itself has little effect on insolation but does control the amplitude of the precession effect. The obliquity cycle produces variations in the degree of seasonality and is most pronounced at high latitudes (Weedon 1993, 2003). The precession index, conversely, determines the direction of inclination when the Earth is at perihelion and aphelion and, consequently, controls the amount of radiation received during each season. The effect of this cycle is strongest at low latitudes and results in latitudinal shifts in the caloric equator and shifting of the boundaries between climate zones (Berger 1978). The interpretation of a sedimentary response to the orbital elements described above assumes regular changes in Earth’s climate systems in response to changes in insolation. However, the resulting changes in climate are not nearly so predictable as the causal insolation changes. For example, the influence of the eccentricity cycle interpreted from many sedimentary sequences is significantly stronger than the precession index, the opposite of what would be predicted from insolation alone (De Boer & Smith 1994). Various feedback systems, such as variations in ice volume or

atmospheric CO2, may be responsible for creating a resonating or oscillating climate system with a frequency matched to the eccentricity cycle. General-circulation-type climate models have only limited ability to predict specific climate paramaters for the distant geological past, due to uncertainties of land-mass position, elevation and atmospheric composition (Kutzbach 1994; Sellwood & Valdes 2006). Nevertheless, various models suggest pronounced climatic effects of orbital changes in insolation on global climate (Kutzbach 1994; Morrill et al. 2001). Implicit in the interpretation of orbital control of sedimentation is the assumption that patterns of sedimentation are more sensitive to small-scale changes in sea-level or changes in climate than to tectonic influences or autocyclic controls. Recognition of orbitally forced cyclicity in the stratigraphic record depends on being able to match the period of sedimentary cycles with frequencies in the Milankovitch band. Simply dividing the time represented within a stratigraphic section by the number of cycles present to yield an average time per cycle should not be considered sufficient to demonstrate Milankovitch periodicity conclusively. The standard method of analyzing the cyclicity of a stratigraphic sequence is spectral analysis, which determines the strength (relative frequency of repetition) of cycles of various periods, although thickness is an imperfect proxy for time in stratigraphic successions. Power spectra are generated by plotting the frequencies of the number of cycles per unit thickness in the strata; the most commonly recurring frequencies will form peaks corresponding to the cycle thickness (Fig. 1). Several methods exist for testing the significance of the identified peaks, for example matching the frequency of cycle repetition to a combination of sine and cosine waves by Fourier transform (see reviews of several methods in Schwarzacher 1993). Bundling patterns of the cycles, caused by the imprint of the eccentricity cycle on the precession cycle, are sometimes cited as evidence of Milankovitch control; that is, a 5:1 bundling ratio may be indicative of cycles with both precessional (with mean 20 ka) and short eccentricity (100 ka) frequencies. One method of identifying this bundling is plotting the cumulative departure of the cycle thickness from the mean against the position of the cycle in the section (Fig. 2). These plots, called Fischer plots, record changes in sedimentation rate, and represent in graphic form the change in accommodation space through time (Fischer 1964); increasing thickness of the cycles up-section, for example, recorded as a positive departure from the mean, presumably resulted from some combination of sea-level rise or increased rate of subsidence.

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Fig. 1. Power spectrum based on depth-rank analysis data from shallow marine carbonates in the Middle Triassic Latemar Massif in the Dolomites, northern Italy (adapted from Preto et al. 2001). Peak intervals were converted to time by assuming precessional forcing. The peak at 21.7 ka was produced by combining a frequency doublet that is predicted for the Triassic.

Triassic sea-level In one of the pioneering works of twentieth-century stratigraphy, Sloss (1963) described six major sedimentary sequences, or cycles, on the North American continent, defined by major episodes of transgression and bounding unconformities. These major cycles of eustasy, lasting from tens of millions to well over a hundred million years, are now commonly referred to as super-sequences because they are at least an order of magnitude

greater in scale than stratigraphic sequences (i.e. an unconformity bounded succession of conformable strata), as they were later defined by sequence stratigraphers (e.g. Vail et al. 1977). These supersequences, also called first-order sub-cycles (Jacquin & Graciansky 1998), are thought to be equivalent to the second-order sea-level cycles that are controlled by the rates of ocean crust formation (Plint et al. 1992). Within the framework of Sloss, the entirety of the Triassic falls within the upper part of the Absaroka sequence.

Fig. 2. This example of a Fischer plot is for a portion of a core of the Dachstein Limestone from the Transdanubian Range, Hungary (adapted from Balog et al. 1997). Identification of the bundling patterns of the cycles (i.e. bundling ratios) ideally permits correlation with Milankovitch frequencies at multiple orders.

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Fig. 3. Stratigraphy of the central Germanic basin, adapted from Bachmann & Kozur (2004), with general sea-level cycles of Aigner & Bachmann (1992).

In a now-classic paper, Haq et al. (1987) compiled relative sea-level data from both outcrop and the subsurface to establish a curve of global sealevel changes since the start of the Triassic. Based on these data, Haq et al. (1987, 1988) concluded that the area of continental emergence at the start of the Triassic was at the maximum level for the entire Phanerozoic. From this minimum, sea-level increased overall through most of the Triassic Period, before falling again in the latest Triassic (see also Hardenbol et al. 1998; Golonka & Ford 2000). This overall trend of Triassic eustasy is well illustrated by the German Triassic. As described by Aigner & Bachmann (1992), the entirety of the German Triassic can be described as a single transgressive–regressive super-sequence (i.e. second-order cycle), best exemplified by the classic succession of the Lower Triassic Buntsandstein, Middle Triassic Muschelkalk and Upper Triassic Keuper (Fig. 3). On this overall trend, according to Haq et al. (1987, 1988), are superimposed four super-cycles, the boundaries of which approximately coincide with the boundaries of the Scythian–Anisian, Ladinian–Carnian, Carnian – Norian and Norian –Rhaetian stages. Haq et al. (1987) included the uppermost Rhaetian with the Hettangian in the succeeding super-cycle. The

super-cycles are, in turn, divided into eleven third-order cycles – five in the Lower Triassic, two in the middle Triassic, and four in the Upper Triassic – each lasting one to several million years. Superimposed on these third-order cycles are the higher-frequency fourth and fifth-order cycles that have been interpreted from essentially every stage of the Triassic, from the system base to the Triassic –Jurassic boundary, and from every conceivable depositional environment, from deep marine to desert.

Case studies South China Many of the best-known and most-studied stratigraphic sections of the basal Triassic occur in South China. For example, the shelfal, mixed carbonate-clastic sediments of the Feixianguan Formation in Guangxi and Guizho provinces provide important sections for study of the Induan ‘extinction aftermath’ (Lehrmann et al. 2001). Li et al. (2007) and Guo et al. (2008) examined the cyclicity of strata in a continuous Upper Permian – Lower Triassic section, the Yinkeng Formation in the Pingdingshan section at Chaohu, Anhui Province.

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Guo et al. in their study recorded the lithology, thickness and magnetic susceptibility of a 44-m section of deep-marine strata, variously comprising sets of limestone–mudstone, limestone –marl– mudstone, or limestone–mudstone –shale, from slightly below the Changhsingian–Induan boundary to above the Induan–Olenekian boundary. Applying spectral analysis to the logged section and wavelet analysis to the magnetic susceptibility data yielded similar results. The authors identified 56 short cycles, which they interpreted as forced by the precession frequency at 19.5 ka (Berger et al. 1992), and 12 cycles of bundling at a ratio of 4.67:1, interpreted as the short eccentricity cycle. Guo et al. further used this cyclostratigraphic interpretation as a basis for calculating a duration of 1.1 Ma for the Induan stage, which accords well with estimates from other means (Gradstein et al. 2004). Lehrmann et al. (2001) examined the cyclicity of the Olenekian carbonates of the Great Bank of Guizhou in the Nanpanjiang basin, South China. These authors found that vertical facies changes in 164 metres of one section defined 83 cycles, or parasequences, that typically shallow upward from subtidal grainstones to intertidal ribbon rock. Changes in thickness of the parasequences throughout the section were interpreted by the authors as reflecting long-term changes in accommodation space within a third-order sequence. Radioisotopic constraints on the duration of deposition of the section suggest that the parasequences represent fifth-order cycles (0.01–0.1 Ma), and a purported 5:1 bundling ratio imply low-amplitude eustatic control of deposition at the precessional frequency (Lehrmann et al. 2001).

The German Triassic Alberti (1834) defined the Triassic for the tripartite nature of the sequence in the Germanic basin (the southern part of the Central European basin), where its thickness exceeds 3000 m. The classic German Triassic stratigraphy consists of the mainly continental Buntsandstein, the marine to marginalmarine Muschelkalk, and the largely continental Keuper (Bachman & Kozur 2004). The poor biostratigraphic record of portions of this sequence long have made precise correlation to the marine realm problematic. However, recent advances in the magnetostratigraphy of the German Triassic and in biostratigraphy based on conchostracans have led to greatly improved correlations to the Tethyan and Boreal realms (Szurlies et al. 2003; Bachman & Kozur 2004; Kozur & Bachman 2006; Szurlies 2007). Cyclicity is a prominent characteristic of much of the strata of the Germanic basin, with cycles of varying thickness interpreted as expression of

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Milankovitch frequencies from 20 ka–400 ka (Bachman & Kozur 2004; Kozur & Bachman 2006). Gaps are evident in the sedimentary section that compromise the ability to identify cycles, particularly in the Middle and Upper Triassic portions of the section. Buntsandstein. Szurlies (2007) divided the entire Buntsandstein into 60 widely correlatable cycles that he interpreted as forced by climate change at the eccentricity frequency. Furthermore, the author derived a duration of 6 Ma for deposition of the Buntsandstein from this cyclostratigraphy. This estimate accords well with the calculation of Menning et al. (2005), which combined the cyclostratigraphy with radio–isotopic constraints on stage boundaries. The stratigraphic trends of the Buntsandstein in the center of the Germanic basin are well-summarized by Bachmann & Kozur (2004) and Kozur & Bachmann (2006). In central Germany, the Lower Buntsandstein comprises roughly 300 m of red-bed siliciclastics and interbedded limestones. Szurlies et al. (2003) subdivided this succession into 20 fining-upward cycles that can be correlated across the Central European basin, and hypothesized that they might have been forced by baselevel fluctuations driven by Milankovitch climate forcing at the eccentricity frequency. Bachmann & Kozur (2004) and Kozur & Bachmann (2006) described the basinal facies of the Lower Buntsandstein, comprising the Calvo¨rde and Bernburg formations, as consisting of 22 cycles of 10 m –20 m thickness of sandstone to oolite, fining upward to shale. Bachmann & Kozur (2004) attributed the cyclicity of these formations, which are dated as upper Changhsingian to lower Smithian, to orbital forcing at the 100 ka (short eccentricity) frequency. The Middle Buntsandstein (Smithian –basal Anisian) encompasses the (in ascending order) Quickborn Sandstone and the Volpriehausen, Detfurth, Hardegsen and Solling formations, which collectively display 35 –40 cycles that are mostly fining-upward. Bachmann & Kozur (2004) interpret these cycles as resulting from the short eccentricity (100 ka) frequency, but cautioned that unconformities within the section prevent use of the cycles from accurately determining the duration of the Middle Buntsandstein deposition. The Upper Buntsandstein in the central Germanic basin consists of the 150 to 300 metre-thick Ro¨t Formation (lower Anisian), which comprises nine fining-upward cycles. Based on their ability to divide these cycles further into five subcycles, Bachmann & Kozur (2004) interpreted these as short eccentricity cycles. Orbitally forced cyclicity is interpreted also in the Buntsandstein from regions at the margins of the Germanic basin. For example, Clemmensen et al. (1994) examined the Middle Buntsandstein

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on the island of Helgoland, in the northwest Germanic Basin. These authors recognized arid-humid climatic cyclicity in the alternating eolian –lacustrine facies and interpreted thin (c. 2.5 m) cycles as forced by the precession cycle, and thicker (11.0 m) cycles as the result of climate forcing at the eccentricity frequency. Significantly, however, these interpretations lacked statistical rigor because they were based solely on the number of cycles counted and a 10 Ma estimate of the duration of the Scythian, now considered unlikely and superceded by more recent estimates (Gradstein et al. 2004). Bourquin et al. (2006) also analyzed the sequence stratigraphy of the Triassic, but in the western Germanic Basin, rather than the centre. These authors defined two major depositional cycles, a Scythian cycle encompassing the Lower to Middle Buntsandstein, unconformably separated from an Anisian –Carnian cycle comprising the Upper Buntsandstein and the entire Muschelkalk. These cycles correspond approximately to supercycles (Upper Absaroka) UAA-1 and UAA-2 of Haq et al. (1987). The overall pattern through most of the Middle Buntsandstein is one of deposition by braided fluvial systems that were laterally equivalent to lacustrine depositional systems situated toward the basin center. Passing into the Upper Buntsandstein, the trend is toward development of lower gradient floodplain and more abundant lacustrine facies, suggesting a closer proximity to the Tethys Seaway late in the Scythian. This is consistent with the interpretation of Haq et al. (1987, 1988) of a gradual transgressive trend through most of the Upper Absaroka sequence. Cycles in the Middle and Upper Buntsandstein are attributed by Bourquin et al. (2006) to fluctuations in climate and/or sediment supply. Muschelkalk. The Middle Triassic Muschelkalk comprises the mainly fully marine Lower Muschelkalk, hypersaline facies of the Middle Muschelkalk and the predominantly marine Upper Muschelkalk. In the central Germanic basin, the Lower Muschelkalk consists mainly of limestones and marls of the Jena Formation of Anisian age. Cyclostratigraphic patterns are interpreted from the Muschelkalk sediments as they are in the Buntsandstein (Kozur & Bachmann 2003, 2006; Bachmann & Kozur 2004). The Jena Formation, for example, includes 21 cycles that Bachmann & Kozur (2004) interpreted as short eccentricity cycles. In the central Germanic basin, the Middle Muschelkalk comprises the Karlstadt, Heilbronn and Diemel formations (Anisian in age), which together contain nine cycles. The Upper Muschelkalk consists of the Trochitenkalk and Meissner formations (upper Anisian through lower Ladinian) and together include 40 short eccentricity cycles.

Vecsei & Duringer (2003) interpreted the Middle to Upper Muschelkalk as part of a long-term, third-order transgressive cycle. The authors recognized fourth-order deepening-upward cycles in the deeper shelf environments, caused by highfrequency sea-level changes superimposed on the longer term eustatic trend. They also noted, however, that contemporaneous cycles in coastal bar and lagoonal environments appear to be shallowingupward, and speculated that the sedimentary systems in these differing environments apparently responded in different ways to sea-level forcing. Menning et al. (2005) calculated a 6.4 Ma duration for total Muschelkalk deposition. Keuper. The high-frequency cyclicity documented in the Muschelkalk continued during the deposition of the largely continental sediments of the Middle to Upper Triassic Keuper (Kozur & Bachmann 2003, 2006; Vecsei & Duringer 2003; Bachmann & Kozur 2004). The Lower Keuper is represented by the Erfurt Formation, which displays eight cycles (Bachmann & Kozur 2004). Fourth and fifth-order cyclicity is evident also in the Ladinian through Norian (through Sevatian) Middle Keuper, which in the central Germanic basin includes the Grabfeld, Stuttgart, Weser formations and Lo¨wenstein (equivalent to Arnstadt) formations (e.g. Bachmann & Kozur 2004). Kozur & Bachmann (2006) identified nine cycles in the (Ladinian) Grabfelt Formation. Reduced rates of sedimentation, condensed sections and substantial stratigraphic gaps prevent clear identification of unambiguous short eccentricity cycles in the remainder of the Middle Keuper, or the Exter Formation (Rhaetian) of the Upper Keuper, although long eccentricity cycles (i.e. 400 ka) have been recognized in the Arnstadt Formation (Bachman & Kozur 2004). Reinhardt & Ricken (2000) studied the periodicity of the decimeter-scale alternating mudstone– dolomite beds of the Steinmergel Middle Keuper in southern Germany, which they interpreted as reflecting fluctuating hydrological conditions on an extensive playa system, driven by climate change. Using spectral analysis of core and outcrop sections, the authors identified cycle intervals that they correlated to precession (19.8 ka), short eccentricity (109 ka), long eccentricity (413 ka) and very long eccentricity (2 Ma) frequencies of Berger & Loutre (1989, 1994). Combining available radio– isotopic dates with the cyclostratigraphy, Menning et al. (2005) calculated a duration for total Keuper deposition of c. 40 Ma.

The Alpine Triassic Fischer & Bottjer (1991) described the Alpine Triassic as ‘the classic ground for the recognition

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of cyclic carbonate platform emergence, presumably a measure of eustatic oscillations’. Triassic sedimentation in the Alpine region was controlled by the interplay of tectonics and eustasy. In the Northern Calcareous Alps, for instance, the Permian through the Early Triassic was characterized mainly by deposition of siliciclastics during a transgression on a passive continental margin, likely formed by Tethyan rifting (Mandl 2000). Carbonate shelf deposition was established on this margin by Middle Triassic, but tectonism, probably during the middle Anisian, caused block faulting that segmented the shelf and created a series of isolated platforms separated by interplatform shelf and basinal environments (Balog et al. 1997; Mandl 2000). Progradation of the platforms continued into the early Carnian, followed by an interval of lowstand and emergence. Rising sea-level during the late Carnian and increased carbonate production filled much of the interbasin palaeotopography. During the earliest Norian, a rapid pulse of sea-level rise accompanied accentuated growth of reefs that rimmed these broad platforms. Latemar. One of the Alpine areas most studied is the Latemar Massif, a carbonate platform of Middle Triassic (upper Anisian –Ladinian) age in the Dolomites of northern Italy. The cyclic facies of the Ladinian-age Latemar Limestone are characterized by sub-metre to metre-scale couplets of subtidal limestone with a thin dolomite caliche cap (Goldhammer 1987; Goldhammer et al. 1987; Hinnov & Goldhammer 1991). The orderly and repeated arrangement of these facies clearly suggests episodes of submergence, during which limestone was deposited subtidally, alternating with episodes of subaerial exposure of the carbonate sediments that led to vadose diagenesis. The cycles appear to be further ordered in upward-thinning bundles of five cycles each. Hinnov & Goldhammer (1991) applied spectral analysis to the stratigraphic section and applied an estimate for the duration of the Ladinian stage to obtain an average value of 21 ka for the cyclic couplets and 96 ka and 110 ka for the modulating eccentricity frequency. Brack et al. (1996) and Mundil et al. (1996) cast serious doubt on the conclusion that Latemar cyclicity was forced only by Milankovitch frequency orbital cycles. The authors of these studies performed single crystal U –Pb dating on volcaniclastic beds in the Buchenstein Beds, the basinal equivalent of the Latemar platform facies that was deposited in a sediment-starved intra-platform basin. Ages from the beds bracketing the equivalent cyclical facies of the Latemar platform indicated that the time interval represented was too short to accommodate deposition of the 598 precession (c. 20 ka) cycles counted by Goldhammer (1987), even with maximum error

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estimates for the dates. Consequently, Brack et al. (1996) and Mundil et al. (1996) concluded that the metre-scale cyclicity of the Latemar succession reflects processes operating on a millennial scale (,8 ka), a higher frequency than predicted by the Milankovitch hypothesis. Egenhoff et al. (1999) reexamined the facies architecture of the Latemar Massif with an eye toward resolving some of these issues. In their view, the overall facies architecture of the build-up (i.e. alternating cyclical and tepee facies) reflects the effects of third-order sea-level cyclicity on the depth of water over the build-up. The authors found that the individual shallowing-upward cycles, defined by marine flooding surfaces at the base and top, are traceable across the build-up for a distance of several kilometers, although thickness and facies vary considerably within individual cycles, due to varying positions on the platform (i.e. intertidal versus lagoonal). They surmised that autocylic processes (e.g. lateral facies migration) are inadequate to explain exposure surfaces that are correlative across the entire platform, and concluded that some order of cyclic sea-level fluctuation likely influenced deposition of the Latemar cycles. However, they left unresolved the issue of the timing of the cycles. An additional perspective was provided by Preto et al. (2001). These authors logged a 160-m section of lagoonal carbonates on the Latemar platform in the so-called Upper Cyclic Facies at cm-scale and applied a depth-rank analysis to the resultant data. Subsequent spectral analysis identified peaks that the authors associated with the precession, obliquity, short eccentricity and long eccentricity cycles. Notably, the authors rejected the issues raised by radio– isotopic dating of the Latemar and equivalent Buchenstein Beds, and stated that ‘The Latemar signature thus constitutes the oldest pristine Milankovitch signature yet observed in the geologic record’. The controversy over the frequency of the Latemar cycles continued as Mundil et al. (2003) produced new U –Pb zircon dates from ash beds in the platform interior to better constrain the interval of deposition. These new data indicated that the average length of time represented by individual shallowing-upward cycles must be less than previously calculated, closer to 4 ka. Zu¨hlke et al. (2003) generated a new set of cyclostratigraphic measurements from the Latemar section, which they analyzed through spectral analysis. Using the age constraints of Mundil et al. (2003), the authors found that although individual cycles occur at a subMilankovitch frequency, stacking patterns (1:4–5) produced higher-order cycles that appear to match the precessional frequency, not the short eccentricity frequency, as previously interpreted (Goldhammer et al. 1987). Moreover, stacking

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Fig. 4. Cyclicity of the Upper Triassic Dachstein Limestone. (a) Classic section of metre-scale cyclically bedded shallow marine carbonates at Loferer Steinberge, southwest of Salzburg, Austria. (b) Detail of Lofer cycles in the Dachstein at a quarry at Tata, Hungary. The top of one cycle occurs at the top of the microbial-laminite below the scale. The laminite is overlain by subtidal carbonate containing thick-shelled megalodontid bivalves (to the left and right of scale).

patterns of 1:9.9 and 1:24.0 identified by power spectra appear to match the obliquity and short eccentricity frequencies. Kent et al. (2004) contributed to the Latemar discussion with magnetostratigraphic data from the Latemar platform succession that they correlated to the Buchenstein Beds, apparently validating the radio –isotopic age correlations between the platform and the very condensed basinal successions. Hinnov (2006) critiqued the work of Kent et al. (2004), particularly in regard to the validity of the correlation between the Latemar platform beds and the Buchenstein Beds, but she offered no effective counter-argument to the constraints provided by the radio –isotopic ages, as pointed out by Kent et al. (2006). Meyers (2008) attempted to provide some resolution to the issue through a new analysis of the Latemar data of Preto et al. (2001) by utilizing a multitapermethod spectral analysis. The fit of the cycles to orbital frequencies was quantified with an average spectral misfit algorithm that was used to test the probability of various sedimentation rates. Meyers’ test rejected the sedimentation rate applied previously by Preto et al. (2001), but obtained a significant fit with a much higher rate that is compatible with the age constraints of Mundil et al. (2003) and Kent et al. (2004). Nevertheless, the spectral analysis of Meyers was able to identify cycles at the precessional, obliquity and short eccentricity frequencies. Lofer cycles. The classic Upper Triassic carbonate platform succession of the Tethyan margin is presented by the Dachstein Formation, which is widely exposed in the Northern, Southern and Eastern Alps of Austria, Italy, Hungary and Slovakia (Fig. 4a). The stratigraphy of the Alpine Triassic is complex.

In some parts of the Northern Calcareous Alps, for example, the Dachstein encompasses the entirety of the Norian and Rhaetian (Fig. 5), whereas in other areas it overlies the Hauptdolomit and is overlain by the Ko¨ssen Formation, or is laterally equivalent to these units, as in the Dolomites (Flu¨gel 1981). The rhythmicity of the bedding of the Dachstein shallow platform carbonates has been long recognized and interpreted as evidence of some ordered cyclicity of depositional processes (Sander 1936), most likely related to sea-level change (Schwarzacher 1948, 1954). Fischer (1964, 1975) described the classic Lofer cycles (named for the Loferer Alps in the Salzburg region of Austria) as deepeningupward transgressive or transgressive– regressive couplets, comprising a basal paleosol, often clayey, overlain by intertidal microbial laminites, and subtidal skeletal wackestone/packstone that may or may not be overlain by another intertidal laminite, all capped by a succeeding palaosol. Fischer (1964, 1975) interpreted these cycles as driven

Fig. 5. Generalized stratigraphy of the Upper Triassic in the Northern Calcareous Alps, from north to south across the Bavarian and Tyrolian nappes (adapted from Mandl 2000).

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by Milankovitch-frequency eustasy. Subsequent researchers have found that the classic cycles of the Dachstein in fact differ considerably from the description by Fischer. Goldhammer et al. (1990), Haas (1994) and Satterley (1996) reinterpreted the classic Northern Alpine cycles as dominantly regressive, shallowing-upward sequences in which a basal palaeosol is overlain directly by subtidal facies, which is in turn succeeded by the intertidal facies (Fig. 4b). Goldhammer et al. (1990), in particular demonstrated that the cyclicity failed to display the orderly stacking patterns predicted by Fischer. Balog et al. (1997), working in the Transdanubian Range of western Hungary, similarly found that the cycles there displayed quite a varied lithostratigraphy, including transgressive, regressive and transgressive-regressive couplets (Fig. 6), and considered the cyclicity as a genuine artifact of highfrequency (fifth-order) eustasy operating in combination with local changes in accommodation space. Up to 3 km of platform carbonates are wellexposed in the Transdanubian Range, where inner platform facies comprise cyclically bedded lagoonal-peritidal carbonates. In the Dachstein Limestone of Hungary, modulating cycles, or stacking patterns of individual Lofer cycles are notably absent (Goldhammer et al. 1990). Haas (1994) examined in detail the extent of lithostratigraphic variations within the cycles throughout the succession and concluded that individual Lofer cycles likely do represent metre-scale sea-level variations at the precessional frequency, but also concluded that no higher-order composite cyclicity is present in the section, likely as a consequence of superimposed variation in subsidence rate. Similarly, Balog et al. (1997, 1999) examined the metre-scale cyclicity of the Upper Triassic (Norian) Main Dolomite, which underlies the Dachstein in the Transdanubian Range (Hungary). Here the cycles are defined, as they are in the Latemar, by subaerial exposure surfaces, which are characterized by dolomitization and caliche (Fig. 6). They interpreted the cyclicity as most closely approximating a precessional frequency forced by sea-level change, but concluded that modulating cycles were only poorly developed (Fig. 2). Reijmer & Everaars (1991) hypothesized that facies deposited as the basinal equivalent of the Dachstein platform should display a similar periodicity as the classic Lofer cycles because basinal sediments would have been derived from the platform at rates corresponding to the platform sediment production. The authors tested a sequence of calciturbidites from the Pedata/Po¨tschen Schichten in the eastern Alps of Austria and found multiple spectral frequencies, some of which seemingly matched Milankovitch frequencies, although others did not.

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Fig. 6. Measured section of the Dachstein Limestone platform carbonates from a core in the Transdanubian Range (Hungary) illustrates a combination of regressive, transgressive, and transgressive– regressive cycles (adapted from Balog et al. 1997).

Satterley & Brandner (1995) and Satterley (1996) presented a rigorous analysis of the classic section at Steinernes Meer, Austria and found that the cycle thickness-frequency distribution displays an exponential pattern which is strongly suggestive of depositional control by non-periodic processes. Satterley (1996) noted in particular that the Lofer cycles on the Steinernes Meer platform displayed limited lateral continuity, low stratigraphic completeness, and a lack of hierarchical bundling. Satterley & Brandner (1995) and Satterley (1996) thus concluded on this basis that autocyclical processes, such as lateral facies migration within a tidal flat island system, and variations in the rate of

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subsidence, were the over-riding controls on stratigraphic patterns of sediment accumulation. Nevertheless, Cozzi et al. (2005) were able to measure 112 cycles in a 271-m thick section of the Dachstein in the Julian Alps, northeastern Italy. Spectral analysis again suggested that these shallowing-upward carbonate cycles were deposited with a periodicity matching precessional frequencies at 19 ka and 22–24 ka, with a potential for 5:1 bundling at 96– 128 ka. Similarly, Schwarzacher (2006) discounted alternative (e.g. autocyclic) mechanisms for genesis of the cycles in his spectral analysis and interpretation of several Dachstein Limestone sections in the Northern Calcareous Alps of Austria. He concluded that cyclicity can be attributed directly to orbital forcing at the precessional and eccentricity frequencies.

North America Newark Supergroup. Van Houten (1962, 1964) first recognized and described an apparent periodicity in the cycles of the sedimentary sequence of the Upper Triassic (Carnian) Lockatong Formation of the Newark basin (Fig. 7a) and proposed that sandstone–mudstone–shale sequences represented transgressive– regressive lacustrine cycles, that is, lake expansion and contraction, in an alluvial– lacustrine basin. Assuming that carbonate-clastic couplets in the dark mudstones recorded annual deposition of varves, Van Houten calculated an average sedimentation rate for the cycles and determined that individual cycles represented an average depositional interval of approximately 20 ka years. Van Houten hypothesized further that these cycles

Fig. 7. Cyclic strata of the Newark Supergroup. (a) Cyclically bedded lacustrine strata from the (Carnian) Lockatong Formation, Newark basin. Section photographed is below the Palisades Sill in Palisades State Park, New Jersey. (b) Strata in the upper Passaic Formation (Norian– Rhaetian) are dominated by fluvial sandstones in this quarry near Clifton, New Jersey. Cliff in the foreground is 10 m high. (c) Typical Van Houten cycle in the (Hettangian) East Berlin Formation consists of pale sandstones surrounding dark, organic-rich mudstone, overlain by red mudstone and sandstone. The light-coloured centre of the dark mudstone is a finely laminated carbonate bed. The section is located near Meriden, Connecticut, in the Hartford basin. The staff at lower right is 1.5 m. (d) Section of the (Norian–Hettangian) Blomidon Formation on the north shore of the Minas Basin, Nova Scotia (Fundy basin), comprises metre-scale alternating beds of mudstone and fine-grained sandstone. Overlying North Mountain Basalt is visible at cliff top at far left. Staff at lower right is 1.5 m.

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were controlled by the precessional signal of the Milankovitch orbital frequencies, although he recognized the imprecision and assumptions inherent in the methodology. Van Houten also recognized groupings of these cycles into thicker compound cycles of five and twenty individual cycles, but he did not speculate as to their origin. The hypothesis that Milankovitch-frequency cyclicity is recorded by the lacustrine cycles within the Newark basin strata was advanced further by Olsen (1984, 1986; Olsen et al. 1989; Olsen & Kent 1996). Olsen, who provided the term ‘Van Houten cycles’ for the ostensibly precession-forced transgressive–regressive cycles in the Newark Group formations, divided the cycles into seven lithologies and assigned a depth ranking to each lithology; that is, the shallowest-water sediments, such as red mudstones bearing vertebrate tracks and desiccation features, were assigned the rank ‘0’ and the deepest-water lithologies, the organicrich laminated mudstones, were assigned rank ‘6’. The recurrence interval of these lithofacies was then examined by Fourier analysis to generate power spectra of the sedimentary cycles in which peak recurrence intervals were statistically tested. The duration of the cycles was calculated by calibrating the cycle thicknesses to the varve counts employed by Van Houten. This permitted assignment of a period of 18 ka to 25 ka to the basic Van Houten precession cycle (Olsen et al. 1989). The power spectra also identified the compound cycles originally identified by Van Houten, and allowed assignment of these to the frequencies of the eccentricity cycles, at 95 ka and 125 ka, the long eccentricity cycle at 400 ka, and the very long cycle of c. 2000 ka. A peak was also identified at 41 ka, which matches the frequency of the obliquity cycle, but this was not statistically significant. Within the limits of resolution of the available data, these cycle periodicities were broadly supported by the radiometric data available at the time. The original work of Van Houten (1962, 1964) was limited to the Lockatong Formation. Olsen extended this downward to include the uppermost part of the dominantly fluvial (Carnian) Stockton Formation. Despite the upward-decreasing prominence of lacustrine facies in the Passaic Formation (presumably Norian to early Hettangian; see below), Olsen extended the interpreted cyclostratigraphy upward to include all of this formation (Fig. 7b), as well as the Jurassic-age Feltville, Towaco and Boonton formations. Olsen & Kent (1996) and Olsen et al. (1996a) continued the investigation of Newark basin cyclostratigraphy by analyzing seven laterally offset, stratigraphically overlapping drill cores obtained from the Newark Basin Coring Project (NBCP). The nearly 7 km of NBCP core provided ostensibly complete sections

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of the Lockatong and Passaic formations, to which the authors applied the depth-rank analysis technique. The authors analyzed the data thus generated by spectral analysis, including both fast Fourier transform (to define peak cycle thicknesses), and space-frequency analysis (to measure shifts in cycle frequency as a function of stratigraphic position). They found that the individual Van Houten cycles at 4–7 m thickness are the most obvious component of cyclicity at the outcrop scale, but statistically they are more weakly expressed than the thicker modulated cycles. In the cored strata, the authors (Olsen & Kent 1996; Olsen et al. 1996a) defined: (1) short modulating cycles, comprising five Van Houten cycles (c. 100 ka); (2) intermediate modulating cycles, which they termed McLaughlin cycles, comprising four short modulating cycles (c. 400 ka long); and (3) long modulating cycles of four to five McLaughlin cycles (c. 2 Ma long). Within the Upper Triassic section of the Newark basin, the authors identified 60 McLaughlin cycles in the upper Stockton through Passaic formations that should therefore represent 24 Ma of sedimentation (Kent & Olsen 1999); for example, the correlation of Kent & Olsen (1999) indicates that the Rhaetian of the Newark basin comprises 15 McLaughlin cycles, encompassing approximately 6 Ma. If the interpretations of the cycle durations are valid, the cyclostratigraphy of the Newark Basin (and by correlation much of the Newark Supergroup) is, in effect, a determinative chronostratigraphy. Given the radiometric dates on the Newark volcanics, which average c. 201 Ma, as a reference datum, cycle counting should yield an absolute age for the strata in these basins. Furthermore, magnetostratigraphic correlation of strata in other basins could allow chronostratigraphic correlation to other sections, for example the Tethyan marine (e.g. Channell et al. 2003; Gallet et al. 2003) or the Somerset Coast of England (see below). In theory, correlation of a chronostratigraphically calibrated section to a biostratigraphically calibrated section then would allow determination of the absolute ages of the stage boundaries. Olsen et al. (1989; see also Olsen 2003; Whiteside et al. 2007) applied this methodology to the other basins of the Newark Supergroup (Fig. 7c), from North Carolina (the Danville –Dan River basin) to Nova Scotia (the Fundy basin). In the Fundy basin, for example, Olsen et al. (1989) interpreted the cycles of sandstone with a ‘sand-patch fabric’ (sensu Smoot & Olsen 1988) and mudstone in the Blomidon Formation (ostensibly Norian– earliest Hettangian) as sediments deposited in a lacustrine –playa system subject to orbitally forced climate fluctuations with predominantly c. 100 ka (eccentricity) cyclicity (Fig. 7d). Mertz & Hubert

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(1990), however, interpreted the cyclicity of the Blomidon Formation as a result of autocylic processes on alluvial fans that bordered the playa system. Kent & Olsen (2000) subsequently recalculated the predominant cycle frequency and concluded that it was obliquity related (e.g. c. 41 ka) on the basis of their palaeomagnetic correlation of the Blomidon Formation to the Newark basin. Olsen et al. (2002; see also Whiteside et al. 2007) used the ordering of the longer period cycles as a basis for intra-basinal temporal correlations, with the volcanics in the respective Newark Supergroup basins serving as a datum. Notably, however, Marzoli et al. (2004) presented data on the age of the basalts in Morocco that call into question the assumption of synchroneity of the CAMP eruptions that underlies the intra-basinal correlations of Olsen et al. (2002). In some basins (e.g. Hartford and Deerfield), the correlations of Olsen et al. (2002) apply mainly to the Jurassic sections because lacustrine facies are rare to absent in much of the Triassic section. Additionally, the authors (Olsen et al. 1996b; Olsen 2003; Whiteside et al. 2007) used the timing of cycles bracketing the volcanics to establish a duration for the extrusive episode in the earliest Jurassic; that is, they concluded that extrusion of the Orange Mountain, Feltville and Preakness basalts in the Newark basin occurred in ,600,000 years. These methods of analysis of sedimentary sections have been applied to other continental Triassic strata in other basins and on other continents that can be correlated to the Newark. Hofmann et al. (2000), for example, described facies and metre-scale cyclicity similar to those of the Blomidon Formation in the apparently contemporaneous Bigoudine Formation of the Argana basin, Morocco. In the Upper Triassic strata of the Branscombe Mudstone Formation (of the Mercia Mudstone group) at St. Audrie’s Bay (England), Kemp & Coe (2007) used image analysis to record changes in colour of the strata, and identified spectral cycles at 26 cm and 116 cm. The authors concluded that the thicker (116 cm) intervals represent sedimentary cycles forced by the c. 100 ka eccentricity frequency, and that the thinner (26 cm) represent the precession frequency at c. 22 ka. Furthermore, the authors established a magnetostratigraphic correlation of this section to the Newark basin section that provides a link to the Newark basin chronostratigraphy. Notably, the age and completeness of the stratigraphic section in the Newark basin is not accepted universally. Kozur & Weems (2005, 2007) examined the conchostracan biostratigraphy of the Newark Supergroup. By correlation with the Germanic Triassic, the authors concluded that the section is neither complete nor appropriately

dated. A substantial regional unconformity is interpreted near the base of the section in the lower Carnian (between the Cordevolian and upper Julian). Of greater significance in the attempt to correlate the continental record to the marine is the lack of the uppermost Norian and most of the Rhaetian in the Newark basin. Kozur & Weems (2005, 2007) found conchostracans of Sevatian (late Norian) age in strata previously dated as Rhaetian (the upper Catharpin Creek and Passaic formations) in the Newark and Culpeper basins, and placed the Triassic– Jurassic boundary within the volcanic sequence, above the Preakness Basalt in the Newark Basin, and above the Sander Basalt in the Culpeper Basin. Colorado Plateau. The classic nonmarine Upper Triassic sequence of North America is presented by the Chinle Group, ranging in age from Late Carnian to possibly Rhaetian. Strata of the Chinle Group, which are exposed across much of the Colorado Plateau, were deposited in a continental retroarc basin on the western edge of the North American craton. Lucas et al. (1997) established correlations between the Upper Triassic marine strata of Nevada and the nonmarine Chinle Group strata of the Four Corners region on the Colorado Plateau (southwestern USA). The authors used these correlations to define unconformity bounded depositional sequences (or third-order cycles) and inferred a eustatic control on their deposition, although Marzolf (1994) previously had argued for a tectonic control. The influence of sea-level fluctuation on alluvial sedimentation is now well accepted (Posamentier & Allen 1993; Schumm 1993; Shanley & McCabe 1994; Atchley et al. 2004), but generally some proximity to the coastline is thought to be required for eustasy to be an effective control on alluvial base level. Studies of the modern Mississippi River system suggest that this influence extends inland no more than a few hundred kilometers (Aslan & Autin 1999; Blum & To¨rnqvist 2000). Given the likely position of the coastline in central-to-western Nevada during Chinle deposition, Cleveland et al. (2007) discounted a eustatic control on Chinle deposition. While accepting the essential sequence stratigraphic architecture of the Chinle Group of Lucas et al. (1997), these authors cited tectonics as the more likely control on the depositional sequences. Tanner (2000) studied the Upper Triassic (Norian) Owl Rock Formation of the Chinle Group on the Colorado Plateau and noted a rhythmicity within the alternating alluvial clastic –palustrine carbonate sequence (Fig. 8). This author speculated that cyclical climate change with the Milankovitch precessional frequency could provide an explanation for the observed rhythmicity through

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Fig. 8. Type section of the Upper Triassic (Norian) Owl Rock Formation exhibits rhythmic interbedding of alluvial clastics and lacustrine–palustrine– pedogenic carbonates. Orbital forcing of the cyclicity of the sedimentation in this formation, through climatic control of base level, is hypothesized, but remains untested.

strengthening and weakening of a monsoonal climate system (which would drive base-level changes), as has been modeled elsewhere for the Late Triassic (Parrish 1993; Kutzbach 1994; Lutz & Ricken 2000). Tanner pointed out, however, that age data are lacking to test the hypothesis quantitatively.

Discussion Preservation of a climate signal implies that climate is a controlling influence on the sedimentary record of every environment in which it is interpreted as present; that is, climate has a recognizable imprint over other factors, including tectonics, long-term eustasy, event stratigraphy, and autocyclic processes. Some depositional settings, such as lakes and evaporite basins, are unquestionably more climate sensitive than others, and thus provide a more faithful archive of climate change. The changing strength of the monsoons over the Pangaean continent, controlled by the precession cycle, has been cited as a dominant control on continental sedimentation by some authors (Olsen 1986; Perlmutter & Matthews 1989; Parrish 1993). In environments in which autocyclic processes operate (e.g. alluvial systems), interpretations of allocyclic controls are likely to be questioned. As Schwarzacher (2000) points out, however, the distinction between allocyclic and autocyclic processes in such settings is not always so clear; climate may exercise an extrinsic control over some autocyclic processes (such as avulsion rate) in fluvial systems, thereby clouding the distinction. Milankovitch-frequency glacio-eustasy is frequently and easily invoked as a dominant factor in

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marine and marginal marine environments for those times in geological history for which there is clear evidence of continental-scale ice sheets. But other mechanisms must be invoked for generating low-amplitude (metre-scale) sea-level changes during non-glacial intervals, such as the early Mesozoic (Frakes et al. 1992). Water flux between mountain glaciers and the oceans with Milankovitch frequencies is a distinct possibility, but would not produce the several-metre sea-level changes required to explain the sedimentary cycles; for example, melting of the entire modern Greenland ice sheet would produce approximately 7 metres of sea-level rise, but there is no evidence for ice mass of this magnitude during the Triassic. The relatively small changes in insolation associated with the orbital cycles also seem inadequate to produce thermal expansion of the ocean of this magnitude. Changing groundwater storage caused by precession-forced climate changes has been hypothesized to cause metre-scale sea-level fluctuations. Jacobs & Sahagian (1993) reasoned that large-scale fluctuations in the water table in northern Pangaea, as indicated by the changing levels of the Newark Supergroup lakes, forced metre-scale fluctuations in sea-level. Indeed, modelling of the Pangaean climate (Kutzbach & Gallimore 1989; Parrish 1993; Kutzbach 1994) suggested that strengthening of monsoonal flow in the hemisphere in which the summer occurs during the perihelion position of Earth would possibly result in a 25% increase in precipitation in tropical and subtropical regions compared to the mean. The suggestion by Jacobs & Sahagian ignores the symmetry of the climate cycle, however. The opposite hemisphere, in which summer occurs at the aphelion position, would see a similar-scale reduction in precipitation. With nearly equal land mass in each hemisphere, the surplus storage of groundwater in one hemisphere logically should be cancelled by a deficit in the other, thus affecting no net change in ocean-water volume. An alternative interpretation of the effects of orbital forcing is given by Mo¨rner (1976, 1984, 1994), who attributed changes in oceanic circulation and sea-level to cyclic adjustments in the differential rotation of the layered Earth system (hydrosphere, lithosphere, asthenosphere, mantle, core) and changes in the gravity potential of the Earth’s surface. This hypothesis remains untested, however, and the issue of the driving mechanism for highfrequency, low-amplitude sea-level change remains largely unresolved.

Carbonate platforms The interpretations of orbital forcing of cyclic sedimentary sequences found on carbonate platforms are not without controversy (for examples of

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arguments and counterarguments, see Kozar et al. 1990; Koerschner & Read 1990; Hardie et al. 1991; Read et al. 1991). A number of researchers have advocated that the ‘Ginsburg model’ (Ginsburg 1971, 1982; Wilkinson 1982) adequately explains the formation of shallowing-upward sequences in carbonates through autocyclic processes. In its essence, this model suggests that carbonate sedimentation on a constantly subsiding shelf will build a shallowing-upward sequence until water depth is too shallow to allow continued carbonate sediment production. A sedimentary hiatus takes place while subsidence continues until a critical threshold depth is reached that allows sedimentation to resume. Hence, repeated shallowing-upward carbonate cycles can occur in the absence of variations in sea-level. Indeed, computer simulations of cyclic carbonate sediments can produce simulated stratigraphic sections similar to those observed in the field with or without input of sea-level changes with Milankovitch frequencies (Goldhammer et al. 1990; Hardie et al. 1991). Algeo & Wilkinson (1988) further argued that as cycle period is dependent on cycle thickness, all cycles between 1– 20 m thick will yield cycle periods in the Milankovitch range (20 ka –400 ka), regardless of origin. Similarly, Drummond & Wilkinson (1993a) contended that the exponential frequency distribution of peritidal carbonate cycles is evidence against periodic accumulation of sediment. Drummond & Wilkinson (1993b) further argued that the assumption that each rise in sea-level produces only one cycle is not necessarily valid. Schwarzacher (2000), however, has countered that purely autocyclic models for peritidal carbonate sedimentation fail to explain the repetition of specific cycle thicknesses or the characteristic bundling patterns so frequently observed. Furthermore, as noted by Schwarzacher and others (e.g. Goldhammer et al. 1987; Egenhoff et al. 1999) the periodic occurrence of subaerial exposure surfaces in these shallow-marine carbonates, in the classic Latemar cycles, for example, is most readily explained by oscillating sea-level.

allow for other mechanisms, for example, that there was a tectonic component to base-level change. De Wet et al. (1998), for instance, noted the importance of tectonics in controlling the development of lacustrine systems in their study of lacustrine carbonates in the Gettysburg basin. Additionally, the classic calculations of cycle duration from varve counts relied on the assumptions that the carbonate–clastic couplet in the dark mudstones did in fact result from seasonal differences in sedimentation during a single year, and that the sedimentation rate estimated from the thickness of these varves can be applied to an entire cycle, which comprises a variety of lithologies, deposited by processes that vary from traction flow to suspension settling. This supposition could be reasonable only if the rate of sedimentation is considered as a function primarily of accommodation space. In such a case, a uniform, time-averaged rate of sediment accumulation may be applicable. Notably, the Newark cyclostratigraphy also presumes completeness of the stratigraphic record; that is, no significant unconformities or hiatuses occurred throughout the history of sedimentation in the basin over a span of millions of years. This assumption is contradicted by the conchostracan biostratigraphy of Kozur & Weems (2005, 2007). The lakes of the East African Rift basins often have been invoked as a modern analog for the Newark basins (e.g. Olsen 1986), given their subtropical setting in elongated half-grabens. These rift lakes differ considerably from their ancient counterparts in numerous ways, however. Water levels in the modern lakes fluctuate seasonally on a scale of metres. Lake levels have fluctuated by as much as 100 m in historic times, and by hundreds of metres since the late Pleistocene, resulting in major unconformities in the stratigraphy (Scholz et al. 1998). Thus, the analogy with the modern African rift lakes suggests that there is a strong potential for major unconformities in the Newark Supergroup cyclostratigraphy, as recognized by Kozur & Weems (2005, 2007).

Newark lakes

Conclusions

Although the operation of orbitally forced climate change with Milankovitch frequencies during the Triassic now seems well-founded, it is noteworthy that several assumptions are implicit in the development of the determinative cyclostratigraphy of the Newark basins. First and foremost amongst these is that the lithofacies changes within the cycles, that is, from red mudstone to grey sandstone to black mudstone, reflect profound changes in water depth within the Newark lakes that were driven entirely by climate change. The model does not

Studies conducted since the mid-1980s have suggested that high-frequency (fourth- and fifthorder) cycles attributable to orbital forcing are ubiquitous in the stratigraphic record of the Triassic. But identification of these cycles by the methods of spectral analysis long has relied on the assumption that the stratigraphic records under examination were essentially complete; this assumption requires testing on a case-by-case basis. Calculations of the cycle periods often has relied on bundling patterns of the cycles and the assumption that these patterns

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resulted from the predictable modulation of the basic cycle, that is, the eccentricity cycle superimposed on the precession cycle. Later application of radio-isotopic dating to some of these sections, however, has demonstrated the weakness of these assumptions. Consequently, the application of cyclostratigraphy to chronostratigraphy should be considered robust only where supporting independent age data are available. Evaluation of the accuracy of the chronostratigraphy derived from the Newark Supergroup cyclostratigraphy, for example, requires establishment of an absolute age that brackets the cyclostratigraphic section, in addition to that of the Newark volcanics. Potentially, a magnetostratigraphic correlation to another section containing volcanic materials will allow a future test of this chronostratigraphy against the more recent biostratigraphy. Until then, age correlations so derived should be regarded as tentative. Finally, many of the arguments on the orbital forcing of cyclical sedimentary sequences result from the assumption of mutually exclusive standpoints, for example, that cyclicity results entirely from orbital forcing, or is completely unrelated to it. But, as demonstrated by Mundil et al. (2003) and Meyers (2008), cyclicity in shallow-marine carbonates may display both Milankovitch and nonMilankovitch frequency components. Assuming that the precession and eccentricity cycles were capable of generating metre-scale changes in sealevel during the Triassic, potential Milankovitchfrequency signals may be preserved, among other signals that are noncyclical, or that operated on frequencies not yet understood. Clearly, there remains much to be learned on the depositional controls of cyclically bedded sediments and their stratigraphic application. I express my gratitude to the volume editor, Spencer Lucas, for the invitation to contribute this chapter. Additionally, I extend my thanks to the following individuals for providing thoughtful and informative reviews that improved the manuscript substantially: Gerhard Bachmann, Heinz Kozur, Karl Krainer, Norman Silberling, James Ogg and Robert Weems.

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Triassic conodonts and their role in stage boundary definition MICHAEL J. ORCHARD Geological Survey of Canada, 625 Robson St., Vancouver, British Columbia V6B 5J3 Canada (e-mail: [email protected]) Abstract: Conodonts have played an important role in the construction of a Triassic timescale. Each of the stage boundaries is reviewed in the context of their evolving conodont faunas. The base Triassic (Induan) is defined by the appearance of Hindeodus parvus, which developed from H. praeparvus; a parallel zonation is provided by Neogondolella species. For the Induan– Olenekian boundary, the appearance of Neospathodus waageni sensu lato within a plexus of similar species is favoured as the defining datum; Borinella and Eurygnathodus also appear about this time. The base of the Middle Triassic Anisian stage lies close to the appearance of Chiosella, with Triassospathodus and Spathicuspus characterizing the late Olenekian, and Gladigondolella tethydis and Nicoraella confined to the Anisian. Proxies for the Anisian–Ladinian boundary, which is defined by an ammonoid, are the first Budurovignathus species. The basal Carnian, also defined by an ammonoid, lies close to the first metapolygnathids, including M. polygnathiformis and M. tadpole. The Carnian– Norian boundary interval is characterized by many new taxa in Canada, but only a few species are common to Tethys, notably Metapolygnathus ex gr. M. echinatus. The Norian–Rhaetian boundary is likely to be based on evolution in Misikella in Tethys, with concurrent changes recognized in North American Epigondolella.

Conodonts arrived late on the stage of early Mesozoic timescale definition, yet they have become increasingly important in helping to define the Global Stratotype Section and Points (GGSP) of Triassic stages. In just 40 years of modern study, these microfossils have been used to define one GSSP (base Induan), feature in two others that are defined (Ladinian, Carnian), and are favoured as indices for all of the remaining boundaries under consideration. This is not an unusual situation in the Palaeozoic, but in the Triassic conodonts have competed with traditional ammonoid indices, unlike later Mesozoic periods wherein conodonts no longer existed. Although there is no denying the pivotal role of ammonoids in Triassic timescale studies, conodonts benefit from their greater abundance, widespread distribution, and great resilience. Hence, they can often provide a continuous bed-by-bed sampling record, and occur in a variety of strata representing marine habitats ranging from deepocean to shallow-shelf, as well as in rocks metamorphosed to at least greenschist grade. Of course, like many other common fossils, they offer no solution to the dating and correlation of nonmarine strata. This paper reviews the current status of conodont succession around each Triassic stage boundary, and shows how they may be used to define Triassic GSSPs (Fig. 1). In the following account, the authorship of species is noted only at their first mention. Illustrated specimens are deposited in the National Type Collection of Invertebrate and Plant Fossils at the Geological Survey of Canada, 601 Booth Street, Ottawa, Ontario K1A 0E8 Canada.

Base Triassic, Changshingian– Induan boundary (Figs 2 & 3) The GSSP of the Permian– Triassic Boundary (PTB) was ratified by the International Union of Geological Sciences in 2001 (Yin et al. 2001). It is defined at the First Appearance Datum (FAD) of the conodont Hindeodus parvus (Kozur & Pjatakova) within the H. praeparvus Kozur– H. parvus– Isarcicella isarcica (Huckriede) lineage (Ding et al. 1996; Kozur 1996) at the base of Bed 27c of Meishan section D, Changxing County, Zhejiang Province, South China. Focus on this boundary led to a substantial revision of Hindeodus and the derivative Isarcicella (Orchard 2007b, p. 97 –98). Forty years ago, only single species of Hindeodus (‘Anchignathodus’ typicalis Sweet) and Isarcicella (‘Spathognathodus’ isarcicus) were identified in PTB strata. The impetus generated by research on the boundary has now resulted in the differentiation of at least 25 species of these two genera, including a subdivision of H. latidentatus Kozur, Mostler & Rahimi-Yazd, the originally proposed precursor of H. parvus, into two subspecies. One of these, H. praeparvus, is now being regarded as the immediate forerunner of the PTB index, whereas H. latidentatus is part of an entirely different Hindeodus lineage according to Nicoll et al. (2002). Similarly, H. parvus and other derivatives have been subdivided, and the phylogeny of the group is becoming increasingly refined (Nicoll et al. 2002; Perri & Farabegoli 2003) with many stratigraphically useful species now identified (Fig. 2). Hindeodus

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 139– 161. DOI: 10.1144/SP334.7 0305-8719/10/$15.00 # The Geological Society of London 2010.

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originated during the Permian. According to Korte et al. (2004), there are additional species that are missing from the latest Permian part of the GSSP succession at Meishan. Apparently, several lineages of neogondolellaceans can be differentiated among PTB populations, many of which were originally submerged in Neogondolella carinata (Clark), a species that commonly occurs in Griesbachian strata. Clarkina may be a suitable name for one of these lineages. Another important line that provides zonal indices links N. nassichuki Orchard with the progressively younger N. krystyni Orchard, N. discreta Orchard & Krystyn, and Sweetospathodus kummeli (Sweet) (Orchard 2007b).

Induan – Olenekian boundary (Figs 4 & 5)

Fig. 1. The Triassic timescale, summarizing some key conodont taxa for boundary definition. Numerical scale after Brack et al. 2005, Lehrmann et al. 2006, Ovtcharova et al. 2006, Galfetti et al. 2007, Kozur & Bachmann 2008 and Schaltegger et al. 2008.

parvus has a worldwide distribution not only throughout Tethys but also in western Pangaea (Paull & Paull 1994) and the Boreal region (Henderson & Baud 1996), and it occurs both in shallow water and pelagic deposits (Kozur 1996). In some biofacies, Hindeodus occurs alone (e.g. Werfen facies), but elsewhere Neogondolella (Clarkina of authors) is also common due to palaeoecological factors, as discussed by Nicoll et al. (2002). In these cases, the succession within Neogondolella provides a parallel zonation, as it does throughout the Late Permian. One particularly useful PTB species of Neogondolella is N. taylorae Orchard, which characterizes the lowest Triassic at Selong (Orchard et al. 1994) and Spiti (Orchard & Krystyn 1998). The forerunner of the latter species is N. zhejiangensis (Mei) (¼ Clarkina praetaylorae Kozur: Kozur 2004; Chen et al. 2008). Reports of N. taylorae in latest Permian strata in Iran (Kozur 2005) and China (Ji et al. 2006) suggest that both that species, and several others first described from the basal Triassic,

The conodont succession around the Induan– Olenekian Boundary (IOB) is best known from two localities: the West Pingdingshan Section in Chaohu, Anhui Province, eastern China (Zhao et al. 2007, 2008a, b), and the Mud section in the Pin valley of Spiti, India (Krystyn et al. 2007a, b; Orchard & Krystyn 2007; Orchard 2007e). The latter locality is favoured as GSSP because of its better ammonoid record. The first appearance of the conodont Neospathodus waageni Sweet sensu lato has been chosen as the best datum for definition, as suggested by Tong et al. (2004a). In Chaohu, this datum lies 26 cm below the FAD of poorly preserved ammonoids assigned to Flemingites and Euflemingites (Tong et al. 2004b; Zakharov 2004). In Spiti, the boundary strata include a Gyronites fauna that is replaced by a succession of flemingitid ammonoid faunas (Krystyn et al. 2007a, b) (Fig. 4). In spite of its widespread distribution, both the sub-speciation and origin of Neospathodus waageni is the subject of ongoing research. In general, the late Induan (Dienerian) conodont record is low in diversity and most commonly consists of specimens assigned to Neospathodus dieneri Sweet. Zhao et al. (2007) distinguished three morphotypes of this species based on the posterior denticle formation, and suggested that morphotype 3 of N. dieneri (with reduced posterior denticles) gave rise to N. waageni. Orchard (2007b) suggested an alternative origin in N. pakistanensis Sweet, which has posterior denticulation similar to that of N. waageni. All these taxa are morphologically similar and differ largely in the characteristics of the posterior denticles and in the relative length: height ratio of the elements. Again, this reflects the initial stages in the major diversification of Neospathodus around the IOB, which almost certainly included changes in their multielement configuration, too. Little is known about the apparatuses of these taxa at the IOB, although later

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Fig. 2. Conodont and ammonoid zones of the Permian– Triassic boundary and the ranges of conodont taxa within the Hindeodus– Isarcicella biofacies (top) and the Neogondolella (¼ Clarkina of authors) biofacies (bottom). Adapted from Orchard & Krystyn 1998, Nicoll et al. 2002, Perri & Farabegoli 2003, Kozur 2005, Ji et al. 2006, Luo et al. 2008 and Chen et al. 2008. Abbreviations: Hin., Hindeodus; Is., Isarcicella; Me., Merrillina; Ng., Neogondolella.

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Smithian waageni morphotypes are believed to be examples of Novispathodus. Zhao et al. (2004) identified three variants of N. waageni, two of which (N. waageni eowaageni Zhao & Orchard and N. waageni waageni) are now regarded as subspecies, and one of which (N. posterolongatus Zhao & Orchard) has subsequently been given species status. These three taxa appeared at slightly different levels in the Chaohu section: N. waageni eowaageni first, followed by N. posterolongatus and then N. waageni waageni. This is not clearly the case in Spiti, where Orchard & Krystyn (2007) emphasized the intraspecific variation of N. waageni by differentiating five early Olenekian morphotypes. Zhao et al. (2008a) added a sixth morphotype. The stratigraphic value of these variants has yet to be clearly demonstrated, so the FAD of N. waageni sensu lato is currently taken as the base of the Olenekian. However, the variants do serve to emphasize the tremendous evolutionary radiation that was taking place at the beginning of the Olenekian (early Smithian) (Orchard 2007b). In the Lower Triassic, China was located in the eastern part of low latitude Tethys, whereas Spiti represents higher palaeolatitude Gondwanaland margin. The differing faunal provincialities in these two regions led to some differences in the conodont faunas around the IOB, particularly the abundance of Neospathodus pakistanensis and Borinella nepalensis (Kozur & Mostler) in Spiti, and the absence of the latter taxon in China. Other conodont guide fossils for the IOB include Neospathodus chii Zhao & Orchard, N. concavus Zhao & Orchard, and N. cyclodontus Zhao & Orchard, which are derivatives of N. dieneri; N. spitiensis Goel, which is the

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successor of N. posterolongatus; and the peculiar platform elements Eurygnathodus costatus Staesche and E. hamadai Koike. The first representatives of Discretella? and Spathicuspus also appear in the early Olenekian (Fig. 4).

Olenekian– Anisian boundary (Figs 6 & 7) The most complete ammonoid succession known from the Olenekian –Anisian boundary (OAB) interval in Tethys is within condensed Hallstatt limestone facies at Desli Caira in Dobrogea, Romania, where the latest Olenekian Deslicairitesbearing ammonoid beds are succeeded by beds with early Anisian Paracrochordiceras and Japonites, and then by an interval characterized as the Aegeiceras ugra beds (Graˇdinaru et al. 2007). Other significant biotic changes involving conodonts, orthocerids, nautilids, and foraminiferids are recognized at this palaeontologically-defined boundary (Graˇdinaru et al. 2007), which is located in between two short normal magnetic polarity intervals, and close to the end of a gradual increase in d13C isotope values. In Nevada, USA, ammonoids of the early Anisian Japonites welteri Zone overlie those of the Spathian Neopopanoceras haugi Zone (Bucher 1989; Guex et al. 2005a, b), whereas in Union Wash, California (Stone et al. 1991), an additional latest Spathian interval, the Courtilloticeras stevensi beds are differentiated (Galfetti et al. 2007). The ammonoid zonation established in western and northern Canada (Tozer 1994) places the Silberlingites mulleri Zone, which is younger than the basal Anisian, above the late Spathian Keyserlingites

Fig. 3. (Continued) Key conodonts from the Permian–Triassic (Changshingian–Induan) boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1– 3. Neogondolella krystyni Orchard. GSC 101739, from Griesbachian of Guling, Spiti (GSC cur. C– 302288; Guling3). 4– 6. Neogondolella discreta Orchard & Krystyn. GSC 101746, from Griesbachian of Guling, Spiti (GSC cur. C–302288; Guling5). 7, 8. Sweetognathodus kummeli (Sweet). GSC 132596, from Dienerian of Guling, Spiti (GSC V– 000426; GU– 6). x100. 12. Neogondolella carinata (Clark). GSC 101734, from Griesbachian of Guling, Spiti (GSC cur. C–302288; Guling3). 13. Hindeodus parvus (Kozur & Pjatakova). Morphotype 3 of Mei 1996. GSC 101761, from Griesbachian of Lalung, Spiti (GSC cur. C– 302291; Lalung3– 1). x100. 14. Isarcicella staeschei Dai & Zhang. GSC 101757, from Griesbachian of Lingti, Spiti (GSC cur. V –000416; L4). x125. 9– 11. Neogondolella nassichuki Orchard. Holotype, GSC 01685, from Griesbachian of Selong, Tibet (GSC cur. C– 301264; U50). 15. Hindeodus latidentatus Kozur, Mostler & Rahmini-Yazd. GSC 101664, from late Changhsingian of Meishan, China (GSC loc. C–158625; D12). 16. Hindeodus praeparvus (Kozur). GSC 101771, from Griesbachian of Lingti, Spiti (GSC cur. V –000416; L4). x125. 17, 18. Neogondolella zhejiangensis (Mei). GSC 101675, from Griesbachian of Selong, Tibet (GSC cur. C–301263; M50). 19–21. Neogondolella taylorae Orchard . Holotype GSC 101673, from Griesbachian of Selong, Tibet (GSC cur. C– 301264; U50).

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Fig. 4. Ammonoid zones and conodont ranges of the Induan– Olenekian boundary in Mud, Spiti, and ranges of conodont taxa in Chaohu, China plotted against bed numbers. Correspondence between the two sections is conjectural. Spiti data based on Krystyn et al. 2007a, b, Orchard & Krystyn 2007, Orchard 2007e, and unpublished data. Chaohu data based on Zhao et al. 2007, 2008a, b. Abbreviations: Ns., Neospathodus; Eu., Eurygnathodus; Bo., Borinella; Dis., Discretella.

subrobustus Zone. A locality in Guandao in Guizhou Province, south China lacks a significant ammonoid fauna but does contain an important conodont succession as well as radiometrically-dated tuff beds (Lehrmann et al. 2006). Published conodont data from each of these localities (Orchard 1994, 2008; Graˇdinaru et al. 2007) show the traditional ammonoid-based OAB closely coincides with the appearance of the conodont Chiosella timorensis (Nogami), whereas upper Olenekian (Spathian) strata are dominated by

Triassospathodus ex gr. homeri (Bender). The conodont C. gondolelloides (Bender), the precursor of C. timorensis, straddles the boundary, appearing first in the Deslicairites beds in Romania, and in the stevensi beds in California. The relatively expanded section at Guandao, China, sampled in detail for conodonts, provides a useful corroboration of the conodont succession (Orchard et al. 2007b). In both the Guandao and Desli Caira sections (Orchard et al. 2007a), parallels may be seen in the range of Triassospathodus homeri and

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Fig. 5. Key conodonts from the Induan– Olenekian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1, 2. Neospathodus pakistanensis Sweet. GSC 120379, from Smithian of Ellesmere Island, Canada (GSC loc. C–303178). 3, 4. Neospathodus posterolongatus Zhao & Orchard. GSC 132571, from Smithian of Spiti (GSC cur. V–000417; M03– 13). x60. 5, 6. Neospathodus spitiensis Goel. GSC 132572, from Smithian of Spiti (GSC cur. V–000420; M03–15). x60. 7, 8. Neospathodus dieneri Sweet, Morphotype 3 Zhao & Orchard. GSC 132573, from Dienerian of Ellesmere Island, Canada (GSC cur. 303171; DIC–F2). 9, 10. Eurygnathodus costatus Staesche. GSC 132574, from Smithian of Spiti (GSC cur. V– 000419; M06–13B). 11. Eurygnathodus hamadai (Koike). GSC 132575, from Smithian of Spiti (GSC cur. V– 000419; M06– 13B). x120. 12, 13. Borinella nepalensis (Kozur & Mostler). GSC 132576, from Smithian of Spiti (GSC cur. V –000418; MO6–13A1). 14, 15. Neospathodus waageni Sweet, Morphotype 5. GSC 132577, from Smithian of Spiti (GSC cur. V –000421; M04–15B).

Gladigondolella carinata into the basal Anisian. In China, Spathicuspus spathi (Sweet) also ranges into the basal Anisian and overlaps with the appearance of Gladigondolella tethydis (Huckriede) and Chiosella? n. sp. A Orchard, which are useful

supplementary Anisian conodont markers in Tethyan regions. Neogondolella ex gr. N. regalis Mosher appears slightly later in both China and Romania, although in North America similar forms appear first within the Spathian (Orchard

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Fig. 6. Ammonoid faunas of the Olenekian–Anisian boundary interval at Desli Caira, Romania and the ranges of conodont taxa based on Orchard et al. 2007a. Conodont ranges from the Nanpangjiang Basin at Guandao based on Orchard et al. 2007b. Abbreviations: Ns., Neospathodus; Tr., Triassospathodus; Cr., Cratognathus; Ch., Chiosella; Gl., Gladigondolella.

2007b, 2008). This boundary interval is marked by a significant faunal change as the proportion of T. homeri and S. spathi drops dramatically (Orchard et al. 2007a, Figs 4 –5). Nicoraella germanica (Kozur) is a distinctive and widespread early Anisian species, and Neospathodus triangularis Bender is a typical late Spathian one. Prior to 1995, the differentiation of Chiosella gondolelloides and C. timorensis was not widely accepted (see Orchard 1995; Graˇdinaru et al. 2006), so older literature is ambiguous in the sense that the two species of Chiosella were often combined. Similarly, at least some early reports of C. timorensis from the Spathian of the USA

(Collinson & Hassenmueller 1978) were based on Triassospathodus homeri (Orchard 1994), which completely dominates collections from the late Spathian Haugi Zone. The latter species is similar in some respects to C. timorensis but is much shorter and lacks platform flanges. It was thought to be ancestral to Chiosella (Bender 1970), but this is not supported by their significantly different multielement apparatuses (Orchard 2005). Chiosella n. sp. A, characterized by a strong terminal cusp, is a form that appears slightly later in Tethys than the other Chiosella species, more or less coincident in both Guandao sections with a ‘second wave’ of Chiosella that is separated from the first by a

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Fig. 7. Key conodonts from the Olenekian–Anisian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1– 3. Gladigondolella tethydis (Huckriede). GSC 132578, from Anisian of Wadi Alwa, Oman (GSC cur. C– 158458; 118C). 4. Nicoraella germanica (Kozur), GSC 101560, from Anisian of Nama, China (GSC cur. V–000424; NMC–104). 5. ‘Neospathodus’ triangularis Bender. GSC 132579, from Spathian of Tobin Range, Nevada, USA (GSC cur. O– 64700).

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narrow interval from which the genus has not been recovered (Orchard et al. 2007a, Figs 4 –5).

Anisian – Ladinian boundary (Figs 8 & 9) The GSSP of the Anisian –Ladinian boundary (ALB) is defined at the appearance of the ammonoid Eoprotrachyceras curionii in the Caffaro River bed near Bagolino in the Italian Dolomites (Brack et al. 2005). The appearance of the conodont Budurovignathus praehungaricus (Kovacs) near the top of the underlying Nevadites secedensis Zone provides a proxy for the ALB. This conodont species probably developed from Neogondolella aequidentata Kozur, Krainer & Lutz and, in turn, is the precursor of the younger Budurovignathus species that are important in Ladinian biochronology. Additional and often more abundant conodont taxa of the ALB beds belong to the groups of Neogondolella constricta (Mosher & Clark), Paragondolella excelsa Mosher, and N. alpina (Kozur & Mostler). However, most of these taxa appear well below the boundary, and several range across it. Other useful ALB taxa, including Neogondolella postcornuta (Kovacs), ‘Ng.’ trammeri (Kozur), and Gladigondolella malayensis Nogami, appear confined to Tethyan successions. In North America, the most complete ALB successions occur in the Prida Formation (Star Peak Group) in the Humboldt Range of northwestern Nevada, specifically at the ‘Fossil Hill’ and ‘Saurian Hill’ sites documented by Silberling & Nichols (1982). This was the type area where Silberling & Tozer (1968) defined (in ascending order) the upper Anisian Gymnotoceras rotelliformis, Parafrechites meeki, and Frechites occidentalis zones, and the lower Ladinian Eoprotrachyceras subasperum Zone, and where more recent work by Bucher & Orchard (1995) intercalibrated the ammonoid and conodont faunas. The Occidentalis Zone at the top of the Anisian originally comprised, in ascending order, the Nevadites hyatti beds, Nevadites humboldtensis beds, Paranevadites furlongi beds, and Paranevadites gabbi beds, although more recently the first of these have been dropped (Ovtcharova et al. 2006). The first occurrence of protrachyceratids has been regarded as defining

the base of the Ladinian stage, which includes a succession of differing Eoprotrachyceras beds, of which the second comprise the Eoprotrachyceras lahontanum beds (Bucher & Orchard 1995). Two levels of significant conodont faunal change are identified within the ALB interval in Nevada, although details of the succession are unpublished. One level corresponds to the base of the Occidentalis Zone (N. hyatti beds), the other to the Eoprotrachyceras lahontanum beds. The first is marked by the appearance of several new taxa, and the reappearance of others. Both Neogondolella cornuta Budurov & Stefanov, a member of the N. constricta group, and Paragondolella liebermani (Kovacs & Krystyn) become more common, and derivatives of the N. szaboi (Kovacs)–N. alpina stock, including N. ex gr. N. pridaensis (Nicora, Kovacs & Mietto) and N. n. sp. B, are often dominant; N. aequidentata also appears. The second distinctive conodont datum occurs in the Eoprotrachyceras lahontanum beds and is marked by the appearance of common Budurovignathus praehungaricus, a proxy for the ALB, rare B. gabriellae Kozur, Krainer & Mostler, and new Paragondolella species.

Ladinian–Carnian boundary (Figs 10 & 11) The Ladinian–Carnian boundary (LCB) is defined at the appearance of the ammonoid Daxatina canadensis at the Prati di Stuores/Stuores Wiesen section in the Italian Dolomites (Manco et al. 2004; Mietto et al. 2007). Unfortunately, conodonts are very scarce, due to the high sedimentation rate in this part of the section, but they become more frequent in higher beds associated with the ammonoid Trachyceras aon, the level at which the base of the Carnian was traditionally drawn. The conodont Metapolygnathus ex gr. M. polygnathiformis (Budurov & Stefanov), characteristic of the Carnian, first occurs close to the Daxatina datum, although it is very rare at the GSSP site. Other areas that have provided key data on the LCB are sections in Spiti, India; New Pass in Nevada, USA; and the Liard Basin in northeastern British Columbia, Canada. The classical sections of the Spiti Valley in the Western Himalayas, India, have shown that the appearance of

Fig. 7. (Continued) 6– 8. Neogondolella ex gr. regalis Mosher. GSC 101543, from Anisian of Guandao, China (GSC cur. C– 306556; GDL50). 9– 11. Chiosella gondolelloides (Bender). GSC 101554, from Anisian of Guandao, China (GSC cur. V– 000423; OU-2). 12– 14. Chiosella timorensis (Nogami). GSC 101558, from Anisian of Guandao, China (GSC cur. C–306556; GDL-50). 15, 16. Triassospathodus ex gr. homeri (Bender), GSC 101563, from Spathian of Desli Caira, Romania (GSC cur. C– 304288; 9036A). 17, 18. Chiosella? n. sp. A. GSC 101552, from Anisian of Guandao, China (GSC cur. V–000425; O40). 19, 20. Spathicuspus spathi (Sweet). GSC 101542, from Anisian of Guandao, China (GSC cur. V–000422; O35).

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Fig. 8. Ammonoid zones and conodont ranges of the Anisian–Ladinian boundary in the GSSP site of Bagolino, Italy, and in Nevada, USA. Based on Brack et al. 2005 (Italy), Bucher & Orchard 1995, and unpublished data (Nevada). No precise correspondence of Anisian zones intended. Abbreviations: Ng., Neogondolella; Para., Paragondolella; Bud., Budurovignathus.

Metapolygnathus ex gr. M. polygnathiformis predates the incoming of Daxatina by several metres, in spite of the slow accumulation rates of the Spiti sections (Balini et al. 2004; Krystyn et al. 2004). South Canyon in New Pass, Nevada, USA, appears to be the best known section for the boundary interval so far described in North America (Silberling & Tozer 1968). There, indices for both the

Frankites sutherlandi (traditionally Ladinian) and Trachyceras desatoyense zones (lowest Carnian) co-occur with both Daxatina and M. ex gr. M. polygnathiformis (Balini et al. 2007; Orchard & Balini 2007). The new definition of the LCB places the upper part of the Sutherlandi Zone in the Carnian. Conodonts from ammonoid-bearing beds of the LCB interval in British Columbia, including

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Fig. 9. Key conodonts from the Anisian– Ladinian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1– 3. Budurovignathus praehungaricus (Kovacs). GSC 132580, from Ladinian of Saurian Hill, Nevada, USA (GSC cur. C– 300239; HB508C). 4– 6. Budurovignathus aff. B. hungaricus (Kozur & Vegh) GSC 120242, from Ladinian of Yukon, Canada (GSC cur. C– 087031; TOR-1). 7– 9. Neogondolella aequidentata Kozur, Krainer & Lutz. GSC 132581, from Ladinian of Saurian Hill, Nevada, USA (GSC cur. C–300327; HB508A).

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Fig. 10. Ammonoid zones and conodont ranges of the Ladinian– Carnian boundary GSSP in Stuores, Italy, and the ammonoid zones and conodont zones and ranges of the interval in British Columbia. Based on Manco et al. 2004, Mietto et al. 2007 (Italy) and Orchard 2007c (B.C.). Abbreviations: Para., Paragondolella; Bud., Budurovignathus; Ng., Neogondolella; Meta., Metapolygnathus.

Fig. 9. (Continued) 10, 11. Neogondolella alpina (Kozur & Mostler). GSC 132582, from Anisian of Saurian Hill, Nevada, USA (GSC cur. C– 201588; HB527). 12, 13. Neogondolella cornuta Budurov & Stefanov. GSC 132583, from Ladinian of Saurian Hill, Nevada, USA (GSC cur. C– 201585; HB524). x65. 14, 15. Neogondolella constricta (Mosher & Clark). GSC 132584, from Anisian of Fossil Hill, Nevada, USA (GSC cur. C– 300212; FH15). 16. Neogondolella ex gr. N. pridaensis (Nicora, Kozur & Mietto). GSC 132585, from Anisian of Fossil Hill, Nevada, USA (GSC cur. C– 300231; FH56). 17. Neogondolella n. sp. B. GSC 132586, from Ladinian of Saurian Hill, Nevada, USA (GSC cur. C–201563; HB504). 18, 19. ‘Neogondolella’ trammeri (Kozur). GSC 132587, from Anisian of Szentkiralyszabadja, Hungary (GSC cur. C–301154; 6– 5– 70). 20, 21. Neogondolella aff. liebermani (Kovacs & Krystyn). GSC 132588, from Anisian of Fossil Hill, Nevada, USA (GSC cur. C– 300214; FH17).

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Fig. 11. Key conodonts from the Ladinian–Carnian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise.

TRIASSIC CONODONTS AND THEIR ROLE IN STAGE BOUNDARY DEFINITION

several new species, were described recently by Orchard (2007c). The interval can be very broadly divided into a lower part characterized by Budurovignathus mungoensis (Diebel) and an upper part with Mosherella newpassensis (Mosher). This division is also quite evident at New Pass in Nevada (Orchard & Balini 2007), although at that locality there is an overlap between the two species that is not observed in British Columbia, probably as a result of insufficient sampling between ammonoid beds. The succession of Budurovignathus and Mosherella species lacks clear speciation events to serve as biochronological markers, whereas the development of Metapolygnathus from Paragondolella and the diversification within the former facilitates a new conodont zonation for the LCB in North America. Paragondolella inclinata (Kovacs) represents the root stock from which one evolutionary trend involved the progressive reduction of the anterior platform. This gave rise, in ascending order, to the inclinata, intermedius, and tadpole zones, with the first subdivided into sulcata and acuminatus subzones (Orchard 2007c). These zones subdivide the North American Sutherlandi ammonoid Zone and adjacent zones (Tozer 1994), with the acuminatus–intermedius zonal boundary closely coincident with the appearance of Daxatina; Metapolygnathus polygnathiformis sensu stricto appears a little before this boundary. The tadpole Zone, as interpreted by Orchard (2007c), extends from high in the Sutherlandi Zone through the lower Carnian Austrotrachyceras obesum and Sirenites nanseni zones. In contrast to North America, several Budurovignathus species are common in the Tethyan sections. This includes B. diebeli (Kozur & Mostler), earlier proposed as a potential index for the LCB. These taxa may be endemic because only B. mungoensis is known from North America. Conversely, Mosherella species, common in North America, appear to be absent in Tethys.

153

Carnian– Norian boundary (Figs 12 & 13) In North America, ammonoid successions in Nevada and British Columbia led to a proposal that the base of the CNB be assigned to the base of the Stikinoceras kerri ammonoid Zone, which overlies the Klamathites macrolobatus Zone (Silberling & Tozer 1968; Tozer 1994). In Tethys, this level is approximately coeval with a boundary drawn between the Anatropites and Guembelites jandianus ammonoid zones (Krystyn 1980). Preliminary biostratigraphic work in western Canada (Orchard 1983; Orchard & Tozer 1997) showed that the base of the Stikinoceras kerri Zone falls within the range of Metapolygnathus primitius (Mosher) and often corresponds to the appearance of Norigondolella navicula (Huckriede). However, the incoming of the latter species, although a useful guide, is an unsatisfactory datum for definition because of apparent facies control on its appearance. This being the case, Orchard et al. (2000) sought better correspondence of faunal changes in ammonoids, conodonts, and radiolarians and suggested that the datum at the base of the Klamathites macrolobatus Zone might serve as a CNB because both M. primitius and M. ex gr. M. communisti Hayashi appeared at about that level in British Columbia. However, much more detailed work at Black Bear Ridge Ridge on Williston Lake in northeastern British Columbia has now shown that there is in fact good correspondence between ammonoid, bivalve, and conodont faunal change at, or close to, the traditional level (Orchard et al. 2001, 2007a; McRoberts 2007). Several new upper Carnian conodont zones are now identified in British Columbia based on forms currently combined as Metapolygnathus (Fig. 12; Orchard et al. 2001; Orchard 2007a, d). The first appearance of Metapolygnathus ex gr. M. echinatus (Hayashi) lies close to the base of the Kerri Zone and coincides with a major faunal turnover whereby subsequent faunas are

Fig. 11. (Continued) 1, 2. Metapolygnathus ex gr. M. polygnathiformis (Budurov & Stefanov). GSC 132589, from Carnian of Brown Hill, BC, Canada (GSC cur. C–301166; BH– B5). 3– 5. Metapolygnathus tadpole (Hayashi). GSC 120361, from Carnian of Liard River, BC, Canada (GSC cur. O–68272). 6, 7. Mosherella newpassensis (Mosher). GSC 101827, from Carnian of New Pass, Nevada, USA (GSC cur. C–306570; 105). 8– 10. Metapolygnathus intermedius Orchard. GSC 120359 from Carnian of Besa River, BC, Canada (GSC cur. C– 305045; NWB–1). 11–13. Metapolygnathus polygnathiformis (Budurov & Stefanov). GSC 120362 from Carnian of Mt. Trimble, BC, Canada (GSC cur. C– 304828; TRIM–2). 14–16. Metapolygnathus acuminatus Orchard. GSC 120358 from Carnian of Mt. Trimble, BC, Canada (GSC cur. C– 304781; TRIM–2A). 17–19. Paragondolella inclinata (Kovacs). GSC 120355, from Ladinian of Brown Hill, BC, Canada (GSC cur. C– 305932; BHE– 1). 20. Budurovignathus ex gr. B. mungoensis (Diebel). GSC 120393, from Ladinian of New Pass, Nevada, USA (GSC cur. C– 306610; SC– B2).

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Fig. 12. Ammonoid zones and conodont zones and ranges of the Carnian– Norian boundary in British Columbia, and the ranges of selected conodonts at Pizzo Mondello. After Orchard 2007a and unpublished (B.C.) and Nicora et al. 2007 (Italy). Abbreviations: Carn., Carniepigondolella; Meta., Metapolygnathus; Ep., Epigondolella; Nori., Norigondolella.

dominated by the primitius group. This datum also corresponds to the appearance of new halobiid species (McRoberts 2007) and is consequently now favoured as the CNB datum in North America. A second GSSP candidate for the CNB is in Tethys, at Pizzo Mondello in Sicily (Nicora et al. 2007). Preliminary conodont data show both similarities and significant differences in the faunal succession compared with North America. This is compounded by unresolved nomenclatural issues. Compared with Canada, Carniepigondolella ex

gr. C. samueli (Orchard) ranges much higher, and Epigondolella ex gr. E. quadrata Orchard appears much earlier. Also, whereas M. primitius, M. n. sp. I, M. n. sp. P, M. n. sp. Q, and M. n. sp. Y (see Orchard 2007a for illustrations) are common at Black Bear Ridge, they are rare or absent at Pizzo Mondello, where M. communisti is far more common. In Tethys, the first occurrence of M. communisti lies close to the base of the Jandianus Zone (Krystyn 1980). The appearance of this species and its associates M. parvus (Kozur) and

Fig. 13. Key conodonts from the Carnian–Norian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1– 3. Epigondolella quadrata Orchard. GSC 132590, from Norian of Black Bear Ridge, BC, Canada (GSC cur. C–426568; BBR –32). 4– 6. Norigondolella navicula (Huckriede). GSC 131172, from the Norian of Black Bear Ridge, BC, Canada (GSC cur. C–426562; BBR –26). 7– 9. Metapolygnathus ex gr. M. echinatus (Hayashi). GSC 131168, from Norian of Pardonet Hill, BC, Canada (GSC cur. C–305903; PHE– 23). 10–12. Metapolygnathus primitius (Mosher). GSC 132591, from Norian of Pardonet Hill, BC, Canada (GSC cur. C– 305903; PHE– 23).

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Fig. 14. Ammonoid zones and conodont zones and ranges around the Norian– Rhaetian boundary in Steinbergkogel, Austria, and ammonoid and radiolarian zones of the interval in North America. Based on Krystyn et al. 2007c (Austria), and Orchard et al. 2007c, d, Carter & Orchard 2007 (North America). Abbreviations: Ep., Epigondolella; On., Oncodella; Mis., Misikella; Pg., Parvigondolella; Zieg., Zieglericonus.

Fig. 13. (Continued) 13, 14, 15. Metapolygnathus n. sp. I. GSC 131166, from Carnian of Black Bear Ridge, BC, Canada (GSC cur. C–305879; BBR –15). 16. Metapolygnathus n. sp. Q. GSC 131240, from Carnian of Black Bear Ridge, BC, Canada (GSC cur. C–305869; BBR– 5). 17. Metapolygnathus n. sp. P. GSC 131174, from Carnian of Black Bear Ridge, BC, Canada (GSC cur. C– 426548; BBR– 13A). 18– 20. Metapolygnathus n. sp. G. GSC 132592, from Carnian of Black Bear Ridge, BC, Canada (GSC cur. C–305868; BBR–4). 21. Metapolygnathus n. sp. Y. GSC 131175, from Carnian of Black Bear Ridge, BC, Canada (GSC cur. C–426553; BBR– 16). 22– 23. Metapolygnathus communisti Hayashi. GSC 132593, from Norian of Black Bear Ridge, BC, Canada (GSC cur. C– 307824; BBR– 21a). 24. Metapolygnathus parvus (Kozur). GSC 132597, from Norian of Black Bear Ridge, BC, Canada (GSC cur. C–307817; BBR–18h).

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Fig. 15. Key conodonts from the Norian–Rhaetian boundary. Catalogue numbers of specimens are given first, followed by age and location, including GSC curation number (cur.); followed by field sample number. All x80 (for which scale bar ¼ 200 microns), unless stated otherwise. 1– 3. Epigondolella englandi Orchard. GSC 95290, from Rhaetian of Laberge, Yukon, Canada (GSC cur. C– 87005; TOT11– 19). x100. 4. Epigondolella bidentata Mosher. GSC 132594, from Rhaetian of Kennecott Point, Haida Gwaii, Canada (GSC cur. C– 158533; R3A). x110. 5. Epigondolella mosheri (Kozur & Mostler), Morphotype A Orchard. GSC 120347, from Rhaetian of New York Canyon, Nevada, USA (GSC cur. C–116526; NY– 13). x110. 6. Epigondolella mosheri (Kozur & Mostler), Morphotype C Orchard. GSC 132595, from Rhaetian of Kennecott Point, Haida Gwaii, Canada (GSC cur. C–150179; SKU–B9). 7, 8. Norigondolella steinbergensis (Mosher). GSC 101517, from Norian of Baja California, Mexico (GSC cur. C– 173391; INA–3). 9, 10. Misikella hernsteini (Mostler), GSC 101513, from Rhaetian of Baja California, Mexico (GSC cur. C–173442; COC–11). 11. Zieglericonus rhaeticus Kozur & Mock. GSC 101511, from Rhaetian of Baja California, Mexico (GSC cur. C– 173442; COC– 11). x150. 12, 13. Misikella posthernsteini Kozur & Mock. GSC 95294, from Rhaetian of Kennecott Point, Haida Gwaii, Canada (GSC cur. C–156526; R18A). 14. Oncodella paucidentata (Mostler). GSC 101509, from Rhaetian of Baja California, Mexico (GSC cur. C–173442; COC–11). x110.

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M. echinatus seem to be a good option for intercontinental correlation.

Norian – Rhaetian boundary (Fig 14 & 15) The Rhaetian stage was introduced for the uppermost Triassic Ko¨ssen Beds in the Northern Calcareous Alps with the brachiopod Rhaetavicula contorta as the guide form. According to Golebiowski (1990), R. contorta begins within Unit 2 of the Hochalm Member of the Ko¨ssen Formation, also the level of the first appearance of Misikella posthernsteini Kozur & Mock (Kozur & Bachmann 2005). However, in places its precursor M. hernsteini (Mostler) occurs at the base of the Ko¨ssen Beds (Golebiowski 1986, 1990; Krystyn 1990). Recent deliberation has considered the relative merits of the FO of the two Misikella species as defining taxa for the Norian-Rhaetian boundary (NRB) at the Steinbergkogel section in Austria (Krystyn et al. 2007b). In North American ammonoid biochronology, the base of the Paracochloceras amoenum Zone, which can be correlated with the base of the Cochloceras suessi Zone in the Tethys (Kozur 1996, 2003), is favoured as a Rhaetian base (Orchard & Tozer 1997). In western North America, Misikella is rare and does not provide a practical guide fossil for boundary definition. Rather, species of Epigondolella are common and apparently range higher into the Rhaetian than they do in Tethys: the base of the Amoenum Zone corresponds closely to the FAD of E. mosheri (Kozur & Mostler) Morphotype A Orchard in both western Canada and Nevada, USA (Orchard & Tozer 1997, p. 685; Orchard et al. 2007c). In British Columbia, rare Misikella posthernsteini are only known from the upper Rhaetian Choristoceras crickmayi Zone (Orchard & Tozer 1997). The most distinctive faunal change for characterizing the base of the Rhaetian in the oceanic realm is the base of the radiolarian Proparvicingula moniliformis Zone (Carter 1993; Carter & Orchard 2004). In the Queen Charlotte Islands, Canada, this radiolarian zone is characterized by the appearance of Epigondolella mosheri Morphotype A, whereas in Baja California Misikella posthernsteini occurs together with radiolarians of the lower P. moniliformis Zone (Whalen et al. 2003; Orchard et al. 2007d). Hence, broad equivalence of the appearance of E. mosheri and M. posthernsteini is implied (see also Kozur 2003). Illustrated specimens were collected by the author and by several other individuals including Tim Tozer, Leo Krystyn, Hugo Bucher, Walter Nassichuk, Beth Carter, Patricia Whalen, Eugen Graˇdinaru, Dan Lehrmann, Wang Hongmei, Wei Jiayong, and Yu Youyi. Peter

Krauss undertook much of the laboratory preparation and SEM photography, and Hillary Taylor helped with microscopy and manuscript preparation. Bob Nicoll and Alda Nicora are thanked for their reviews. This paper summarises achievements that were supported by the Geological Survey of Canada, the Subcommission on Triassic Stratigraphy, and IGCP 467 (2002– 2007). ESS Contribution number 20090242.

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Triassic radiolarian biostratigraphy LUIS O’DOGHERTY1*, ELIZABETH S. CARTER2, SˇPELA GORICˇAN3 & PAULIAN DUMITRICA4 1

Facultad de Ciencias del Mar, Universidad de Ca´diz, 11510 Puerto Real, Spain 2

Geology Department, Portland State University, 97207 – 0751 Portland, USA

3

Paleontolosˇki Insˇtitut Ivana Rakovca ZRC SAZU, Novi trg 2, SI – 1000 Ljubljana, Slovenia 4

Dennigkofenweg 33, CH – 3073 Guemligen, Switzerland *Corresponding author (e-mail: [email protected])

Abstract: This paper summarizes 30 years of research on the biostratigraphy of Triassic radiolarians and presents a correlation of currently-used radiolarian zonations established in North America, Europe, Japan and Far East Russia. An up-to-date stratigraphic distribution of all hitherto described and still valid Triassic genera is provided. This new range chart consists of 282 genera and allows an accurate dating to substage level. It also clearly manifests general trends in radiolarian evolution through the Triassic. The end-Permian extinction, the most severe extinction in the history of radiolarians, was followed by a long recovery until the early Anisian. The middle and late Anisian were then characterized by a rapid explosion of new morphologies. Maximum generic diversity was attained during the early Carnian, but the first severe extinctions also occurred in the Carnian. A progressive decline of diversity took place through the Norian and Rhaetian, and ended in a mass extinction around the Triassic– Jurassic boundary.

Since the revolution that signified the use of the hydrofluoric acid method in extracting radiolarians from hard siliceous rocks, radiolarians have proven to be of great importance in reconstructing the stratigraphy of the Mesozoic, particularly the Triassic System. This methodology, discovered independently by Dumitrica (1970) and Pessagno & Newport (1972), together with the building of a new taxonomic system for Mesozoic radiolarians, opened the way to stratigraphic progress and zonal correlation. This contribution reviews the different biostratigraphic scales proposed in the past 30 years and discusses the validity of correlation. The first radiolarian range charts were proposed almost simultaneously in North America, Japan and Europe (Pessagno et al. 1979; Nakaseko & Nishimura 1979; Yao et al. 1980a; De Wever 1982). One of the most important aspects to be mentioned is the apparently low provincialism displayed by these zonations through the Triassic when compared to the classical schemes proposed for the Jurassic and Cretaceous. Another characteristic feature of these zonations is the accurate calibration to the standard chronostratigraphic stages and substages, which are established basically by means of ammonites and conodonts. But the most striking aspect is the time resolution of these biostratigraphic scales, which are better than those proposed for the Jurassic –Cretaceous. However, the potential use of Triassic radiolarians for stratigraphic purposes

remains largely untapped. The reason for this is the great generic diversity shown by this group during the Triassic, probably the largest known in the history of the group, as has been illustrated in a recent stratigraphic and taxonomic review of genera (O’Dogherty et al. 2009a). The aim of this paper is also to show that the stratigraphic distribution of genera is a powerful tool that allows dating at the substage level and opens new insights for future research.

State of the art Historical review of Triassic radiolarian biostratigraphy In the late 19th and early 20th centuries, the Triassic System was the least studied Mesozoic system in terms of radiolarians. Compared to relatively numerous works on Jurassic and Cretaceous radiolarians (e.g. Ru¨st 1885; Squinabol 1903, and others), virtually nothing was known about Triassic radiolarians until 1970. A rare exception was Ru¨st (1892), who described 21 Triassic species (most from Felso¨o¨rs in Hungary, a well-known radiolarian locality today). The beginning of Triassic radiolarian research thus practically coincides with the discovery of etching techniques for siliceous rocks (Dumitrica 1970; Pessagno & Newport 1972) and with the development of scanning electron microscopy.

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 163– 200. DOI: 10.1144/SP334.8 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Early studies of Triassic radiolarians focused on description of new species and intended to present either all representatives of certain families (Kozur & Mostler 1972, 1978; Dumitrica 1978a, b, 1982a, b, c) or radiolarian assemblages of certain stratigraphic levels (De Wever et al. 1979; Pessagno et al. 1979; Nakaseko & Nishimura 1979; Kozur & Mostler 1979, 1981; Dumitrica et al. 1980). The first radiolarian range charts were proposed in the late 1970s and early 1980s almost simultaneously in North America, Japan and Europe. The hitherto published radiolarian zonations are discussed below and summarized in Figure 1.

Development of zonal schemes and correlation In North America, the study of Triassic radiolarians was initiated by Pessagno et al. (1979), who established two zones, Capnodoce Zone and Pantanellium silberlingi Zone, for the latest Carnian and Norian in Baja California. These zones were further emended and refined by Blome (1984) based on material from different sites in western North America (Oregon, Baja California and Queen Charlotte Islands). Blome (1984) divided the Capnodoce Zone into three subzones, that is, Justium novum, Xipha striata and Latium paucum subzones. He erected a new zone for the middle and late Norian, the Betraccium Zone, which was subdivided into two subzones, Pantanellium silberlingi Subzone (equivalent to former zone of Pessagno et al. 1979) and Betraccium deweveri Subzone. The latest Norian and Rhaetian was studied by Carter (1990, 1993) in the Queen Charlotte Islands. On the basis of unitary associations (Guex 1977, 1991), Carter (1993) recognized Blome’s Betraccium deweveri Zone at the base and divided the Rhaetian stage in two formal zones. These are the Proparvicingula moniliformis Zone, which represents the lower Rhaetian, and the Globolaxtorum tozeri Zone, representing the upper Rhaetian. The Propavicingula moniliformis Zone is subdivided into two distinct assemblages, that is, assemblages 1 and 2, with the second assemblage consisting of four subassemblages, that is, 2a to 2d. The Globolaxtorum tozeri Zone consists of a single assemblage (assemblage 3). Radiolarians older than late Carnian were reported only as single occurrences from tectonically complex areas (e.g. Blome et al. 1988; Blome & Reed 1992; Cordey 1998), so no radiolarian zonation has been established for the Early and Middle Triassic in North America. Blome et al. (1988) also created a zone below their Capnodoce Zone, namely the Pseudostylosphaera Zone, but this zone is rather vague and

has not been employed. In the Middle and Upper Triassic of Oregon, Yeh (1989) recognized four assemblages separated by long empty intervals. At present, the zonations of Blome (1984) and Carter (1993) are the most widely used radiolarian zonations for the Norian and Rhaetian (Fig. 2, see also De Wever et al. 2001). In addition to their high resolution, which enables dating to substage level, these zonations provide ranges of a great number of taxa, which are regularly encountered elsewhere in the Tethys. In Japan, Nakaseko & Nishimura (1979) established three assemblage zones, that is, Tripocyclia cf. acythus, Emiluvia? cochleata, and Capnuchosphaera theloides assemblages, and assumed that all three were Late Triassic in age. Further research revealed that only the Capnuchosphaera theloides assemblage is Late Triassic (Norian) in age, whereas the Tripocyclia cf. acythus and Emiluvia? cochleata assemblages are late Anisian –early Ladinian and late Ladinian, respectively (Kozur & Mostler 1994, Ramovsˇ & Goricˇan 1995, Kozur et al. 1996). In the 1980s, many Japanese researchers independently proposed radiolarian biostratigraphic schemes for the Middle to Upper Triassic (Yao et al. 1980a, b, 1982; Yao 1982, 1990; Kishida & Sugano 1982; Nishizono & Murata 1983; Igo & Nishimura 1984; Kishida & Hisada 1986; Sato et al. 1986; Yoshida 1986; Sashida et al. 1993; Nishizono 1996; for correlations see Sugiyama 1997). The most commonly used scheme in the 1980s and early 1990s was that established by Yao (1982), who divided the Ladinian to Rhaetian interval into three rather long-ranging assemblages. Pre-Ladinian radiolarian stratigraphy was first established by Sugiyama (1992), who presented three new assemblages for the Spathian to middle Anisian. Later, Sugiyama (1997) established 18 zones for the Spathian to the top of the Triassic. The lower two zones are assemblage zones, and the others are defined by the first or last occurrences of index taxa. Sugiyama’s zonation is the only one that spans practically the whole Triassic Period with relatively high resolution. It is based on occurrences of 233 taxa in 21 continuous sections of siliceous rocks from central Japan, and the proposed zones are recognized globally in low latitudes. In comparison with other low-latitude radiolarian zonations, Sugiyama’s zonation bears the longest and most continuous record of radiolarian events. On the other hand, because the investigated successions consist of siliceous claystone and bedded chert, radiolarian preservation is moderate, diversity is relatively low and other age-diagnostic fossils are rare. The zones TR 1 (Spathian), TR 2B (mid Anisian) and TR 4B (upper Ladinian –lower Carnian) are directly calibrated with co-occurring

TRIASSIC RADIOLARIAN BIOSTRATIGRAPHY AUTHOR(S)/YEAR Nakaseko & Nishimura 1979 Pessagno et al. 1979

Yao et al. 1980a, b, 1982; Yao 1982, 1990

TIME INTERVAL late Anisian to Norian latest Carnian to early late Norian Ladinian to Rhaetian

REGION

3 assemblages: Tripocyclia cf. acythus, Emiluvia? cochleata, Capnuchosphaera theloides 2 Oppel zones: Capnodoce Zone, Pantanellium silberlingi Zone

central Japan

3 assemblage zones: Triassocampe deweveri , Triassocampe nova , Canoptum triassicum stratigraphic distribution of 44 taxa

Ladinian to Norian

Kishida & Sugano 1982; Kishida & Hisada 1986 Nishizono & Murata 1983

late Anisian to Rhaetian

Greece, Sicily, Turkey SW and central Japan

Ladinian to Norian

Kyushu, Japan

Igo & Nishimura 1984 Blome 1984

Norian to Rhaetian latest Carnian to Late Norian

Kishida & Hisada 1985

NorianRhaetian?

central Japan Oregon, Baja California, Queen Charlotte Islands central Japan

Sato et al. 1986

late Anisian to Norian

Kyushu, Japan

Yoshida 1986

Carnian to Rhaetian

central Japan

Yeh 1989

Ladinian to ?Rhaetian

Oregon

Cheng 1989

Anisian, late Carnian? to Norian Anisian to Rhaetian Rhaetian

Philippines

Spathian to Rhaetian Spathian to Rhaetian

Far East Russia Philippines

late Norian Rhaetian

NE China

Carter 1990, 1993

Bragin 1991 Tumand a 1991

Yang & Mizutani 1991

Fig. 1. (Continued)

ZONES AND TYPE OF ZONES

SW and central Japan Baja California

De Wever 1982

Yeh 1990

165

Philippines Queen Charlotte Islands

4 assemblage zones: Eptingium manfredi, Emiluvia? clochleata, Capnodoce anapetes, Spongosaturnalis multidentatus 3 assemblages: Archaeospongoprunum compactum, Emiluvia? cochleata, Capnodoce anapetes 2 assemblage zones: Capnodoce anapetes Capnodoce sarisa, Canoptum triassicum 2 zones: Capnodoce Zone (Oppel Zone divided in 3 interval subzones), Betraccium Zone (Oppel zone divided in one Oppel and one interval subzone) (Fig. 2)

Palaeosaturnalis multidentatus Assemblage divided in Canoptum aff. triassicum and Canoptum lubricum subassemblages 5 range and concurrent range zones: Archaeospongoprunum compactum, Emiluvia? cochleata, Capnuchosphaera triassica, Capnodoce, Betraccium deweveri 6 interval zones: Capnuchosphaera, Capnodoce, Acanthocircus-Pseudoheliodiscus, Betraccium deweveri , Livarella-Canoptum, Justium cf. novum 4 assemblages: Pseudostylosphaera magnispinosa, Poulpus carnicus, Corum parvum, Orbiculiforma sp. A 3 assemblages: Pseudostylosphaera japonica, Pseudoheliodiscus sp. F, Betraccium deweveri 3 assemblages: Busuanga chengi, Trialatus megacornutus, Livarella sp. A 6 successive assemblages (defined by Unitary Associations) grouped in 2 zones: Proparvicingula moniliformis Zone, Globolaxtorum tozeri Zone (Fig. 2) 7 interval zones, upper 3 zones divided in subzones (Fig. 2) 8 interval zones: Pactarentinia koikei , Hozmadia altipedaria, Pseudostylosphaera japonica, Triassocampe deweveri , Muelleritortis cochleata, Capnuchosphaera, Capnodoce, Livarella Livarella - Canoptum rhaeticum Assemblage

166 AUTHOR(S)/YEAR Sugiyama 1992

Yeh 1992 Sashida et al. 1993

L. O’DOGHERTY ET AL. TIME INTERVAL Spathian to middle Anisian Norian to Rhaetian late Anisian to Norian

REGION central Japan Philippines central Japan

Kozur & Mostler 1994

Anisian to Rhaetian

Kozur 1995; Kozur & Mostler 1996; Kozur et al. 1996

late Anisian to Ladinian

Yeh & Cheng 1996 Nishizono 1996

Rhaetian early Anisian to late Norian

Philippines Kyushu, Japan

Sugiyama 1997

Spathian to Rhaetian

central Japan

Tekin 1999

Ladinian to Rhaetian Anisian to Ladinian

Turkey

Feng et al. 2000, 2001 Bragin 2000

late Olenekian to late Rhaetian

Bragin 2007 Moix et al. 2007

Norian early Tuvalian

Italy, Hungary, Austria Austria, Bosnia

south and southwest China Far East Russia

Cyprus Turkey

ZONES AND TYPE OF ZONES 3 assemblage zones: Parentactinia nakatsugawaensis , Hozmadia gifuensis, Triassocampe coronata 2 assemblages: Betraccium deweveri , Livarella longus 5 assemblages: Pseudostylosphaera japonica, Pseudostylosphaera helicata, Cryptostephanidium sp., Capnuchosphaera sp., Betraccium sp. 9 zones based on FADs or LADs of index and several other species (Fig. 2) subdivision of zones established by Kozur & Mostler 1994 (Fig. 2; fu ll names of subzones: Tiborella florida, Yeharaia annulata , Oertlispongus primitivus, Oertlispongus inaequispinosus, Ladinocampe annuloperforata , Ladinocampe vicentinensis, Pterospongus priscus, Spongoserrula rarauana, Spongoserrula fluegeli) Parabipedis pessagnoi Assemblage one unnamed zone and 5 interval zones based on FADs of index taxa: Pseudostylosphaera compacta, Plafkerium cochleatum, Capnuchosphaera triassica, Capnodoce, Betraccium deweveri 18 zones: 2 assemblage zones (TR 0 and TR 1) and 16 zones defined by FADs or LADs of index species (Fig. 2) stratigraphic di stribution of 332 taxa 4 new interval zones in the Anisian: Triassocampe dumitricai , T. coronata inflata , T. coronata coronata, T. deweveri Pseudostylosphaera fragilis Beds (upper Olenekian) and 14 zones: Hozmadia gifuensis (lower Anisian), Triassocampe diordinis (middle Anisian), Triassocampe deweveri (upper Anisian), Triassocampe scalaris (uppermost Anisian-lowermost Ladinian), Oertlispongus inaequispinosus (lower part of lower Ladinian), Falcispongus falciformis (upper part of lower Ladinian-lower part of upper Ladinian), Muelleritortis cochleata (middle-upper part of upper Ladinian), Tritortis kretaensis kretaensis (lower Carnian), Capnuchosphaera theloides (upper Carnian), Capnodoce crystallina (lowermiddle Norian), Lysemelas olbia (lower part of upper Norian), Betraccium deweveri (upper part of upper Norian), Livarella densiporata (lower Rhaetian), and Globolaxtorum tozeri (upper Rhaetian) stratigraphic distribution of 101 species Spongotortilispinus moixi Zone, taxon range zone

Fig. 1. (Continued) Development of Triassic radiolarian biochronology from the first publications in 1979 to the present. Note that in this table the Anisian and Ladinian are separated at the ‘historical’ boundary, that is, at the base of Reitziites reitzi Ammonoid Zone.

TRIASSIC RADIOLARIAN BIOSTRATIGRAPHY

167

RADIOLARIAN ZONES AND SUBZONES

Norian Capnodoce

Upper Triassic

Betraccium

Rhaetian Betraccium deweveri

Carter 1993 Globolaxtorum tozeri

Ass. 3

Proparvicingula moniliformis

Ass. 2 Ass. 1

Europe

Japan

Far East Russia

Kozur & Mostler 1994,1996; Kozur et al. 1996

Sugiyama 1997

Bragin 1991

TR 8D Haeckelicyrtium breviora Livarella densiporata TR 8C Skirt F TR 8B Praemesosaturnalis pseudokahleri

Betraccium deweveri

Livarella gifuensis

Betraccium deweveri

TR 8A Praemesosaturnalis multidentatus group

Pantanellium silberlingi Latium paucum

TR 7 Lysemelas olbia

Xipha striata

TR 6B Trialatus robustusLysemelas olbia

Capnodoce ruesti

TR 6A Capnodoce-Trialatus Justium novum Nakasekoellus inkensis

Triassocampe nova

Blome 1984

Canoptum triassicum

North America

CHRONOSTRATIGRAPHIC UNITS

Capnodoce antiqua

Capnuchosphaera lea TR 5B Poulpus carcharus

Tetraporobrachia haeckeli

Carnian

TR 5A Capnuchosphaera Tritortis kretaensis

Middle Triassic

Ladinian

Anisian-Ladinian boundary

Ladinocampe multiperforata

S. fluegeli S. rarauana P. priscus

TR 4A Muelleritortis cochleata

L. vicentin.

TR 3B Yeharaia elegans group

L. annuloperf.

Spongosilicarmiger O. inaequisp. italicus O. primitivus Spongosilicarmiger Y. annulata transitus T. florida

Anisian

Tetraspinocyrtis laevis

Lower Triassic Scythian

Parasepsagon robustus

Olenekian

Sarla dispiralis

TR 4B Spongoserrula dehli Muelleritortis cochleata

Plafkerium cochleatum

Yeharaia elegans

TR 3A Spine A2 Triassocampe deweveri TR 2C Triassocampe deweveri TR 2B Triassocampe coronata group TR 2A Eptingium nakasekoi group TR 1 Parentactinia nakatsugawaensis Ass. TR 0 Follicucculus-Parentactinia Ass.

Triassocampe diordinis

Hozmadia

"Stylosphaera" fragilis

Induan

Fig. 2. Correlation of Triassic radiolarian zones and subzones (for full names of European subzones see Fig. 1).

conodonts, but all the other zones are calibrated through correlation with zonal schemes established in other regions (Sugiyama 1997). In Europe (including Turkey), the first range chart for Triassic radiolarians was produced by De Wever (1982), who compiled the ranges of 44 Ladinian to Norian species based on previous works and his data from Greece, Sicily and Turkey (De Wever et al. 1979; De Wever 1982), but he did not define new zones. Although extensive systematic studies were carried out in Europe (Dumitrica 1978a, b, 1982a, b, c; Dumitrica et al. 1980; Kozur & Mostler 1972, 1978, 1979, 1981, 1983; Lahm 1984; Goricˇan & Buser 1990), no formal radiolarian zones were proposed. The reason for this delay was the lack of continuous radiolarian-bearing successions suitable for biochronological studies. Radiolarian occurrences in the western Mediterranean are limited in time and restricted to short pelagic intervals within relatively shallow-water deposits. However, these radiolarian assemblages come mostly from siliceous limestones and are thus generally well preserved and diverse

with ammonoids and conodonts providing age control. The first radiolarian zones were formally introduced by Kozur & Mostler (1994), who constructed a zonation on the basis of several sections from Hungary, Italy and Austria. They established nine zones for the Anisian to Rhaetian (Fig. 2). At that time, the lower Ladinian assemblages were by far the best studied and were divided into two zones (Spongosilicarmiger italicus and Ladinocampe multiperforata), which were subdivided into three and two subzones, respectively. Kozur (1995) and Kozur et al. (1996) elevated the rank of the Spongosilicarmiger transitus Subzone (i.e. lower subzone of the Spongosilicarmiger italicus Zone) to zonal level, defined two subzones within this zone and lowered the age to late Anisian. Kozur & Mostler (1996) further divided the upper Ladinian Muelleritortis cochleata Zone into three subzones. In the late 1990s important biostratigraphic work was carried out in Turkey by Tekin (1999), who did not introduce new zones but presented a range chart with 332 Ladinian to Rhaetian species. Kozur

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(2003a, b) presented an up-to-date correlation of radiolarian zones with ammonoid and conodont zones. He selected his own radiolarian zones for the Anisian, Ladinian and Carnian but used Japanese and North American zones for the Lower Triassic, Norian and Rhaetian. However, some radiolarian zones indicated in this publication have never been defined or described (e.g. Stigmosphaerostylus turkensis Zone in the upper Induan, Muelleritortis firma Zone in the Ladinian, Tritortis kretaensis dispiralis Subzone in the uppermost Ladinian, Squinabolella? trispinosa and Laxtorum perfectum zones in the Rhaetian). Therefore, in the present paper we reproduce only the zones proposed by Kozur in his previous publications. Overall, the Middle Triassic and Carnian radiolarian zonation of Kozur and co-workers is very precise (Fig. 2), but new findings of well preserved radiolarian faunas still allow for further refinement. For example, based on excellently preserved material from SE Turkey a new Spongotortilispinus moixi Zone has recently been inserted in the lower Tuvalian (Moix et al. 2007) and a zonal gap in the upper Ladinian–lowest Norian is expected to be filled based on the study of the radiolarian fauna from the chert member of the Zula Formation in Oman (Blechschmidt et al. 2004; Dumitrica & Hungerbu¨hler 2007), which encompasses the upper Anisian to lowest Norian interval. Some other radiolarian zonations have been proposed outside the aforementioned regions. For Far East Russia (Sikhote– Alyn, Koryak Upland, Sakhalin), Bragin (1991) proposed a zonation that covers almost the entire Triassic. The Spathian to Rhaetian interval is divided into seven zones, with the upper three zones subdivided into six subzones (Fig. 2). These zones are well dated with co-occurring conodonts, but in comparison with other zonations, include only a small number of radiolarian taxa, that is, the stratigraphic distribution of only 25 species is included in the range chart. For this reason, the zonation has rarely been used outside Russia. The zonation was later emended (Bragin 2000) to include 15 zones, with 9 zones newly defined (Fig. 1). Recently, Bragin (2007) presented the stratigraphic distribution of 110 Norian species from southern Cyprus and correlated the assemblages to his zones established in Far East Russia. In the Philippines, Yeh & Cheng (Cheng 1989; Yeh 1990, 1992; Yeh & Cheng 1996) described several Middle and Upper Triassic assemblages that contain typical Tethyan faunas. These assemblages were extracted from bedded cherts, and, lacking independent age control, their ages could be determined based solely on radiolarians whose range had previously been established in other regions (for correlation see Sugiyama 1997). Another zonation for the Philippines was proposed

by Tumanda (1991), who divided the Spathian to Rhaetian into eight interval zones. Since the early 1990s, extensive research on Triassic radiolarians has been carried out in Thailand (e.g. Sashida & Igo 1992; Sashida et al. 1997, 2000a, b; Kamata et al. 2002; Feng et al. 2005), China (e.g. Yang & Mizutani 1991; Feng 1992; Feng et al. 2000, 2001; Xia & Zhang 2000; Yao & Kuwahara 2000; Wang et al. 2002, 2005; Feng & Liang 2003) and Indonesia (Sashida et al. 1999). Because the existing radiolarian range charts apply well in these areas, new local zonations have only exceptionally been proposed (e.g. Yang & Mizutani 1991; Feng et al. 2000, 2001). Very few radiolarian localities have been studied in high latitudes, and most are concentrated in New Zealand where Early, Middle and Late Triassic radiolarians have been documented (Grapes et al. 1990; Aita & Bragin 1999; Takemura et al. 2002, 2003; Hori et al. 2003; Kamata et al. 2003; Takemura & Aono 2007; Takemura et al. 2007b; Kamata 2007; Kamata et al. 2007). This country is especially well known for Permian–Triassic boundary faunas at Arrow Rocks (Takemura & Aono 2007; Takemura et al. 2007a, b; Kamata 2007; Kamata et al. 2007). Middle Triassic high-latitude radiolarians have also been reported from the Omolon Massif in NE Siberia (Aita & Bragin 1999), and higher paleolatitudes were assumed for Late Triassic radiolarians from the Brooks Range in northern Alaska (Blome 1987). High-latitude faunas clearly differ from their low-latitude counterparts but, because data are very scarce, a separate high-latitude zonation has not been developed yet. De Wever et al. (2001) combined zonations from different regions in order to obtain a complete succession of low-latitude radiolarian zones for the entire Triassic Period. They selected the zones of Kozur & Mostler (1994, 1996) and completed these zones with those of Sugiyama (1997) for the Olenekian to middle Anisian, and Blome (1984) and Carter (1993) for the Late Triassic. Herein, we correlate the currently used zonations for their entire extent (Fig. 2).

The genera revision project The Mesozoic Working-Group of the International Association of Radiolarian Paleontologists (InterRad) has just completed a detailed taxonomic revision of Mesozoic radiolarians at the generic level (O’Dogherty et al. 2009a, b). The aim of this project, which began in 2006, was to compile and review the taxonomy of all existing genera as a basis for a refined Mesozoic radiolarian stratigraphy. The ‘Mesozoic Generic revision project’ was comprised of 11 scientists from 8 countries. This working group met twice in May 2006 and

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May 2007 in two-week duration meetings that provided a forum for taxonomic discussions necessary to the making of a refined catalogue of genera. The basic purpose was to provide the scientific community with a catalogue of type-species in hopes of clarifying the correct generic assignment of Mesozoic species (more than 6000 species have been described for the Mesozoic alone). Nonetheless, taxonomy without stratigraphy has little significance and for this reason the generic catalogue also contains precise information on stratigraphic ranges. Generic ages are referred not to radiolarian zones or subzones, but to ICS stage subdivisions that can be correlated easily to faunal stages, a subdivision classically used by most of the researchers working on Triassic radiolarians (Figs 2 & 3). The Working-Group review has shown that more than 900 genera have been published since the early work of Ehrenberg (1838), in which the genus Cornutella (a common Jurassic to Recent genus) was described. Many of these nominal genera are regarded as valid names according to the rules of the International Code of Zoological Nomenclature (ICZN), but a considerable number have to be treated as nomina dubia, synonyms or homonyms. The project consists of two distinct parts: the Triassic, and the Jurassic –Cretaceous. This division is justified by the low number of genera common to each part (only 30 genera cross the Rhaetian– Hettangian boundary: Fig. 4). The Triassic part contains 381 described genera of which only 26% must be treated as invalid taxa according to ICZN rules. The compilation of species and genera has shown a very low number of nomina dubia (4%), which underscores the existence of the few publications dealing with Triassic radiolarians before the 1970s. This means that nearly all taxonomic publications since that time have held to a precise systematic concept (especially at the generic and specific level) and, in general, have good illustrations of the type-species. The Working-Group has revised and updated the stratigraphic distribution of 282 valid Triassic genera providing an accurate biostratigraphic chart for the Triassic that is reproduced in Figure 3. This range chart allows for dating samples at substage level by using the identification of characteristic taxa at the generic level. The reason this is so powerful is that the biostratigraphic ranges of many taxa are relatively short, that is, nearly 75% of the genera show a duration of less than four Triassic substages. This is, in part, the imprint of a homogeneous systematics, inherited by the low number of workers involved in both taxonomy and stratigraphy of the Triassic Period. A rapid analysis of authorship indicates that 75% of the generic systematics has been produced by only four authors: P. Dumitrica, H. Kozur, E. Pessagno and U. Tekin. In summary,

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we can say that the accurate taxonomy used in the systematics of Triassic genera, together with the rapid evolution of radiolarians during the Middle Triassic (De Wever et al. 2003, 2006), are the main reasons for the good biostratigraphic resolution, even at the generic level.

Triassic range chart of genera The discussion that follows is sufficient to outline our present understanding of generic occurrences of radiolarians through Triassic time, highlighting major evolutionary events in the history of the group. The genera discussed below were selected because they are sufficiently distinctive to be recognizable, and have been employed at a number of localities and in zonations, by more than one author. This information is based on the range chart (reproduced in Fig. 3) resulting from the new taxonomic revision of Triassic genera made by the InterRad Mesozoic working group discussed above (O’Dogherty et al. 2009a, b). For more comprehensive information on radiolarian systematics and terminology the reader is referred to the latest monograph on this group of microfossils (De Wever et al. 2001). Basic information for a less demanding audience is also available online (http://www.radiolaria.org).

Induan The beginning of the Triassic is marked by the relatively common presence of Permian survivors (see below) and the occurrence of the first and primitive forms of monocyrtids (Tripedocorbis) and dicyrtids (Triassospongocyrtis).

Olenekian Radiolarian diversity starts to increase measurably during the Late Olenekian with Entactinaria diversifying more rapidly than other groups. This interval records the first occurrence of primitive Eptingiidae such as Spongostephanidium (?Pentabelus) simultaneous with the initial occurrence of most simple Multiarcusellidae (Tiborella). The first Triassic spicular Entactinaria (Archaeosemantis and Parentactinia) occur through this interval. Among nassellarians, this period is characterized by the occurrence of the first primitive multicyrtids, characterized by bearing a low number of segments (Anisicyrtis), and the earliest representatives of the monocyrtid Poulpidae (Hozmadia). Also, it records the initial development of new spicular forms (Verticiplagia). The Olenekian–Anisian boundary is marked by the final disappearance of typical Paleozoic families (Follicucullidae, Latentifistulidae).

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Fig. 3. Range chart of Triassic radiolarian genera including family position and re-illustration of the type-species presented in chronological order. This table represents the main goal of the generic revision project undertaken by the Mesozoic Working Group of InterRad (O’Dogherty et al. 2009a). The FAD and LAD numbers refer to stage and

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Fig. 3. (Continued) substage numbering used in O’Dogherty et al. (2009a). Radiolarian orders are indicated by letters; Albaillellaria (A); Latentifistularia (L); Entactinaria (E); Spumellaria (S); Nassellaria (N). The triangles indicate that the genus either ranges down to the Permian, or upward to the Jurassic.

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Early Anisian Originations during the Early Anisian are not as well documented as in the Early Triassic, but it is likely that the first appearance of advanced forms of Eptingiidae (Eptingium and Cryptostephanidium) took place at this time, as well as the origin of two of the most diverse Triassic families of multicyrtid nassellarians: the Planispinocyrtiidae (Planispinocyrtis) and Ruesticyrtiidae (Triassocampe).

Middle Anisian The Middle Anisian (Pelsonian and Illyrian) is marked by a huge diversification, especially among nassellarians. It records the first occurrence of many mono- and dicyrtid genera of the families Tripedurnulidae (Tripedocassis, Baratuna and Tripedurnula), Poulpidae (Triassobipedis, Eonapora and Neopylentonema), and the origination of the families Nabolellidae (Fueloepicyrtis), Ultranaporidae (Muellericyrtium, Hinedorcus and Silicarmiger) and Spongosilicarmigeridae (Spongosilicarmiger and Nofrema). The multicyrtids also experienced a large diversification, and new families such as the Tetraspinocyrtidae (Tetraspinocyrtis), Monicastericidae (Monicasterix) and Bulbocyrtiidae

appeared for the first time. Several new generic occurrences include the Ruesticyrtidae (Paratriassocampe, Yeharaia, Pararuesticyrtium and Annulotriassocampe) and Planispinocyrtiidae (Spinotriassocampe and Triassocyrtium). The Middle Anisian was also a period of high diversification in Entactinaria. New families appearing at this time are the Pentactinocarpidae (Pentactinocapsa and Pentactinorbis) and Heptacladidae (Pseudosepsagon, Heptacladus, Parentactinosphaera, and other undescribed genera) together with the mass-occurrence of the subfamily Hindeosphaerinae (Parasepsagon, Pseudostylosphaera, Sepsagon and Hindeosphaera). Spumellarians start to be frequent in the middle Anisian, and while important occurrences are rare, an interesting modification is the initial development of twisted spines in a simple spongy form (Monospongella) belonging to the Sponguracea, a superfamily in which this characteristic is extremely common. The Pyloniacea, one of the most important spumellarian groups in both Mesozoic and Cenozoic, appeared during this period; the oldest representative is the genus Patrulius. A word of caution is needed, however, when interpreting the outstanding diversification during the Middle Anisian as this could be an artifact of

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Fig. 4. Faunal spectra at the Triassic– Jurassic boundary. The FAD and LAD numbers refer to stage and substage numbering used in O’Dogherty et al. (2009a, b). Same letter key as Figure 3.

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Fig. 4. Continued.

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preservation. Extremely well-preserved Triassic faunas first appear in this interval, for example, the Pelsonian assemblage from Cristian in Romania (Dumitrica 1982c, 1991 etc.) and the Illyrian assemblage from Felso¨o¨rs in Hungary (Kozur & Mostler 1994). The first occurrences of rare and delicate forms (e.g. Fueloepicyrtis, Tripedocassis, Tripedurnula, Pentactinocapsa, Patrulius) may also be related to better preservation. The number of resistant genera that newly appear is, nevertheless, extremely high, which indicates that the diversification rate in the middle Anisian was indeed the highest during the Triassic.

Ladinian The occurrence of heavily ornamented (foliaceous) Oertlispongidae (Pterospongus, Steigerispongus, Scutispongus and Spongoserrula) is one of the most important events during the Ladinian. The multicyrtid families Pseudodictyomitridae (Triassocingula and Corum) and Canoptidae (Canoptum and Multimonilis) make their first occurrence also. This interval further records an important bloom of Entactinaria bearing strong, twisted spines (Tritortis and Muelleritortis) and advanced Austrisaturnalinae (Ornatisaturnalis, Hungarosaturnalis, Praeheliostaurus).

Late Anisian The early late Anisian (starting with the Reitziites reitzi Ammonoid Zone) is easily recognized by the occurrence and diversification of the genus Oertlispongus; nassellarians reach their maximum diversification during this period. The late Anisian is characterized by the occurrence of the typical scalariform multicyrtid genus Ladinocampe and the first nassellarians bearing a skirt-type distal segment (Deflandrecyrtiidae: Deflandrecyrtium and Dreyericyrtium). The typical dicyrtid Foremanellina and the scalariform Annulobulbocyrtium, genera characterized by a very large and rounded cephalis, also appear at this time.

Early Carnian Major changes in the composition of Triassic radiolarian assemblages occurred during the early Carnian (note that a twofold subdivision is used herein for the Carnian, because Cordevolian is considered to be part of the Julian) and a huge turnover, both at the generic and family level (De Wever et al. 2006), took place. Generic diversity reaches its maximum extent, especially among spumellarians, and with it came the first blooming of spumellarians bearing twisted spines (Sarla, Spongotortilispinus), hollow spines (Capnuchosphaera), branched spines (Kahlerosphaera) and the origin of a particularly

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common group during the Mesozoic, the true Saturnalidae. The latter are represented by three early branches: the Archaeoacanthocircidae (Huglusphaera), Heliosaturnalinae (Pseudoheliodiscus) and Italosaturnalinae (Annulosaturnalis). However, five genera of the Oertlispongidae went dramatically extinct within this interval. This is remarkable in that the Oertlispongidae are considered the stock of all Triassic Saturnalidae (Kozur & Mostler 1990; Dumitrica & Hungerbu¨hler 2007). The early Carnian also witnessed new morphological organizations within the arm-bearing Pyloniacea, with the appearance of the first genus having four arms, for example, Triassocrucella, which can be considered the oldest Hagiastridae. Hexaporobrachia, an easily recognized form with six latticed tubular arms also occurs in this interval. In nassellarians, the genus Praeprotunuma first appears, as well as the genus Trialatus, the last representative of the Tetraspinocyrtiidae. In the Entactinaria, the most interesting innovation is the occurrence of the Spongosaturnaloididae (Spongosaturnaloides and Ploechingerella), a group very close to the family Eptingiidae that possess a simple or multiple saturnalid type ring. At this time several genera of the Multiarcusellidae (Austrisaturnalis, Quadrisaturnalis and Hungarosaturnalis) develop a saturnalid-like test with the ring resulting from the junction of cortical arches.

nassellarians (Triassocampe, Pararuesticyrtium, Papiliocampe) and all representatives of the Xiphothecaellidae. Caphtorocyrtium is characteristic of the early Norian and is the only known Planispinocyrtiidae to possess a distal skirt-like chamber (quite common in Ruesticyrtiidae) and probably the last multicyrtid genus displaying this morphology, which began in the middle Anisian. Among spumellarians, it is interesting to note the occurrence of the genus Triarcella, which develops a saturnalid-like ring by connecting the three carinated main spines.

Middle Norian This interval records a drastic diversity drop. It is marked by severe extinctions affecting many primitive multicyrtids (Ruesticyrtiidae, Bulbocyrtiidae), and the last representatives of the Capnodocinae and Capnuchosphaeridae families went extinct at this time. Only three new genera appeared, the incertae sedis multicyrtid Lysemelas, the new Pentactinocarpidae with cortical shell Braginella and the new Saturnalinae Mesosaturnalis.

Late Norian

The Pantanelliidae first appeared during this interval. The genus Pantanellium is the oldest representative and represents the main stock from which all the other Mesozoic pantanelliids are derived. The interval also records the first appearance of pantanelliids bearing strongly twisted-spines (Betraccium), but these occur only in low latitudes. An interesting form in this group is the genus Capnodoce, the first pantanelliid with hollow spines. The Capnuchosphaerinae become very diverse with Dicapnuchosphaera, Monocapnuchosphaera, and Nodocapnuchosphaera appearing at the end of the Carnian. At the early –late Carnian transition the genus Archaeoacanthocircus appeared, an easily recognizable primitive saturnalid with a flat and very broad ring bearing four spines. The genus Xipha (¼ Nakasekoellus) and Mostlericyrtium are two easily recognizable nassellarians appearing at this time also. Only one Entactinaria, the genus Xenorum, originated within this interval; it is the first robust Eptingiidae bearing twisted spines.

New multicyrtid nassellarians displaying typical Jurassic patterns are common in the late Norian. These are represented by two genera of the family Canoptidae (Canoptum and Laxtorum). The latter first appeared in the late Norian, whereas the former arose in the Ladinian but did not become abundant and diverse until the late Norian; both genera survived the end-Triassic extinction. Following the Carnian, no new families appeared for the remainder of the Triassic with the exception of the Livarellidae, which produced several dicyrtid representatives (Ayrtonius, Citriduma and Livarella). Aside from these, dicyrtids are rarely recorded in the Late Triassic. Another interesting form is the genus Bipedis, belonging to the common Jurassic Ultranaporidae. This interval saw the rapid diversification and acme of pantanelliids bearing strongly twisted-spines (Betraccium) and of saturnalids with a 3- or 4-bladed ring, for example, the genus Octosaturnalis. In general, the late Norian is marked mainly by the appearance of new forms of nassellarians and spumellarians while the Entactinaria are represented only by the appearance of Ferresium. This will be a common feature during the Rhaetian where the assemblages are almost entirely dominated by nassellarians and spumellarians.

Early Norian

Early Rhaetian

The Early Norian is marked by severe extinctions affecting many primitive (scalariform) multicyrtid

The base of the Rhaetian is marked by the appearance of new representatives of the Canoptidae

Late Carnian

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(Globolaxtorum, Neocanoptum and Proparvicingula). The latter genus seems to be the direct ancestor of the Jurassic family of Parvicingulidae because it displays an offset arrangement of pore frames quite close to the family. In spumellarians, a remarkable evolutionary adaptation is noted whereby the genus Serilla evolves from Ferresium by acquiring concave sides and a strongly triangular test (Carter & Guex 1999). Orbiculiformella, a common genus in the Jurassic, also appears at this level.

Late Rhaetian The end of the Triassic is marked by the extinction of many families (see below). Only two spumellarian genera originate during this interval (Pseudacanthocircus and Tipperella) and both cross the T/J boundary.

Radiolarians, mass-extinctions and problems at Triassic boundaries Problems exist with radiolarian faunas at both the base and top of the Triassic. Takemura et al. (2007a, b) have only recently discriminated endPermian from earliest Triassic radiolarian assemblages, whereas at the top of the Triassic System, few radiolarian zonations cross the Triassic–Jurassic boundary (e.g. Yao et al. 1980a, b; Yao 1982; Yoshida 1986; Sugiyama 1997). Since both boundaries are characterized by severe extinctions related to major ecological perturbations, a thorough evaluation of boundary events is far beyond the scope of this paper. This problem has partly been addressed in recent studies dealing with radiolarian taxonomy as well as short-term changes in diversity and relative abundances (De Wever et al. 2006; Carter 2007). The changes in taxonomic structure have been compared to fluctuations in geochemical composition of sediments and to faunal turnovers in other fossil groups (see De Wever et al. 2006 for a recent review).

The end of Palaeozoic groups and the collapse of radiolarian diversity The Permian –Triassic boundary records the greatest extinction known in the fossil record. Radiolarians were severely damaged at all taxonomic levels, and undoubtedly the extinction must have played an important role in the evolutionary history of this group of protists during the Triassic. According to the data reviewed by De Wever et al. (2001, 2003, 2006), no new orders of radiolarians appeared after the Permian –Triassic crisis. The only exception might be the Nassellaria, if we consider that Mesozoic nassellarians are not direct

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descendants of Paleozoic representatives, all of which disappeared by the end of the Carboniferous. At the family level, the crisis seems to have produced little effect because of the 18 families that existed during the Permian (De Wever et al. 2006), 13 cross the critical interval. This is rather different at the generic level, where the end-Permian extinction caused a decrease of more than 50% (Kuwahara & Yao 2001; Umeda 2002; Yao & Kuwahara 1997), although real diversity varies from one author to another. Nevertheless, the decline in generic diversity unquestionably starts in the late Lopingian, but the scope of this faunal turnover is still precarious for two main reasons: (1) the scarcity of papers covering the Permian– Triassic interval (see De Wever et al. 2006); and (2) the lack of a taxonomic revision of Late Paleozoic genera. The scarcity of data is tightly bound to the abrupt disappearance of radiolarian cherts from deepmarine sections of South China, Japan, and Western Canada in the Late Permian (Isozaki 1997). Radiolarians reappeared only in the Spathian following a gap of 7– 8 Ma, but the mechanisms responsible for the cessation of siliceous sedimentation are not entirely understood. Logically, the abrupt change in radiolarians is amplified by this lithological turnover, and hence the current information on Early Triassic radiolarians should be considered quite incomplete. The development of a satisfactory common taxonomic system for the Late Permian–Early Triassic is also hampered by a ‘frontier effect’ that occurs equally at major boundaries of the Mesozoic. In other words, the taxonomic criteria used for classification change over the Eras, resulting in a somewhat artificial taxonomy. This evidently results because the groups of authors working on Late Palaeozoic and Early Triassic radiolarians are quite distinct in most cases. In this sense, De Wever et al. (2001) stressed the importance of homogenizing the taxonomic criteria used for the Palaeozoic, Mesozoic and Cenozoic, in order to update our biostratigraphic knowledge.

The long Early Triassic recovery True recovery following the end-Permian extinction does not begin until the middle Anisian when new genera bloom and assemblages become highly diverse. The Middle Triassic was undoubtedly the main epoch for radiolarian radiation. More than a third of the total number of radiolarian families recognized from the Cambrian to Present originated during this important period of plankton radiation (De Wever et al. 2006). Similarly, generic diversity starts to increase progressively in the early middle Anisian, and attains a maximum in the early Carnian with the occurrence of more than 135 genera (Fig. 3). However, recovery following the

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aftermath of the end-Permian extinction is marked by rare occurrences, poor preservation and a paucity of new forms (Sashida 1983, 1991; Sugiyama 1992, 1997; Suzuki et al. 2002; Hori et al. 2003; Kamata 2007; Kamata et al. 2007; Takemura et al. 2007a, b). These faunas mostly contain species of entactinarians (Entactinia, Hegleria, Parentactinia, Pseudostylosphaera, Tiborella, Glomeropyle), some sparse mono- and dicyrtid nassellarians (Archaeosemantis, Hozmadia, Tripedocorbis, Triassospongocyrtis), spicular nassellarians (Archaeosemantis, Verticiplagia) and rare spumellarians (Thaisphaera, Pegoxystris). Recently, some representatives of Middle Triassic spumellarians belonging to the families Intermediellidae (Tetrapaurinella, Paurinella), Gomberellidae (Tamonella) and Oertlispongidae (Pararchaeospongoprunum) have been reported from the uppermost Permian of China (Feng et al. 2006); however, they have not been found yet in Lower Triassic material. Kozur et al. (1996) and Kozur (1998, 2003a) have suggested that the long phase of recovery may be due to immigration from southern high latitude cold waters (less affected than those of high northern and low latitudes) rather than from tropical waters, which would be a faster process. This process would be slow (c. 5 Ma; late Induan to early Anisian) because it would require first the adaptation of some elements of the cold water fauna to warm tropical conditions. But, as stressed by De Wever et al. (2006), real data neither support nor contradict this interpretation, they are just insufficient at the present time.

The Anisian – Ladinian boundary The GSSP for the Ladinian Stage has only recently been established (Brack et al. 2005). The decision on the boundary ammonoid zone thus postdates all hitherto published radiolarian zonations for the Middle Triassic (Fig. 1). This GSSP was discussed for several decades, and two options were strongly debated. The question was whether to define the base of the Ladinian at the base of the Reitziites reitzi Ammonoid Zone or at the base of the Eoprotrachyceras curionii Ammonoid Zone. The FAD of Eoprotrachyceras curionii (Mojsisovics) was finally approved, and the GSSP was ratified by the IUGS Executive Committee in Spring 2005. Based on the radiolarian record, Kozur vigorously advocated that the best level for the base of the Ladinian would be the FAD of Reitziites reitzi (Bo¨ckh) (Kozur & Mostler 1994; Kozur 1995, 2003a, b). He pointed out that radiation of radiolarian taxa was explosive from the Paraceratites trinodosus to the Reitziites reitzii zones and consequently, that the resolution of radiolarian

biochronology in this interval is as high as that of the ammonites. In contrast to this, no distinct change in the radiolarian fauna is recorded near the base of the E. curionii Zone. Other radolarian researchers mainly followed Kozur’s arguments. They agreed that the base of the Spongosilicarmiger italicus Radiolarian Zone (corresponding to the Reiziites reitzi Ammonoid Zone) is well marked by a major radiolarian turnover. In addition, they favoured this boundary for practical reasons, that is, because the marker species are easy to find even in poorly preserved material. The lineage of Oertlispongidae, a family with a high biochronological value for the Ladinian, diversifies at this level, and the species below and above this boundary have dissolution-resistant spines, which are easy to determine and always common. Consensually, the Reitziites reitzi Ammonoid Zone was considered as the base of the Ladinian in radiolarian papers published until 2005 (e.g. Sashida et al. 1999; Dumitrica 1999; Feng & Liang 2003; Goricˇan et al. 2005). The officially accepted GSSP at the base of the Eoprotrachyceras curionii Ammonoid Zone lies much higher than the base of the R. reitzi Zone and is of crucial importance to Middle Triassic radiolarian dating. First, one has to bear in mind that early Ladinian ages based on radiolarians in literature older than 2005 should be ‘translated’ to late Anisian. Second, the boundary should be displaced in existing radiolarian zonations. In Figure 1, where the historical review is presented, we retain the early Ladinian in the old sense of this term. In Figure 2 the new position of the Anisian –Ladinian boundary is emphasized. In the range chart of genera (Fig. 3) developed herein, we wanted to preserve the obvious faunistic changes at the base of the R. reitzi Zone. For this reason, the Anisian in the range chart consists of three intervals, but the Ladinian is not subdivided. The upper Anisian contains two radiolarian zones formerly attributed to the lower Ladinian; these are the Spongosilicarmiger italicus and Ladinocampe multiperforata zones. The Ladinian in the range chart basically is equivalent to the Muelleritortis cochleata Zone. We note that the base of the Ladinian as now defined is not distinct in the radiolarian fauna. The Anisian –Ladinian boundary is correlated to the upper part of the Ladinocampe multiperforata Zone, that is, to the upper part of the Ladinocampe vicentinensis Subzone (Kozur 2003a, b). Because the radiolarian zones do not allow an exact correlation with ammonoid zones at this level and because the stratotype at Bagolino in the Southern Alps (Brack et al. 2005) contains no radiolarians, the only usable solution remains to consider the top of the Ladinocampe multiperforata Zone as the top of the Anisian.

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The Carnian – Norian boundary The Carnian –Norian boundary (CNB) is one of four Triassic boundaries that are not yet defined. In North America, this boundary traditionally has approximated the base of the Kerri Zone, that is, the level between the Macrolobatus and Kerri zones (Silberling & Tozer 1968), whereas in Tethys the boundary is located between the Spinosus and Jandianus zones (Krystyn et al. 2002). Currently there are no proposals for the Carnian –Norian boundary before the Triassic Subcommission for consideration, but two candidates and at least two levels are under discussion (M. J. Orchard, pers. comm. 2008): 1.

2.

The thick hemipelagic limestone succession at Pizzo Mondello, Sicily (Muttoni et al. 2001, 2004; Nicora et al. 2007) has neither a precise level nor a marker taxon proposed for definition and the magnetostratigraphic marker previously employed as a proxy apparently does not correspond with any significant fossil events. However, Nicora et al. (2007) consider the FAD of Epigondolella quadrata (¼ E. abneptis) as a possible marker. Most radiolarians illustrated from this section (Nicora et al. 2007) are clearly early Norian, with only a few taxa mentioned as Carnian, and depending on where the boundary is eventually placed, some of these could be early Norian also. Orchard (2007a, b) proposed a section Black Bear Ridge, in the Williston Lake area in northeast British Columbia, for the Carnian– Norian boundary based on the excellent fossil succession. This section contains a continuous outcrop of Ludington and Pardonet formations deposited in a deep marine slope setting on the north-western margin of Pangea and also contains ammonoids (Tozer 1965) and abundant bivalves. A significant bivalve turnover (McRoberts 2007), and a dip in carbon isotopes (Williford et al. 2007) are both coincident with the appearance of Metapolygnathus echinatus (occurring in both North America and Tethys) and with the ascendance of typical M. primitius; this lies close to the base of the traditional Kerri Zone and is favoured as the datum for GSSP definition in Canada. Radiolarians are also expected to play a role in Carnian– Norian boundary studies at Black Bear Ridge as they have previously in the Charlotte Islands (Carter & Orchard 2000). Presently, only one late Carnian collection is known from the Black Bear Ridge area (Carter, pers. comm.), but collections from many localities in the Charlotte Islands provide insight to the boundary interval, as they are linked through

3.

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M. echinatus, which occurs in both areas. A full multi-authored proposal for the section at Black Bear Ridge is expected in 2009. Since the late 1970s, Kozur (1980, 2003a, b) has advocated the FAD of Epigondolella quadrata as a marker for the Carnian –Norian boundary. This taxon is common in Tethys and in the upper part of the Kerri Zone and upper M. primitius conodont zone (Orchard 1983) in North America. These ideas are presently cloaked in taxonomic uncertainty since the introduction of several new and transitional taxa and have not lead to a proposal.

Given the uncertainty regarding future definition of the Carnian –Norian boundary, it is difficult at this time to accurately determine radiolarian ranges surrounding this particular level.

Triassic– Jurassic boundary faunas The Triassic –Jurassic boundary marks a significant palaeontological event in earth history as intense physical and chemical forces affected both terrestrial and marine faunas (Tanner et al. 2004). In the seas, conodonts and ceratite ammonoids disappeared almost completely with only a few known holdovers in the earliest Hettangian, bivalves gradually evolved across the boundary, pollen changed significantly (McElwain et al.1999) and radiolarians suffered a dramatic extinction that signalled a productivity collapse in the oceans (Ward et al. 2001). Prior to 1989, knowledge of Rhaetian radiolarians was known only from scant reports in Japan (Yao 1982; Yao et al. 1980a, b; Kishida & Sugano 1982; Kishida & Hisada 1986; Yoshida 1986), China (Kojima & Mizutani 1987), Austria (Kozur & Mostler 1981; Kozur 1984), Oregon (Yeh 1989) and New Zealand (Spo¨rli & Aita 1988; Spo¨rli et al. 1989) while several collections of probable Hettangian age were known from the Queen Charlotte Islands (Pessagno & Blome 1980; Pessagno & Whalen 1982), east-central Oregon (Pessagno & Blome 1980) and a few taxa were illustrated from Japan (Hori 1990). The impetus for detailed Triassic –Jurassic boundary studies began in 1987 with the finding of long successions of Rhaetian and Hettangian radiolarians dated by co-occurring conodonts and/or ammonites at Kennecott Point and Kunga Island, Queen Charlotte Islands (Carter et al. 1989; Carter 1990; Tipper & Carter 1990; Tipper et al. 1994). Soon after, many species were described and zonation based on unitary associations led to the establishment of two radiolarian zones for the Rhaetian (Carter 1993) and seven zones for the Hettangian– Sinemurian (Carter et al. 1998).

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Since that time, Rhaetian faunas have been found around the globe (see review in Carter 2007), and Hettangian faunas are known from the Northern Calcareous Alps (Kozur & Mostler 1990), Hungary (Kozur 1993; Pa´lfy & Doszta´ly 2000; Pa´lfy et al. 2007), Italy (Bertinelli et al. 2004), Japan (Hori 1992; Sugiyama 1997), Montenegro (Goricˇan 1994) and southern Turkey (Tekin 1999, 2002). However, earliest Hettangian faunas are still very rare and confirmed only to the Queen Charlotte Islands and Japan (Carter & Hori 2005; Longridge et al. 2007a, b), Montenegro (Goricˇan 1994) and Peru (Carter 1993). The following discussion of Triassic –Jurassic boundary faunas is based on collections from the Queen Charlotte Islands, as these are the most diverse and completely-documented fauna known and support postulation by Hallam & Wignall (1997) that the best evidence for catastrophic change at the end of the Triassic may come from microfossils.

Rhaetian radiolarian faunas and the end-Triassic extinction Over 160 described or informal species are present in the diverse Rhaetian fauna of the Queen Charlotte Islands (Carter 1993). Approximately half disappear towards the end of the Proparvicingula moniliformis Zone (early Rhaetian), but the others together with newly arising species range upward into the Globolaxtorum tozeri Zone (late Rhaetian) and many continue into uppermost beds of the Rhaetian. Then, over the space of less than a metre, nine families, at least 27 genera (Fig. 4), and nearly all Rhaetian species disappeared (Longridge et al. 2007a, b). For radiolarian correlation with ammonite and condont faunas see also Longridge et al. (2007 a, b). A lesser scale pattern of extinctions together with the first appearance of Hettangian taxa (discussed below) has also been observed in Japan, providing evidence that the radiolarian crisis at the end of the Triassic was not local, but rather global in extent (Carter & Hori 2005). The taxa most severely affected by the extinction were architecturally complex forms, for example, Eptingium, Icrioma, Nabolella, Paratriassoastrum, Pentactinocarpus, Praecitriduma, Tetraporobrachia, and short-ranging genera confined to the Rhaetian such as Serilla, Kungalaria and Globolaxtorum. It is note worthy that nearly all forms with highly twisted or spiralling spines disappeared. This morphological characteristic is exceptionally well displayed in Upper Triassic radiolarians, but is not found in Hettangian taxa whose peripheral spines are generally primitive and rod-like. Genera surviving the boundary are typically forms with

conservative morphologies such as Amuria, Archaeocenosphaera, Canoptum, Crucella, Orbiculiformella, Pantanellium and Paronaella. Guex (1993, 2001) has postulated that conditions of ecological stress may have contributed to simplification and/or reduced size in protists such as foraminiferans, silicoflagellates and radiolarians. The end-Triassic radiolarian fauna substantiates these ideas, as it is clear from perturbations in geochemical signatures (McRoberts et al. 2007; Pa´lfy et al. 2007) that ecological conditions were very unstable at that time. The latest Rhaetian fauna is rife with individuals that are sharply reduced in size and multicyrtid nassellarians having a reduced number of chambers (see also Carter & Guex 1999). This trend is further accentuated in the earliest Hettangian where the fauna is dominated by small-sized spumellarians and rare multicyrtid nassellarians with fewer chambers. This evidence supports O’Dogherty & Guex (2002), who suggested that spumellarians are more extinction-resistant than nassellarians and thus are more likely to dominate post-extinction faunas. Small spumellarians also dominate the recovery fauna following the endPermian extinction (e.g. Kakuwa 1997, English abstract: p. 76). See also Matsuoka (2007), who assessed feeding mechanisms in living multisegmented nassellarian radiolarians and determined that the diminishment of available prey may contribute to the rarity of these forms. Should these same feeding modes have existed in the Mesozoic, they too may have played an important role in contributing to the lack of nassellarians during times of ecological stress.

The early Hettangian radiolarian fauna The low diversity earliest Hettangian fauna is composed of primitive indeterminate spumellarians and entactiniids, rare nassellarians and usually only a single species of a new and/or surviving genus. A few Rhaetian holdovers are present in the very lowest beds but these are always small, rare and disappear quickly. The persistence of these shortranging holdovers may support ideas that the extinction was rapid, but not instantaneous. This trend is also seen in an exceptionally abundant and well preserved mixed fauna from Kennecott Point in the Queen Charlotte Islands (sample R2, section II in Longridge et al. 2007b) composed of c. 90% Hettangian species and ,10% Rhaetian holdovers (Longridge et al. 2007b). This sample occurs well above Rhaetian faunas and below lowest Hettangian ones that lack Rhaetian holdovers. Survival faunas have generally been described as producing low diversity, simple forms that are dominated by highly abundant geographically and environmentally widespread species and can also

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include blooms of opportunistic taxa that thrive only in difficult environmental conditions (Erwin 1998, 2001). The earliest Hettangian radiolarian fauna of the Queen Charlotte Islands is dominated by simple forms with spongy or irregularly latticed meshwork and long, rod-like rather than triradiate and/or twisted spines. Most spherical spumellarians and entactiniids lack definable organized structure, are extremely varied morphologically, and occur with floods of Archaeocenosphaera laseekensis, a spherical form with simple hexagonal pore-frames and no spines. Pantanellium tanuense, a probable opportunist, is also abundant in all but the very oldest samples. These two species suggest that oceanic productivity may have been restored sufficiently in the earliest Hettangian to allow vast numbers of these simple forms to proliferate. Nassellarians such as Canoptum merum, Droltus hecatensis, Bipedis elizabethae and indeterminate parahsuids are rare and very small. Over time the number of ‘indeterminate spherical spumellarians’ (possibly short-lived endemics) reduces as the fauna rebuilds. True recovery following the endTriassic extinction does not begin until the middle –upper Hettangian when new genera appear and assemblages become more diverse. Approximately 20 genera, most arising in the Carnian, Norian and Rhaetian, survived the end Triassic crisis, and at least six new genera first appear in the lower Hettangian together with other indeterminate forms. Most surviving genera are present in the Hettangian, but others, for example, Citriduma and Gorgansium, having sporadic distribution, do not reappear until the Pliensbachian; hence their absence in the Hettangian may not be that unusual. Genera barely surviving the extinction have been termed ‘Rhaetian holdovers’ (Longridge et al. 2007b) and include Fontinella, Deflandrecyrtium, Livarella and Proparvicingula. These genera are probably not true survivors because they are very rare, small in size, and all disappear completely by the middle Hettangian, at least in material from Queen Charlotte Islands. Other surviving genera are present in the Hettangian but disappear before the end of the Lower Jurassic. These include Haeckelicyrtium, Laxtorum, Parentactinia and Tipperella: the latter is abundant in the early Hettangian but gradually disappears as the fauna rebuilds; Parentactinia disappears at the end of the Hettangian, Laxtorum in the lower Pliensbachian; and Haeckelicyrtium is rare in the Hettangian and Sinemurian, common in the Pliensbachian, but unknown thereafter. Still other genera survived to have their major radiation in the Jurassic and Cretaceous, with the most well known being Canoptum, Crucella, Orbiculiformella, Pantanellium and Paronaella. Finally, the group least affected by

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events may be the saturnalids; Mesosaturnalis, Palaeosaturnalis, Praehexasaturnalis, Pseudacanthocircus and Pseudoheliodiscus all survived the end Triassic crisis, and a few species even crossed the boundary, for example, Pseudacanthocircus troegeri and Mesosaturnalis acuminatus.

Current status of the Triassic – Jurassic boundary The system boundary between the Triassic –Jurassic has been difficult to define owing partly to sea level fall, which caused widespread gaps in sedimentation and facies breaks at the end of the Triassic (Hallam 1997; Hallam & Wignall 2000), and to the decreasing diversity and/or extinction in fossil groups at this time. Over the past 20 years four candidates where the boundary was present and could be recognized were proposed as Global Stratotype Section and Point (GSSP) for the T–J boundary, and this number was expanded to six in 2007. Ammonites have traditionally been the markers for Jurassic GSSPs, but owing to the rarity of ammonites in the boundary beds, carbon isotope anomalies and radiolarians were also proposed as potential markers. For radiolarians, the section on the southeast side of Kunga Island, Queen Charlotte Islands was proposed by Carter & Tipper (1999) and updated by Longridge et al. (2007a, b). There, the entire Rhaetian to middle/ lower upper Hettangian sequence is exposed over a distance of c. 150 m with over 70 radiolarian collections documenting the faunal succession. Other proposed localities and their indicators were the New York Canyon area, Nevada [FAD Psiloceras spelae (Lucas et al. 2007a, b) and carbon isotopes (McRoberts et al. 2007)], St. Audries Bay, Somerset, UK and Larne, Northern Ireland [FAD Psiloceris planorbis (Simms & Jeram 2007; Warrington et al. 2008, respectively)], and Kuhjoch, Austria (FAD P. cf. spelae, von Hillebrandt et al. 2007). The task of choosing the best candidate for GSSP was the mandate of the Triassic –Jurassic Boundary Task Group (TJBTG) of the International Subcommission on the Jurassic System (ISJS). In March –April 2008, the FAD of Psiloceras spelae was chosen as the marker in the Kuhjoch section, Tyrol, Austria. In June this decision was passed by the voting members of ISJS, and in August, the Ferguson Hill section in Nevada was approved as ASSP (Auxilliary Stratotype Section and Point). Ratification by the International Commission on Stratigraphy (ICS) will comprise the final step in this decision. In the words of G. Bloos, Secretary of the TJBTG, in his final summary to task group members, ‘each section will remain an important stratigraphic

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landmark in the future’. Thus, Kunga Island will remain the standard for the radiolarian succession across the Triassic –Jurassic boundary.

Summary Usefulness and limitations During the past 30 years much effort has been paid to the biostratigraphy of Triassic radiolarians, with most publications focusing on faunas from low and middle latitude sequences. A few high latitude assemblages have been described also, but the faunas are much less diverse and the localities few and geographically limited. The main zonal units discussed in this paper (Fig. 2) represent our current knowledge of the biostratigraphy of radiolarians whose applicability and traceability has been tested satisfactorily in numerous geographic areas. The resolution of these zonal schemes is quite satisfactory, but is still less precise than scales based on ammonites and conodonts, especially in intervals devoid of continuous radiolarian-bearing sequences, that is, the Lower Triassic. But, the future is promising because the numerical and morphological diversity of genera is extremely high through the Triassic as evidenced by the range chart (Fig. 3), which offers a tool hitherto unknown in the biostratigraphy of the group. The range chart, comprised of 282 revised genera, undoubtedly presents some oversimplification because it is shown in entire (rather than partial) stages or substages increments and this implies, falsely, that many first and last occurrences happen at exactly the same time. However, even with these limitations the range chart constitutes a powerful tool for dating at substage resolution.

Future perspectives More than 1700 species have been formally described for the Triassic. This great taxonomic diversity unquestionably includes a large quantity of synonyms, but also provides an optimal background for refining the resolution of future zonations based on the range of species. But, it is important to emphasize, that the taxonomic approach must also reflect phyletic relationships and not merely external geometric ones as in the Haeckelian system. Special attention must also be given to critical boundaries (particularly those bounding the Mesozoic) in order to reduce the distortion/disconnection produced by groups of authors working either side of a particular boundary who may employ a distinctly different approach to classification. Improving the common taxonomy around these boundaries will be an active field in the coming years and only possible if, in forthcoming InterRad

meetings, new working groups are engaged to do this task. Although the history of Triassic radiolarian research is rather long, many countries remain largely unexplored. New discoveries of extremely rich faunas are still to be expected. The most recent examples of such discoveries are the Tuvalian assemblages from Turkey and Oman (Dumitrica & Hungerbu¨hler 2007; Kozur et al. 2007a, b, c; Moix et al. 2007), that contain many new genera and species. These new researches will be the key to bringing new resolution to Triassic biostratigraphic scales in the future. We thank Rie Hori, Heinz Kozur, James Ogg, and Lawrence Tanner for their thoughtful reviews of the manuscript. Many of their suggestions have been followed. This manuscript is an outgrowth of the genera revision project by the Mesozoic Working Group of INTERRAD (O’Dogherty et al. 2009a, b). We thank all additional members of the group for the measure of participation each has contributed toward this final paper on the Triassic. This study forms part of the results obtained in Research Projects: CGL2005– 02500, financed by the DGI (Direccio´n General de Investigacio´n, Spain), and by the RNM– 027 Research Groups, Junta de Andalucı´a, Spain). Further funding for this work was provided by research program P1– 0008 of the Slovenian Research Agency.

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Biochronology of Triassic bivalves CHRISTOPHER A. MCROBERTS Department of Geology, State University of New York at Cortland, P.O. Box 2000, Cortland, NY 13045, USA (e-mail: [email protected]) Abstract: Substantial advances by numerous researchers over the past 20 years have made it possible to develop a composite biochronological scheme for the Triassic based on the bivalves Claraia, Peribositria, Enteropleura, Daonella, Halobia, Eomonotis and Monotis. These bivalves exhibit temporal durations nearly equal to ammonoids and conodonts. Widely distributed across the Tethys, Panthalassa and Boreal regions, these bivalves occur in a wide variety of marine facies and water depths, but are most notable for their thick shell accumulation in deeper-water oxygen deficient environments. They were most likely resting or reclining benthos, may have housed chemosymbionts, and were part of episodic opportunistic palaeocommunities in or near oxygen deficient settings. A new biochronological zonation for bivalves is presented that encompasses the entire Triassic and is integrated with standard ammonoid schemes. The Lower Triassic is characterized by 2– 3 zones of Claraia, most notably from the eastern Tethys representing the entire Induan and lower portion of the Olenekian. Later in the Olenekian, and most notably from the Boreal realm, species of Peribositria (included by some workers within Bositra) provide useful zonal indexes. The Middle Triassic is well represented by Enteropleura (Middle Anisian) and Daonella (Upper Anisian through Ladinian) in the Tethys and North America with significant occurrences throughout the circum-Pacific and Boreal realms. The Upper Triassic can be subdivided into 8 –13 bivalve zones based on the succession of Halobia, Eomonotis and Monotis sensu lato species with best representation in the Tethys, Boreal and eastern Panthalassa regions.

Following the end-Permian mass extinction, benthic marine ecosystems underwent a remarkable reorganization associated with the post-extinction recovery and subsequent adaptive radiation during the Mesozoic. Among shelly benthos, this reorganization resulted in the diminishment of typical Late Palaeozoic brachiopod, pelmatozoan, and stenolaematid bryozoan dominated level-bottom marine communities in favor of communities characterized by bivalve and gastropod molluscs, bony fishes, gymnolaemate bryozoans, echinoids and certain crustaceans, representing an increase in trophic variety ranging from deep infaunal suspension feeders to active nektonic carnivores (Vermeij 1977; Sepkoski 1981). A significant component of this essentially modern fauna are bivalve molluscs, which in the Triassic dominate many level-bottom, reefal, and pelagic settings (McRoberts 2001; Fraiser & Bottjer 2007). Among the bivalves of the Triassic radiation, those belonging to the genera Claraia, Peribositria, Enteropleura, Daonella, Aparimella, Halobia and Monotis, are arguably the most abundant macrofauna in deep-water marine facies of the Triassic (Fig. 1). Often referred to as ‘flat clams’ because of their very thin shell width and narrow valve convexity, these bivalves generally have widespread distributions and very high species turnover rates making them excellent biochronological macrofossils of the Triassic. Their biochronological value

was recognized soon after their discovery, and within the past 30 years they have become standard components of biochronological schemes (e.g. De Capoa Bonardi 1984; Polubotko 1986; McRoberts 1997; Zakharov et al. 1997; Yin 2003). Even though bivalve zonations have been provided for specific regions and or limited portions of the Triassic, this paper provides the first global summary of bivalve zonation for the entire Triassic (Figs 1, 3, 5, 6).

Taxonomic and nomenclature issues Perhaps the largest problem with the use of Triassic bivalves in biochronological studies is a combination of the multitude of available species names, many of which are in open nomenclature, and a high degree of uncertainty in species and higher taxonomy of this heterogeneous grouping of bivalves. These issues are due to several factors including: poor understanding of significant morphological traits and their variation in natural populations, their frequently poor preservation resulting in a misinterpretation of such traits, and palaeogeographical/palaeooceanographical provincialism that is in many cases unrecognized. As construed herein, Triassic ‘flat clams’ represent a morphologically and taxonomically diverse group belonging to at least three separate bivalve

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 201– 219. DOI: 10.1144/SP334.9 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Temporal distribution of Triassic flat clam bivalve genera with biochronological utility.

clades distributed across four family-group taxa: the Posidonidae, Halobiidae, Monotidae and Pterinopectinidae. Further compounding the taxonomic clutter has been the introduction of lower-level groups (genera and subgenera) which have proliferated in recent years but which have often been constructed on limited morphological evidence without regard to natural morphological variation within species or populations or phylogenetitc considerations. The primary genus-group taxa used here are broadly construed (sensu lato) and include Claraia,

Ellesmerella, Enteropleura, Daonella, Halobia, Eomonotis, and Monotis. Brief summaries of these accepted taxa and comments on group membership and affinities are provided in Appendix 1 and representative species are illustrated in Figures 4 and 7. The sheer number of species in some of these genera is another subject of concern. For example, following Mojsisovics’ (1874) impressive monograph, Kittl (1912) alone recognized 66 species of Daonella and 116 species of Halobia. More recent workers have not fared much better and the pace of

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naming species has accelerated. In the past 50 years, workers have introduced more than 50 new species of Monotis, more than 75 new species of Daonella, and more than 200 new species of Halobia. Not surprisingly, a vast majority of these new species are based on wholly inadequate material – either very small sample sizes of only one or two specimens and without regard to natural morphological variation within populations, or they are based on taphonomically compromised specimens. While a thorough assessment of individual species validity is beyond the scope of this paper, an attempt has been made to include only species that have biochronological value, are morphologically distinctive, are based on sufficient material, and are generally accepted by modern workers.

Palaeoecological and palaeogeographical controls on biochronology Although the mean duration of bivalve species is estimated to be c. 15 Ma (Stanley 1979), Triassic flat clam species are closer to the average durations for ammonoid species (1 –2 Ma or less). What makes these bivalves different from other level-bottom bivalves likely lies in their evolutionary palaeoecology which can, in the absence of modern analogs or genetic information, be inferred from morphological adaptations, their unique facies occurrence, and analogous yet anecdotal similarities from other non-pelagic bivalves.

Palaeoecology A significant attribute of Triassic flat clams is their episodic occurrence in monospecific or paucispecific shell accumulations. Although many of these shell beds occur in deep-water oxygen deficient settings, they also occur in significant abundance in deep so called ‘pelagic’ or ‘filamentous’ limestone (e.g. Hallstatt facies) that represent fully oxygenated marine settings (McRoberts et al. 2008). In addition, many examples are known from presumably shallower-water environments from Russia, New Zealand and Japan among other regions (see Miroshinikov & Burii 1969; Ando 1987; Campbell 1994). Within these settings, Triassic flat clams occur within a continuum of accumulation modes that are interpreted to be nearly in situ to occasional or episodic shell accumulations with varying amounts of re-working and timeaveraging. Although some taxa (e.g. Claraia and Enteropleura) typically form thin shell beds less than 10 cm thick, the shell beds comprised of other taxa (e.g. Halobia and some Monotis) can achieve thicknesses of greater than 2 m (McRoberts et al. 2008; Grant-Mackie, pers. comm.). A ubiquitous

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feature of the bivalve accumulations is that they are almost exclusively monospecific, or, more rarely, composed or two or three species. Similar shell beds are often labeled as ‘stagnation deposits’ (e.g. Seilacher et al. 1985) and may contain an admixture of palaeoecological and/or time-averaged associations. Following Kidwell & Bosence (1991), many are best described as either census assemblages or within-habitat time-averaged assemblages. These accumulations can be interpreted as in situ biogenic deposits representing episodic population expansion or immigration. An extensive literature exists on the hypothesized living habit of Triassic flat clams (see Gruber 1976; Ando 1987; Wignall & Simms 1990; Campbell 1994; Schatz 2005; Waller in Waller & Stanley 2005 for useful summaries). While there certainly exists morphological disparity between genera, and no one-size-fits-all approach with respect to their autecology, Triassic flat clams are most likely either freely resting or reclining without byssal attachment (e.g. Peribositria, Enteropleura, Daonella and Halobia) or attached with a feeble or weakly functioning byssus (e.g. Claraia, Eomonotis, and Monotis). Although it has also been proposed that several of these genera were pseudoplanktonic and attached to drifting seaweed, wood, or other epiplanktic host (e.g. Jefferies & Minton 1965; Tozer 1982; Wignall & Simms 1990; McRoberts 1997), shell-bed attributes, facies occurrences, and lack of demonstrably attached specimens point to a benthic life mode for all of the genera considered here. A benthic living habit, however, does not in itself explain their unusual occurrence nor their biochronological utility. It is also possible that some Triassic flat clams harbored symbiotic bacteria that enabled them to live in the presence of hydrogen sulfide in dysaerobic palaeoenvironments. Seilacher (1990) proposed that the anterior tube of Halobia served as a sulphur pump and perhaps served as the locus for chemosynthetic bacteria. This view was recently advocated by Waller (in Waller & Stanley 2005) who reinterpreted structures on Halobia’s anterior tube interior as radial mantle muscle scars and thus suggesting open passage from the tube to the mantle cavity. It should be noted, however, that neither Daonella (ancestral taxon to Halobia) nor other halobiids, monotids, or claraids exhibit this shell tube, yet occur in similar settings and shell accumulation modes. It should be further noted that extant bivalves with sulphur-oxidizing or methanotrophic endosymbionts (e.g. lucinid, thyasirid and solemyid bivalves) typically have relatively large gills which host the endosymbionts which would not be accommodated in the slender valve cavity of the Triassic forms (most notably halobiids and claraids). Furthermore, known chemoautotrophic endosymbionts

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and methanotrophs occur where both sulfide and limited oxygen co-exist (Scott & Cavanaugh 2007), and that free oxygen is necessary for metabolic function employing glycolysis (Oschmann 1993). Hydrothermal vent and, to a lesser extent, methane seep faunas are notably dominated by endemic species (Tunnicliffe 1991) which is atypical for these Triassic bivalves. Lastly, although the Triassic flat clams are notable for their denselypacked accumulations, which resemble some of the dense faunal communities surrounding methane seeps, the shell beds are laterally continuous in exposures up to one kilometer in lateral extent, which is atypical for modern chemo-symbiotic communities whose distribution is patchy. Based on facies occurrences and shell-bed attributes, many of the Triassic flat clams can be best described as belonging to opportunistic populations which Levinton (1970) described as unstable populations that are not resource limited but primarily controlled by the physical and not the biotic environment. Opportunistic fossil populations appear to be a major component of the exaerobic biofacies (sensu Savrda & Bottjer 1987, 1991) that are characterized by organic-rich laminated sediments and epibenthic macrofossil assemblages in oxygen deficient (dysaerobic) benthic conditions. Claraids, halobiids and monotids in particular fit the opportunistic model in several ways as demonstrated by several empirical studies on similar biofacies. Wignall (1993) established a positive correlation between decreasing oxygen gradients and decreasing benthic diversity, reaching zero at essentially the anaerobic/dysaerobic boundary. Gradient analyses of Jurassic oxygen-deficient faunas by Etter (1995) demonstrate the macrofaunal succession of equilibrium species being displaced by low diversity, high abundance opportunistic species and again by low abundance higher diversity chemosynthetic specialists at the lower end of the dysaerobic/ anaerobic interface. Many of these Triassic flat clam palaeocommunities are interpreted to have inhabited and dominated environments near a threshold oxygen minimum boundary which other shelly benthos found unsuitable. The spatial distribution and mode of occurrence of these bivalve palaeocommunities conform to an incumbent dominance model (e.g. Rosenzweig & McCord 1991; Sepkoski 1996) where the first individuals of a species to occupy a niche, dominate and hold the spot until replaced or removed from that niche. It has also been proposed that rapid fluctuations in environmental conditions tend to favor eurytopic species with high genetic diversity (Sanders 1968; Bretsky & Lorenz 1970). The episodic nature of many of the shell beds suggests that Triassic flat clams also appear to have exhibited low resistance to environmental perturbations but are resilient in being able to recover

quickly (see Tang 2001 for a review of resistant v. resilient species). With respect to flat clams in general, and Monotis in particular, Shaner & McRoberts (2000) proposed that once colonization occurs, the bivalve assemblages dominate substrate to the exclusion of other benthic organisms. This dominance may continue even if oxygen conditions become favorable. This model of episodic replacement of bivalve assemblages provides ample room for the repeated colonization of different flat-clam species and even genera along oxygen or other resource threshold boundaries.

Palaeobiogeography The assembly of the supercontinent Pangaea by Triassic time had resulted in more than half of the Earth’s surface being covered by the Panthalassa Ocean, which spread from pole to pole, and an eastwest tropical seaway – the Tethys (Fig. 2). Within these broad palaeogeographical constraints, geographical distribution of Triassic bivalves with biochronologic significance is controlled by the location of suitable environments. Major constraints in the distribution of bivalves are bathymetry, substrate, temperature, chemistry (e.g. dissolved oxygen), and oceanographical currents needed for dispersal. Palaeocurrents within Panthalassa and the Tethys can be determined through actualistic principles applied to palaeogeographical reconstructions and have been summarized by several workers (e.g. Tollmann & Kristan-Tollmann 1985; Yan & Zhao 2001). In general, the major palaeocurrents of the Triassic can be summarized as: (a) a low latitude north and south equatorial current from the margin of Pangaea moving westwards across Panthalassa and into the Tethys with an equatorial counter current moving eastward; (b) a clockwise gyre in northern Panthalassa feeding warm currents along the northeast margin of Pangaea and a coldwater current moving down the northwest margin of Pangaea; and (c) a counterclockwise gyre in southern Panthalassa bringing warm water across southeastern Tethys, Australia and perhaps as far south as New Zealand and cooler water northwards adjacent to the Andean margin of Pangaea. It should be noted, however, that distributional patterns of many Triassic macrofauna are quite complex and may require additional dispersal models, such as the Pantropical model of Newton (1988) or tectonic displacement or some combination thereof (e.g. Silberling 1985; Stanley & Gonza´lez-Leo´n 1995). Dispersion of bivalve species across Panthalassa and the Tethys would most likely have been through passive transport of larvae via surface currents. Although some Triassic flat clams may have had planktotrophic larvae (McRoberts 2000), almost nothing is known about the larval development of most species. Regardless, even short-lived planktic

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Fig. 2. Biogeographical distribution of significant occurrences of Triassic bivalves. Abbreviations for informal regions include: WT, Western Tethys; NIM, Northern Indian Margin (e.g. Salt Range, Kashmir); CI, Cimmeria (e.g. Tibet); NC, North China; SC, South China; IND, Indonesia; JP, Japan; MA, Maori; AN, Andean; EP, Eastern Panthalassa including pericraton regions (e.g. Nevada and northeastern British Columbia) and allochthonous terranes (e.g. Wrangallia, Stikinia); BO, Boreal (Arctic Alaska, Arctic Canada, Greenland, Svalbard, northern Siberia); OK, Okhotsk (e.g. northeastern Siberia). Palaeogeographic reconstructions modified from Blakey (2005).

larva could be transported great distances – for example, larvae having a pelagic stage of 2–6 weeks could be transported up to 500 km by only moderate coastal currents; longer pelagic larval

stages (e.g. 3 months or more) could effectively cross entire deep ocean basins (e.g. Sheltema 1977). Except for the broadly construed Tethyan, Panthalassan, and Boreal realms, Triassic seas are

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not easily differentiated into biogeographical provinces. With few notable exceptions discussed later, Triassic flat clams are distributed across most regions in which marine strata are preserved. Although too numerous to list here in detail, Triassic pelagic bivalves commonly occur in the Western Tethys (including the northern and southern Alps, Sicily, Carpathians, and into the Balkans). Major occurrences in the central and eastern Tethys are also significant, with many species known from both the southern branch of the Tethys northern Indian margin (e.g. Salt Range) as well as from the central and northern branches (e.g. Caucasus, Crimea, Kashmir, Spiti, Nepal and Tibet). In the far eastern Tethys, significant occurrences of flat clams are also known from both the North and South China Blocks, Timor, and in many localities in what is now southeast Asia (e.g. Thailand, Vietnam, etc.). The Boreal region appears to be a center of diversity among bivalve species having significant occurrences in northern and northeastern Siberia, northern Alaska, the Canadian Arctic Archipelago, northeast Greenland and Svalbard. The northwest margin of Panthalassa includes Primorye and the terranes of Kamchaka and Korayak. Many of these regions contain mixtures of faunas from more than one province. For example, bivalves of low to mid palaeolatitudes of western offshore Pangaea (e.g. Nevada and northeastern British Columbia) contain a mixture of Boreal

and Tethyan elements (see discussion in Silberling 1985; Newton 1988).

Biochronology and bivalve zones of the Triassic This section presents a summary of the temporal distribution and global/interregional correlation of bivalve (chiefly flat clam) occurrences having demonstrably biochronological significance during the Triassic. These data are based on an extensive literature compilation and first-hand study of field and museum collections, and integrated, to the extent possible, with other biochronologically important fossils, most notably ammonoids and conodonts. The temporal ranges of species are depicted here as zones and, except where noted, these are total range zones and are based on the ranges of individual species. These zones, however, remain tentative and informal as they lack appropriate stratal and geographical definitions. Where possible, the bivalve ranges are resolved to standard ammonoid zones as depicted in the zonal charts (Figs 3, 5 & 6). In species whose range is demonstrably shorter than an ammonoid zone (e.g. uppermost Olenekian Ellesmerella, uppermost Anisian Daonella or uppermost Norian Monotis) the zonal limits may be artificially extended upwards and downwards for the entire interval.

Fig. 3. Bivalve zones of Early Triassic Claraia, Peribositria and Ellesmerella for the western and eastern Tethys (including both northern Indian margin and South China), and Boreal realm. Sources for bivalve zonation discussed in text. Ammonoid zones modified from Orchard & Tozer (1997), Zakharov et al. (1997) and Kozur (2003).

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Temporal assignments rely on integration of several zonal schemes from various taxa taken across widely distributed regions as discussed in each appropriate section below. Stage and substage nomenclature and definitions are largely taken from Ogg (2004) and in several cases have been modified as recent advances in selection of GSSPs in defining the standard chronostratigraphic scale of the Triassic.

Early Triassic As with other parts of the Triassic, biochronological schemes of the Early Triassic are largely based upon conodonts and ammonoids. Following the formal definition of the base of the Triassic with the first occurrence of the conodont Hindeodus parvus from the base of bed 27c at Meishan South China (Yin et al. 2001), the Triassic Subcommission on Stratigraphy has recently adopted the use of subdividing the Early Triassic into two stages, the Induan and Olenekian. Progress has been made at defining the base of the Olenekian at the Mud section of Spiti by the first occurrence of the conodont Neospathodus waageni sensu lato which is less than one metre above the base of the Rohillites rohilla ammonoid Zone (Krystyn et al. 2007a). The Lower Triassic is subdivided into between 10 and 15 ammonoid zones (Kozur 2003) and recent estimates suggest that the entire Early Triassic was of only 5 Ma or so (Mundil et al. 2004; Ogg 2004) in duration. As mentioned earlier, the Early Triassic is dominated by low diversity but locally abundant bivalve faunas, which largely include representatives of Permian hold-over clades (e.g. Permophoridae, Myalinidae and Pterinopectinidae). Included in this group are species Eumorphotis and Promyalina, and most importantly the flat clam Claraia that has long been recognized as a zonal taxon for the Early Triassic. Although Claraia has often been regarded as a cosmopolitan taxon, a majority of its species are of local or regional importance. It should also be noted that Claraia species are quite rare in the lowermost Triassic Otoceras beds and are often found above the disaster fauna (e.g. microbialites and Lingula zone) associated with the end-Permian mass extinction. Furthermore, individual species ranges (and subsequent zones) for Claraia are somewhat long when compared to ammonoid ranges of the same interval, yet they are comparable to bivalve zones of the Middle and Late Triassic. In spite of this, zonal schemes for Claraia have been developed for a number of Tethyan, Boreal, and circum-Pacific regions (Fig. 3). Of the numerous successions, those of the eastern Tethys, including the Northern Indian Margin and South China are the best known and integrated into other biochronological scales using

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conodonts and ammonoids. Nakazawa (1977, 1981) developed a zonation based on Claraia from the uppermost Permian through Lower Olenekian (Meekoceras beds of the Smithian) from Kashmir and Iran. Nakazawa listed 33 Claraia species belonging to 4 species groups. A complimentary zonation from south China was provided by Yin (1985, 1990) who identified six Claraia acme zones distributed across much of the Lower Triassic. Presented here is a composite zonation beginning with a zone of Claraia dieneri Nakazawa correlative to the Otoceras woodwardi Zone that is well established in Kashmir, South China, and the Perth Basin of Australia (Nakazawa 1977; He et al. 2007). A zone of C. wangi Patte associated with the conodont Isarcicella isarcica and the ammonoid Ophiceras sinense and largely restricted to South China and Vietnam is in part correlative with C. dieneri Nakazawa, yet ranges into somewhat younger strata (Yin 1990; Tong & Yin 2002). Claraia griesbachi (Bittner) and C. concentrica (Yabe) also occur in the Lower Induan (Ophiceras –Lytophiceras ammonoid Zone) of Anhui Province China (Tong et al. 2004). Claraia wangi Patte has been reported from several localities in south China and is frequently associated with the ammonoid Ophiceras and the conodont Isarcicella isarcica but never with lowermost Induan ammonoids Otoceras woodwardi nor the conodont Hindeodus parvus. This east Tethyan zone can be correlated to the westernTethys, where at the Tesero section in the Italian Dolomites, Claraia griesbachi (Bittner) first appears about 15 –20 m from the base of the Mazzin Member of the Werfen Formation (e.g. Broglio Loriga et al. 1986; Broglio Loriga & Cassinis 1992; Wignall & Hallam 1992). Tethyan Claraia can be correlated to North America, where several species including C. clarae (Emmrich) (Fig. 4h) and C. stachei Bittner occur in the central Rocky Mountains in the western United States (Newell & Kummell 1942; Ciriacks 1963; Newell & Boyd 1995) and also northern Rocky Mountains in British Columbia and Alberta Canada (Tozer 1970). It should be noted, however, that reports of Permian C. stachei (e.g. Muromtseva 1984) have not been adequately illustrated, and their validity remains in question. Generally above the sequence of C. stachei Bittner is a zone of C. aurita (Hauer) that has proven to be of regional significance extending from the western Tethys eastward into the Praecaucasus, the northern Indian Margin, South China and northeastern Russia and with a little more precision to the Tirolites harti ammonoid Zone (Gavrilova 2004). Above the Claraia zones of the Induan and lowermost Olenekian are several demonstrable zones of Peribositria that are largely relegated to the Boreal realm. Several Peribositria species are

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Fig. 4. Representative Lower and Middle Triassic pelagic bivalves. All scale bars ¼ 1 cm. (a) Daonella elegans McLearn, Upper Ladinian, Frankites sutherlandi Zone from Toad Formation, Halfway River area, British Columbia, Canada (SUCP collection). (b) Daonella subarctica Popov, Upper Ladinian, Botneheia Formation, Svalbard (SM M1285, from Campbell 1994, pl. 5, fig. 2). (c) Daonella moussoni (Merian), Upper Ladinian, Prototrachyceras gredleri Zone (PIMZ collection, from Rieber 1969, pl. 2, fig. 3). (d) Daonella cf. D. sturi (Benecke), Upper Anisian, Parafrechites meeki Zone, Favret Formation, Lone Peak, Nevada (SUCP collections). (e) Enteropleura jenksi Hopkin & McRoberts, Middle Anisian, Balatonites shoshonensis Zone, from Favret Formation, New Pass Range, Nevada (USNM 526057, from Hopkin & McRoberts 2005, fig. 3.1). (f ) Peribositria mimer (Oberg). Upper Olenekian, Otuk Formation, Brooks Range, Alaska (SUCP collection). (g) Ellesmerella aranea (Tozer), Upper Olenekian (Smithian) Svalbard (SM G1281, from Tozer & Parker 1968, pl. 25, fig. d). (h) Claraia clarae (von Hauer), Lower Induan, Brazeau River, Alberta (GSC 14196, from Newell & Boyd 1995, fig. 17.7; see also Tozer 1961, pl. 28, fig. 3).

relegated to the lower Olenekian including the most common forms P. borealis Kurushin & Trushchelev and P. mimer Oeberg (Fig. 4f ), which comprises two distinctive zones of the upper Smithian. Tozer (1963) reports P. mimer Oeberg from the Euflemingites romunderi Zone of the Toad Formation of northeastern British Columbia, from the same interval in the Blind Fiord Formation of Ellesmere Island in Arctic Canada and likely from the Yukon area. Other species of Peribositria, including P. backland (Wittenburg) from the lower Olenekian from Siberia and Svalbard (Kurushin & Trushchelev 1989), are rather poorly constrained but are known to co-occur with the ammonoids Arctoceras blomstrandi and Euflemingites and Hedenstroemi equivalents of the Early Olenekian (Kurushin & Trushchelev 1989). Somewhat longer in duration

is P. siberica Kurushin (1980) which seems to have some biochronological value (Zakharov et al. 1997), yet it is long-ranging, encompassing nearly the entire Spathian (Bajarunia euomphala through Parasibirites grambergi ammonoid Zones) from central Siberia, Arctic Canada, and British Columbia (Kurushin & Trushchelev 1989). In high to mid northerly latitudes, Ellsemerella aranea (Tozer) (Fig. 4g) comprises the youngest bivalve zone of the Olenekian and is fairly well constrained by ammonoids of the Keyserlingites subrobustus ammonoid Zone. Tozer (1973) summarized the distribution of this species which is now known from several localities in Svalbard (Tozer & Parker 1968), and Ellesmere Island in the Canadian Arctic (Tozer 1961, 1967), northeastern British Columbia (Tozer 1963, 1970) and northeastern Siberia

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Fig. 5. Bivalve zones of Middle Triassic Peribositria, Enteropleura and Daonella for the western Tethys, North America and Boreal realms. Sources for bivalve zonation discussed in text. Ammonoid zones modified from Mietto & Manfrin (1995), Orchard & Tozer (1997), Zakharov et al. (1997), Kozur (2003) and Monnet & Bucher (2005).

Fig. 6. Stratigraphical ranges of Late Triassic, Halobia, Eomonotis, and Monotis for western Tethys, North America and Boreal realms. Ammonoid Zones modified from Orchard & Tozer (1997) Zakharov et al. (1997), Kozur (2003) and Gallet et al. (2007).

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(Vozin & Tikhomirova 1964; see also Zahkarov et al. 1997).

Middle Triassic The Middle Triassic comprising the Anisian and Ladinian stages lasted approximately 10– 12 Ma (Brack et al. 2005; Lehrmann et al. 2006; Ovtcharova et al. 2006). The Anisian, whose base is as yet undefined, contains 7–13 ammonoid zones (Silberling & Tozer 1968; Tozer 1994; Mietto & Manfrin 1995; Orchard & Tozer 1997; Monnet & Bucher 2005). As shown in Figure 5, the Ladinian, whose base is now accepted to be the first occurrence of Eoprotrachyceras curionii (Brack et al. 2005), can be subdivided into approximately seven standard ammonoid zones (Tozer 1994; Mietto & Manfrin 1995). No global or regionally applicable bivalve zones have been identified for the Lower through Middle Anisian. The first globally useful bivalve zone of the Middle Triassic occurs in the upper part of the Middle Anisian (Balatonites shoshonensis ammonoid Zone) where several species of Enteropleura (Fig. 4e) appear in different and widely-spaced regions including the west Tethys, North America and China (Waller in Waller & Stanley 2005; Hopkin & McRoberts 2005; Chen & Stiller 2007). Beginning at the base of the Gymnotoceras weitschati ammonoid Zone, a succession of Daonella species permits subdivision of the Upper Anisian through uppermost Ladinian into approximately 5–7 Daonella zones of regional and global significance. Perhaps the best-represented succession of daonellas is known from the Monte San Giorgio section, Switzerland (Rieber 1968, 1969; Schatz 2004), and from the Italian Dolomites (Brack & Rieber 1993). A composite succession begins with Daonella sturi (Benecke) in the Upper Anisian (Paraceratites trinodosus Zone) overlain by the Daonella elongata Mojsisovics group (co-occurring with D. pseudomoussoni Rieber) that dates as uppermost Anisian and lower Ladinian (Reitziites retzi through Eoprototrachyceras curionii Zones). Above this is a distinct zone containing D. moussoni (Merian) (Fig. 4c) which ranges close to the base of the upper Ladinian (Prototrachyceras gredleri Zone), although it may extend a short time above its uppermost boundary. This is in turn followed by two successive Daonella zones: D. picheri Mojsisovics overlain by D. lommeli (Wissmann) that are well represented in the Southern Alps and provide useful correlations to elsewhere in the Tethys and North America. A North American sequence of daonellas can be constructed using the well documented succession from western Nevada (e.g. Silberling & Tozer 1968; Silberling & Nichols 1982) and northeastern British Columbia, Canada (e.g. Tozer 1967, 1970).

The sequence provides several links to the Tethyan succession (e.g. D. moussoni (Merian) near the middle/upper Ladinian boundary) and with the Boreal Realm (e.g. D. americana Smith and D. dubia (Gabb) in the upper Anisian). In North America, the youngest Daonella of the Ladinian, D. elegans McLearn, is well constrained to the Frankites sutherlandi ammonoid zone from northeastern British Columbia (Tozer 1967) and possibly from western Nevada (Balini et al. 2007). Although they remain poorly documented, several occurrences of Daonella are known from lower Carnian strata overlapping with true Halobia.

Upper Triassic The Upper Triassic is subdivided into three stages: in ascending order Carnian, Norian and Rhaetian. The basal Carnian has recently been defined as the base of the Daxatina canadensis ammonoid Zone (Mietto et al. 2007). Several potential candidates are being considered for a base-Norian GSSP including ammonoids (e.g. Stikinoceras kerri) slightly above the first occurrence of Epigondolella primitius datum and below several other potential conodont-based levels including first occurrence of Epigondolella quadrata (see Orchard 2007). Of the three potential levels for the base Rhaetian, the conodont Misikella posthernsteini is the leading candidate (Krystyn et al. 2007b). The Carnian through middle Norian can be subdivided into several zones based on Halobia species (Fig. 6). Zonal schemes are provided for the western Tethys Alpine-Mediterranean belt (taken from Gruber 1976; Cafiero & De Capoa Bonardi 1982; De Capoa Bonardi 1984), western North America including the allochthonous and autochthonous elements of British Columbia, southeastern Alaska, Oregon and Nevada (McRoberts 1997, personal observations), and the Boreal regions encompassing northeastern Siberia (Kiparisova et al. 1966; Polubotko 1980, 1984, 1986), Arctic Alaska (Blome et al. 1988; Dutro & Silberling 1988), Arctic Canada (Tozer 1961, 1967), and Svalbard (Campbell 1994). It should be noted that other regions including China (Chen 1964), SE Asia and Melanesia (e.g. Krumbeck 1924; Vu-Khuc 1991; Chonglakmani & Grant-Mackie 1993) and New Zealand (e.g. Marwick 1957; Campbell 1994) have significantly contributed towards our understanding of Halobia biochronology and provide valuable interregional and global correlations. The Early Carnian Trachyceras desatoyense through Sirenites nanseni ammonoid zones and equivalents are rather poorly represented by pelagic bivalves in most regions. It is perhaps best represented by Halobia zitteli Lindsto¨m (Fig. 7k) in the Boreal regions of Russia, Canada, and

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Svalbard (e.g. Tozer 1961; Polubotko 1980; Campbell 1994). In NE Siberia, H. zitteli Lindsto¨m may be restricted to the earliest Carnian (Trachyceras desatoyense Zone) and is subsequently overlain by four halobiid zones including H. zhilnensis Polubotko with H. talajaensis Polubotko, H. popowi Polubotko, H. ornatissima Smith with H. subfallax Efimovae, and H. asperella Polubotko (Polubotko 1980, 1986; Zakharov et al. 1997). In western North America, however, H. zitteli Lindsto¨m is poorly represented, and H. rugosa Mojsisovics is more characteristic of the lower Carnian, yet it too may extend into the upper Carnian extending into the Klamathites macrolobatus Zone or even Stikinoceras kerri Zone where it co-occurs with H. superba Mojsisovics and H. radiata Gemmellaro (McRoberts 1997, 2007, personal observations). The base of the Norian stage (as yet undefined, but here taken as the first occurrence of the ammonoid Stikinoceras kerri) closely corresponds to the first occurrence of Halobia styriaca (Mojsisovics) (Fig. 7i) and H. beyrichi (Mojsisovics) and similar species best represented in low palaeolatitudes with notable occurrences in the western and eastern Tethys and eastern Panthalassa (Gruber 1976; De Capoa Bonardi 1984; McRoberts 1997). A temporally-limited zone spanning the CarnianNorian boundary is based on H. radiata Gemmellaro that is well represented across the Tethys, Boreal, and low to mid latitude strata in western North America (De Capoa Bonardi 1984; Nicora et al. 2007; McRoberts 2007). Above the zone of H. beyrichi (Mojsisovics) and H. styriaca (Mojsisovics) are a suite of similar forms referable to H. cordillerana Smith (Fig. 7g), H. obruchevi Kiparisova in NE Russia (e.g. Kiparisova et al. 1966; Polubotko 1984) and H. hochstetteri Mojsisovics in New Zealand (Campbell 1994). In the western Tethys, this interval is characterized by a sequence of zones in ascending order: H. mediterranea Gemmellaro, H. darwini Cafiero & De Capoa Bonardi, H. halorica Mojsisovics, H. norica Mojsisovics, and H. distincta Mojsisovics (Gruber 1976; De Capoa Bonardi 1984). Of particular importance is H. halorica Mojsisovics, a short-lived species found in the western Tethys, Russia and in North America which is well constrained to the middle Norian Drepanites rutherfordi into lowermost Mesohimavitites columbianus Zones and equivalents (De Capoa Bonardi 1984; McRoberts 1997). Eomonotis contains at least 10 species, primarily in the circum-Pacific, and notably absent in the western Tethys, that are all largely restricted to the upper part of the Middle Norian (e.g. below the conodont Epigondolella bidentata). The oldest Eomonotis, E. daonellaeformis Kiparisova and slightly younger E. pinensis (Westermann) (Fig. 7e) are found in association with Halobia near the first

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appearance of the conodont Epigondolella postera and ammonoid equivalents (e.g. Mesohimavitites columbianus II Zone) in western North America and with E. scutiformis (Teller) at approximately the same level in far-east Russia and New Zealand (Westermann & Verma 1967; Bychkov et al. 1976; Grant-Mackie 1980; Grant-Mackie & Silberling 1990). Above Eomonotis, the Upper Norian (Gnomohalorites cordilleranus Zone and equivalents) is characterized by numerous species of Monotis (sensu lato) that have proven to be of global biochronological significance (e.g. Ichikawa 1958; Westermann 1973; Bytchkov et al. 1976; Grant-Mackie 1978; Tozer 1980; Silberling 1985; Silberling et al. 1997). Although there are about 30 well-defined species of Monotis that demonstrably occur in the upper Norian, only a few are noteworthy because of their regional and global correlation potential. These include the distinctive M. subcircularis Gabb in the circum-Pacific, Monotis ochotica (Keyserling) across the circumPacific and into the Boreal Realm, and Monotis salinaria (Schlotheim) common in the Tethyan Province and perhaps elsewhere in the circumPacific. Some authors have differentiated an uppermost Norian into two zones, a lower consisting of M. subcircularis Gabb (Fig. 4f ) and an upper zone of M. ochotica (Keyserling) and M. alaskana Smith in certain mid to high-latitude settings (e.g. Grant-Mackie & Silberling 1990; Silberling et al. 1997). It should be noted that both M. subcircularis and M. ochotica provide valuable correlations to the upper Norian of South America (e.g. Jaworski 1922; Westermann 1970; Geyer 1973) which are mostly lacking in marine strata during earlier Triassic times. It has been long assumed that Monotis experienced mass extinction close to the Norian/Rhaetian boundary as defined by first appearance of the conodont Misikella posthernsteini and its ammonoid equivalents (various Paracochloceras and perhaps Sagentites reticulatus). Recent finds in the Hallstatt limestone in Austria record two surviving and dwarfed species of Monotis above this event, one of which, Monotis rhaetica McRoberts et al. (Fig. 7d) is new (McRoberts et al. 2008). Although species of Otapiria are known from strata as old as Carnian and continue through the Early Jurassic, they take particular temporal significance in the Norian and Rhaetian. For example, Otapiria ussuriensis (Voronetz) is the name bearer of a well established zone in the Middle Norian of Siberia (Kiparasova et al. 1966; Zakharov et al. 1997) and Otapiria dissimilis (Cox) is distinctive of the Rhaetian of New Zealand (Marwick 1957; Campbell 1997) and the closely related Otapiria alpina Zapfe (Fig. 7b) is restricted to the Rhaetian of the Austrian west-Tethys (Zapfe 1973).

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Fig. 7. Representative Upper Triassic pelagic Bivalvia. All scale bars ¼ 1 cm. (a) Tosapecten efimovae Polubotko, Rhaetian, Olomon River area, NE Siberia, Russia. (TsNIGR collection, same specimen as in Bychkov et al. (1976, pl. 71, fig. 4). (b) Otapiria alpina Zapfe, Rhaetian, Zlambach Formation, Rossmoosgraben, Austria (NHMW 1973/ 1623, from Zapfe 1973, pl. 1, fig. 1). (c) Rhaetavicula contorta (Portlock), Rhaetian, Ko¨ssen Formation, Hochalm, Austria (SUCP collection). (d) Monotis rhaetica McRoberts et al., Lower Rhaetian, Paracochloceras suessi Zone, Hernstein, Austria (NHMW 2007z0111/0012, from McRoberts et al. 2008, pl. 1, fig.15). (e) Eomonotis pinensis (Westermann), Middle Norian, Mesohimavites columbianus III Zone, Peace River area, British Columbia, Canada (SUCP collection). (f ) Monotis subcircularis (Gabb), Upper Norian, Gnomahalorites cordilleranus Zone, California (USNM 32492, see Martin, 1916, pl. 30, fig. 2). (g) Halobia cordillerana Smith, Middle Norian, Drepanites rutherfordi Zone, Peace River area, British Columbia, Canada (SUCP collection). (h) Halobia lenticularis Gemmellaro, Probably uppermost Carnian, Pizzo Mondello, Sicily (UNIMI FNP67/68, from Nicora et al. 2007, pl. 2, fig. 1). (i) Halobia beyrichi Mojsisovics, Lower Norian, Stikinoceras kerri Zone, Vancouver Island, British Columbia, Canada (GSC 13712. ( j) Halobia styriaca (Mojsisovics), Lower Norian, Guembelites jandianus Zone, Sommerkogel, Austria (SUCP collection). (k) Halobia zitteli Lindstro¨m, Lower Carnian, Axel Heiberg Island, Arctic Canada (GSC 85851).

BIOCHRONOLOGY OF TRIASSIC BIVALVES

Other Late Triassic bivalves of notable biochronological significance With few exceptions, the uppermost stage of the Triassic is characterized not so much by the emblematic pelagic bivalves, but by level bottom bivalve assemblages with long-ranging species. A few notable species, however, have proven biochronologic value but are mostly relegated to the western Tethys or the Boreal realm of northeastern Russia. The distinctive pteriid bivalve Rhaetavicula contorta (Portlock) (Fig. 7c) is perhaps the most commonly recognized bivalve and key index of the Rhaetian in shallow-water facies of the western Tethys, southern Germany and across northwestern Europe. Golebioski (1990) provided a most useful summary of the facies, geographical, and temporal distribution of R. contorta (Portlock). Based on recently published high-precision chemostratigraphic correlations between the NW European and Tethys sections (see Hesselbo et al. 2002; Kuerschner et al. 2007), and accepting the base of the Rhaetian to be the first occurrence of the conodont Misikella posternsteni (Krystyn et al. 2007b) which is present in the basal Hochalm Member of the Ko¨ssen Formation (Kozur 1996), all true R. contorta (Portlock) are now interpreted to be confined to the Rhaetian. Other claims of R. contorta in the eastern Tethys and circum-Pacific (e.g. Healey 1908; Muller & Ferguson 1939) may indeed be Rhaetian, but have yet to receive satisfactory study to warrant assignment to R. contorta (Portlock). In the Boreal realm and in far-east Russia in particular, the Rhaetian is characterized by the pectinid Tosapecten. Although Tosapecten sensu stricto appears as early as the Carnian of Japan and other east-Asian, western Panthalassa, and Boreal regions, other representatives of the family group such as Janopecten and Nevadapecten occur in the Middle Triassic (see Arkhipov & Truschelev 1980; Waller in Waller & Stanley 2005). Tosapecten efimovae Polubotko (Fig. 7a) has been a key index for the Rhaetian strata of Boreal far-east Russia (e.g. Polubotko & Repin 1990). This zone encompassing the entire Rhaetian can be subdivided into a lower Camptonectes nanus subzone and an upper Tosapecten efimovae subzone (Polubotko & Repin 1990; Dagys & Dagys 1994). Zakharov et al. (1997) seem to have inverted the Tosapecten effimove and Camptonectes nanus subzones.

Conclusions The reorganization of marine benthic communities following the end-Permian extinction marked the onset of an essentially modern fauna dominated

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by bivalve molluscs during Triassic time. The radiation of bivalves during the Triassic coincided with the origination and proliferation of many pelagic ‘flat clam’ species, which, because of their high evolutionary turnover and widespread distribution, makes them one of the best macrofossils for biochronological studies. Based on more than a century of work on pelagic bivalves of the Triassic in conjunction with detailed taxonomic study, high-resolution stratigraphic sampling and integration with more traditional biostratigraphic schemes using ammonoids and conodonts, the Triassic Period can been temporally subdivided into more than 30 discrete bivalve zones with a mean duration of typically less than 2 Ma. These bivalve zones can be identified across most of the marine Triassic with significant occurrences in the Tethyan Panthalassa and Boreal realms. Still, there remain many uncertainties in taxonomy of the zonal bivalves and the temporal limits of the established zones. I would like to thank the numerous colleagues who have made their collections available to me. Funding for aspects of this research were made available through the National Science Foundation (EAR–9706040 and EAR– 0518605) and the donors to the Petroleum Research Fund of the American Chemical Society. I. Polubotko, J. Grant-Mackie and Yin Hong-Fu made available photographs and additional information on Triassic bivalves which greatly aided in the preparation of this manuscript. S. Damborenea, N. Silberling, D. Taylor and B. Waterhouse provided helpful reviews that substantially improved the manuscript. This is a publication of IGCP Project 467, Triassic Time and Correlations.

Appendix 1. Taxonomic Notes Family Pterinopectinidae: Claraia sensu lato Claraia Bittner is a cosmopolitan pectinoid genus originating in the Permian but achieving peak diversity in the Early Triassic. Stemming back to Bittner (1901), it was recognized as a key index of the Early Triassic, and since that time, there have been various opinions as to the familial affinity of Claraia. Most recent workers now consider them to be pectinoids aligned with the Pterinopectinidae (e.g. Zhang 1980; Gavrilova 1995, 1996; Newell & Boyd 1995; Waterhouse 2000, 2008). Following recent workers (e.g. Newell & Boyd 1995; Yang et al. 2001; He et al. 2007) Pseudoclaraia Zhang, and Claraioides Fang are considered to be junior synonyms for Claraia. Other newly proposed clariid genera are based on either surfical morphology (e.g. Rugiclaraia Waterhouse) or represented by poorly known endemic genus-group taxa (e.g. Periclaraia Li & Ding and Epiclaraia Gavrilova) and can be included within current concepts of Claraia. Of the approximately 75 species names for Claraia, 30 or so are considered valid.

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Family Monotidae: Monotis sensu lato and Eomonotis A diverse group of which several genus-level taxa have biochronological significance in mostly Upper Triassic strata Eomonotis Grant-Mackie and Monotis Bronn sensu lato (here including the subgenera Monotis Bronn, Entomonotis Marwick, Pacimonotis Grant-Mackie & Silberling and Maorimonotis Grant-Mackie) represent a coherent clade arising from a likely Meleagrinella ancestor during the Early or Middle Norian. Waller (in Waller & Stanley 2005) considered the Monotioidae to be a superfamily within the Pectinoida and included the families Asoellidae, Monotidae and Oxytomidae. Waterhouse (2001, 2008) relegated the Monotidae to suborder rank and suggested they were derived from a eurydesmatoid ancestor. Although it is acceptable practice to subdivide this species group into several subgenera and/or species groups that may have some phylogenetic significance (Ichikawa 1958; Grant-Mackie 1978; Grant-Mackie & Silberling 1990) none of the proposed genus-groups have been adequately tested with a rigorous cladistic analysis. Of these proposed names, this paper follows the lead of Tozer (1980) in recognizing two broadly construed genus-group taxa: Monotis and Eomonotis.

Family Posidoniidae: Peribositria and Ellesmerella Waller (in Waller & Stanley 2005) elevated Posidoniodea to superfamily and included the families Posidoniidae and Halobiidae. Included in this clade are five genus-group taxa including the posidoniid Peribositria Kurushin & Trushchelev and the halobiid genera Enteropleura Kittl, Daonella Mojsisovics, Aparimella Campbell, and Halobia Bronn. As noted by several workers (e.g. Campbell 1994), the traditional classifications of the Posidonidae (sensu Cox & Newell 1969) contain an artificial grouping of several thin-shelled taxa that are now considered to belong to several different clades. Waterhouse (2008) placed some of these taxa within the Pterinopectinidae along with Claraia (e.g. Halobia, Aparimella, Daonella and some Enteropleura) and some (including true Jurassic Bositra) with several Enteropleura into his new genus Wallerobia (order Ostreoida). A more conservative approach used herein considers all to be pterioids. The Early and Middle Triassic contains several flat clams which have global biochronologic utility and have been variously placed in Posidonia Bronn, Peribositria Kurushin & Trushchelev, Bositra de Gregorio, and even Daonella Mojsisovics (e.g. Kurushin 1980; Waller in Waller & Stanley 2005). Waller (in Waller & Stanley 2005) placed most of the species attributed to Posidonia and Peribositria into an expanded view of Bositra. Because of some outstanding issues regarding morphological detail of the ligament systems of the Triassic forms, and, less so, because the large stratigraphic gap between

the Early and Middle Triassic forms and the Jurassic (Oxfordian) representatives, Peribositria is provisionally retained until further evidence can be brought to light. At present, about 15 species of Peribositria are known ranging from base of the Olenekian through the Middle Anisian. Although most other species historically attributed to Posidonia are rather poorly known, the radial-ribbed Posidonia aranea Tozer from the Olenekian of the Boreal province is sufficiently distinct from other so called Posidonia and Peribositria to warrant recognition. Waterhouse (2008) used this species as the basis for his new aulacomyellid genus Ellesmerella that he interpreted to be closely aligned to Bositra (etheripectinoid order Ostreoida). Although Ellesmerella appears to be valid (its sole species does not appear to belong in an existing genus-group taxon), it is perhaps better placed with the pterioid families Posidoniidae or Halobiidae (sensu Waller in Waller & Stanley 2005).

Family Halobiidae: Enteropleura, Daonella, Aparimella and Halobia Enteropleura Kittl, here considered to belong to the Halobiidae, is an interesting genus narrowly restricted to the late Middle Anisian. Like other Triassic flat clams, there is some controversy regarding its affinity, but most who have addressed this matter consider them to be the stem group for the Halobiiidae derived from the Posidoniidae (Encheva 1978; Campbell 1994; Waller in Waller & Stanley 2005; Chen & Stiller 2007). Departing from the generally accepted view, however, is Waterhouse (2008) who placed some Enteropleura (including the type species Enteropleura gumbelli Mojsisovics) with etheripectiniids in the Ostreoida and others, including his new Wallerobia (including Enteropleura jenksi Hopkin & McRoberts) within his new family Bositridae. Approximately five species of Enteropleura are known from Europe, Asia, and North America (Chen & Stiller 2007). A complex evolutionary transition from Daonella Mojsisovics to Halobia Bronn (either directly or via Aparimella Campbell) occurred near the base of the Carnian. For example, Daonella elegans McLearn passes into Halobia gr. H. zitteli Lindstro¨m within beds that containing both of the ammonoids Trachyceras and Frankites. At about the same time Daonella nitainae McLearn is found below Halobia daonellaformis McRoberts from similar aged strata in British Columbia. The taxonomy of species groups, subgenera and genera here included within Halobia sensu lato are provided by Kittl (1912), Encheva (1978), Gruber (1976), Polubotko (1984, 1988), Campbell (1994) and McRoberts (1997). Although there exist more that 300 available species names for Halobia, slightly more than 30 species of Halobia are likely valid with significant occurrences in Alpine-Mediterranean (Austria, Italy, Balkans) and circum-Pacific (Canada).

BIOCHRONOLOGY OF TRIASSIC BIVALVES

Repository abbreviations Repositories for illustrated specimens as follows: GSC, Geological Survey of Canada, Ottawa; NHMW, Naturhistorisches Museum, Vienna; PIMZ, Pala¨ontologischen Instituts und Museum, University of Zurich; SM Sedgwick Museum, Department of Earth Sciences, Cambridge University; SUCP, SUNY Cortland Paleontology Collections, Cortland, New York; TsNIGR, Central Geological Research Museum, St. Petersburg; UNIMI, Dipartimento di Scienze della Terra ‘Ardito Desio’, Univeristy of Milano; USNM, National Museum of Natural History, Smithsonian Institution, Washington, D.C.

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Triassic ammonoid biostratigraphy: an overview MARCO BALINI1*, SPENCER G. LUCAS2, JAMES F. JENKS3 & JUSTIN A. SPIELMANN2 1

Dipartimento di Scienze della Terra “A. Desio”, Via Mangiagalli 34, 20133, Milano, Italy

2

New Mexico Museum of Natural History and Science, 1801 Mountain Rd. NW, Albuquerque, NM 87104-1375, USA 3

1134 Johnson Ridge Lane, West Jordan, UT 84084, USA

*Corresponding author (e-mail: [email protected]) Abstract: The Triassic chronostratigraphic scale was built on two centuries of research on ammonoid biostratigraphy and biochronology. Two Triassic stage bases and all of the Triassic substages are currently defined by ammonoid bioevents. The study of Triassic ammonoids began during the late 1700s, and in 1895, Edmund von Mojsisovics, Wilhelm Waagen and Carl Diener published an essentially complete Triassic chronostratigraphic scale based on ammonoid biostratigraphy. This scale introduced many of the Triassic stage and substage names still used today, and all terminology of stages and substages subsequently introduced has been based on ammonoid biostratigraphy. Early Triassic ammonoids show a trend from cosmopolitanism (Induan) to latitudinal differentiation (Olenekian), and the four Lower Triassic substage (Griesbachian, Dinerian, Smithian and Spathian) boundaries are globally correlated by widespread ammonoid biotic events. Middle Triassic ammonoids have provinciality similar to that of the Olenekian and provide a basis for recognizing six Middle Triassic substages. Late Triassic ammonoids provide a basis for recognizing three stages divided into five substages. The main uncertainty for the future of Triassic ammonoid biostratigraphy is not the decline of the ammonoids as a tool for dating and correlation of Triassic strata but, rather, the dramatic decrease in the number of specialists, due to the lack of replacement of experienced palaeontologists who started their activity in the 1950s and 1960s.

The Triassic Period was a special time in the history of ammonoids (i.e. the representatives of the subclass Ammonoidea). The group barely survived the Permian/Triassic extinction and then it was very nearly destroyed again by the end-Triassic extinction, when only the precursors of Psiloceras survived. However, the time between these two dramatic events was one of the most interesting in the evolution of the Ammonoidea, and the group provided some of its best evolutionary performances, with an impressive succession of radiations and crises, including the first experiment in heteromorphic coiling. The resulting Triassic ammonoid record consists of three orders, about 80 families, about 700 valid genera and an estimated 5000 valid species. A few Triassic ammonoid families (e.g. Sageceratidae, Sturiidae, Gymnitidae, Ptychitidae, Isculitidae, Cladiscitidae, Arcestidae, Sphingitidae and Joannitidae) consist of smooth, relatively longranging forms that are informally defined as ‘leiostraca’, whereas the majority of families belong to the group ‘trachyostraca’, which includes the ornamented, fast-developing and short-ranging forms. The first description of Triassic ammonoids was in the late 1700s, and within a century of that, ammonoid biostratigraphy became the basic building block of the Triassic chronostratigraphic scale.

Indeed, the Standard Global Chronostratigraphic Scale of the Triassic (Fig. 1) was, until the 1990s, wholly based on ammonoid biostratigraphy. Today, ammonoids still provide a strong basis for Triassic marine correlations and are used to define some of the Triassic stages and all of the substages (see Lucas 2010). Given their central role in the Triassic chronostratigraphic scale, we provide a review of the history and current status of Triassic ammonoid biostratigraphy.

History of Study Tozer (1984) provided a detailed history of the use of ammonoids in the development of the Triassic chronostratigraphic scale. Additional discussion of aspects of this history can be found in Zittel (1901), Silberling & Tozer (1968) and Lucas (2010). Here, we provide a brief historical review emphasizing the development of the ammonoid biostratigraphic scale.

Beginnings After the first description of a Triassic ammonoid (‘Ammonite nodosa’) by the French naturalist Jean

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 221– 262. DOI: 10.1144/SP334.10 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. The Triassic chronostratigraphic scale.

Guillaume Bruguie`re (1750–1799), the first 60 years of the 1800s saw numerous descriptions of Triassic ammonoids from the Alpine regions of western Europe (Tozer 1984). This work involved extending the ‘Trias formation’ of Alberti (1834) from its largely nonmarine ‘type section’ in southwestern Germany into the marine strata of the Alps. Key to this extension was correlation of the ceratite record of the German Muschelkalk (the marine middle portion of the type Triassic) to the Alps, which resulted in recognition of an ‘Alpine Muschelkalk’ (e.g. Hauer 1850).

Mojsisovics’ contribution The key figure in the subsequent development of Alpine Triassic ammonoid biostratigraphy was Austrian palaeontologist Edmund von Mojsisovics (1839–1907), whose 40-year career began in the 1860s (Rosenberg 1958; Tozer 1984). Mojsisovics worked as a palaeontologist at the Geological Survey of Austria (Geologische Reichsanstalt), where he

quickly became the world authority on Triassic ammonoids. He collected many ammonoids, especially from the Alps, but he also studied all the previous collections as well as several new collections sent to him from locations as far-flung as the Olenek River in eastern Siberia, Japan and the Himalayas (e.g. Mojsisovics 1869, 1874, 1882a, b, 1888a, b, 1893, 1896, 1902). Mojsisovics published detailed and lavishly illustrated monographs (Fig. 2), describing more than 1000 new species (about 20% of the Triassic ammonoid species so far described) and 111 new genera. In addition to this impressive contribution to the understanding of Triassic ammonoids, Mojsisovics was also the first to establish a biostratigraphic scale of the Triassic. To do so, he applied the same approach used by Albert Oppel (1831– 1865) to recognize a succession of ammonoid zones in the Jurassic of central Europe (Oppel 1856–1858), applying it to the Triassic successions of the ‘Juvavische Triasprovinz’ (Northern Alps) and to the ‘Mediterranen Triasprovinz’ (Northern and Southern Alps, Hungary and Dinarids). However, the definition of the succession of Triassic ammonoid zones took Mojsisovics considerably more time when compared to Oppel’s contribution to the Jurassic, and required several intermediate steps (e.g. Mojsisovics 1873, 1879, 1882b, 1893), mostly because of the relatively complex tectonic setting of the Triassic successions in the Alps combined with complex facies changes resulting from carbonate platformbasin transitions and condensed successions (also see Tozer 1971, 1984). One crucial step was in 1882, when Mojsisovics published a comprehensive study of the ‘mediterranen Triasprovinz’ for which he provided an extremely clear and well-documented description of ammonoid zones based on faunas recognized in several facies and localities (Mojsisovics 1882b; Fig. 3). This scale, covering the interval between the Early Triassic and the earliest Late Triassic, combined a new taxonomy (several index species were formalized in the same monograph) and description of faunas within a framework of correlations between localities and different facies. The components of this structure were closely linked and this provided, at least at that time, some certainty to Mojsisovics’ biostratigraphy. The ammonoid zones recognized by Mojsisovics are very typical Oppel zones that were recognized by grouping together all taxa found in the same facies into a single fauna, then by comparing faunas from different sites/facies. Faunas with some common elements were regarded to be coeval, and then included in the same ammonoid zone. The zones defined in such a way have a biostratigraphic basis but also a chronostratigraphic significance and a

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Fig. 2. Plate 9 of Mojsisovics’ (1882b) monograph on the ammonoids of the Triassic ‘Mediterranean province’. Figures 1– 3 are ‘Meekoceras reuttense’, figure 4 is ‘Ceratites beyrichi’, figures 5– 6 are ‘Ceratites elegans’, figure 7 is ‘Ceratites rothi’ and figure 8 is ‘Ceratites boeckhi’.

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Fig. 3. The summary of Mojsisovics’ (1882b) biostratigraphic scale based on localities from the Southern and Northern Alps, Hungary and the Dinarids, with a minor contribution from southern Russia for the Lower Triassic. The C. binodosus and C. trinodosus zones were based on the most complex correlation of facies and localities. In the following 100 years they became the most controversial part of Mojsisovics’ scale.

very high potential for correlation. However, later in the twentieth century it became clear that the accuracy of these correlations was sometimes relatively low. The long-lasting efforts by Mojsisovics to emulate for the Triassic System what Oppel did for the Jurassic culminated with the 1895 contribution co-authored with Carl Diener (1862–1928) and Wilhelm Heinrich Waagen (1841–1900) (Mojsisovics et al. 1895). In this paper, the authors integrated the results of research carried out by Mojsisovics and Diener, mostly on the Middle and

Upper Triassic ammonoid zones of the Alps (primarily Austria, Italy and Bosnia) with Waagen and Diener’s work on the Lower Triassic of the Salt Range (Pakistan) and the Lower and Middle Triassic of the Himalayas (India). This integrated succession of ammonoid zones (Fig. 4) was used to define Substages, Stages and Series for the first complete chronostratigraphic scale of the Triassic System (see Lucas 2010), which had great influence on the development of the stratigraphy of the Triassic during the twentieth century. Several Substages and Stages introduced by Mojsisovics et al. (1895)

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Fig. 4. The Triassic ammonoid-based bio-chronostratigraphic scale of Mojsisovics et al. (1895).

are still commonly used in the present version of the Standard Global Chronostratigraphic Scale for the Triassic (cf. Ogg 2004a) if not formally defined.

From Mojsisovics to Spath The ammonoid bio-chronostratigraphic scale of Mojsisovics, even in his last version, had some weak points due to the incorrect stratigraphic position of some zones and the difficulty of separating other zones due to faunal similarities and ambiguous taxonomic separation of their index species, which would have required tens of years to be resolved. The incorrect reconstruction of the zonal succession affected especially the Carnian– Norian (see Tozer 1984 for details) and the Ladinian Dinarites avisianus Zone, which was located above the Protrachyceras curionii Zone in 1895 (Mojsisovics et al. 1895) and later moved to below it by Assereto (1969).

The faunal similarity, combined with ambiguous separation of Ceratites binodosus and C. trinodosus, deeply affected the development of Anisian bio- and chronostratigraphy, until Assereto (1971) clearly demonstrated the separation of the two faunas. At the beginning of the twentieth century another problem was soon emphasized. The known ammonoid record was recognized as being discontinuous and incomplete, and some authors integrated the ammonoid zones with bivalve, brachiopod and even crinoid zones, filling the gaps between ammonoid zones, or replacing them. This approach was employed by Arthaber (1905), who was the first to suggest a threefold subdivision of the marine Triassic, and by Pia (1930), who coined some substages that afterwards became very popular. A totally different solution to the subdivision of the Triassic was provided by British Museum palaeontologist Leonard Frank Spath (1888–1957), who emphasized

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the usefulness of ammonoids by subdividing (Spath 1934) the entire Triassic into ages, named for the most important ammonoid families (e.g. Beyrichitan and Paraceratitan for what was regarded as Anisian by Arthaber in 1905). Spath’s ‘ages’ were not conceived as hierarchically equivalent to stages, but rather as a rank equivalent to substages (Spath 1934, tables IV and V), further divided into zones based on faunas, almost equivalent to Mojsisovics’ Oppel zones. Spath’s approach was in part influenced by Williams (1901) and represents the first application of a biochronologic scheme to the Triassic by means of ammonoids.

The North Americans take the lead The middle of the twentieth century saw a historical change in the leadership of Triassic biostratigraphic studies. For almost a century, the ammonoid biostratigraphy of the Triassic was developed in the Germanic basin and in the Tethyan realm, whereas at the beginning of the 1960s the North American specialists E. T. Tozer and N. J. Silberling advocated the North American ammonoid scale as the best and the most complete in the world. The studies of Triassic ammonoids in North America actually started much earlier, in the nineteenth century, with the pioneering work of Alpheus Hyatt (1838–1902) and James Perrin Smith (1864– 1931). But, one peculiar feature that made the North American tradition so important is the continuity of the school, not affected by the disasters of WWI and WWII, as were the European schools. Hyatt and Smith published important monographs on the taxonomy of the new ammonoid faunas discovered in the western U.S. (Hyatt & Smith 1905; Smith 1914, 1927, 1932), but for the biostratigraphy they tried to apply the Tethyan scale to the North American successions (also see Lucas 2010). Their efforts were continued by F. H. McLearn in Canada and S. W. Muller (a student of Smith) in the U.S., who discovered several ammonoid localities between the 1930s and 1950s, and then by E. T. Tozer and N. J. Silberling (a student of Muller), who were very active from the 1950s to the 1990s. The first presentation of the North American biostratigraphic scale, totally based on ammonoids and documenting the complete chronostratigraphic subdivision of the Triassic System, was in 1967– 1968. Tozer (1967) first described a biostratigraphic scale for the Canadian Triassic, consisting of 31 zones. Silberling & Tozer (1968) presented a biostratigraphic correlation chart of the most important marine successions of Nevada, Idaho, California, Oregon, British Columbia, Alaska and Arctic Canada. The Canadian scale was extremely innovative, and its merits included: (1) ammonoid documentation from the basal Triassic to the

Rhaetian; (2) good control and documentation of the superposition of the zones, which was often poor in Mojsisovics’ scales; (3) clear indication of type localities of the zones as well as of their index species and faunal composition; (4) clear illustration of the type of biozone used in the scale, conceived of as “bodies of rock characterized by an assemblage of ammonoids of which one characteristic species is chosen as index” (Tozer 1967: p. 10); and (5) high biostratigraphic resolution, even if the ammonoid record was recognized as discontinuous. Further refinements of this scale (Tozer 1971, 1974, 1978, 1981b, 1984) lead to the final version (Tozer 1994) consisting of 37 zones (11 for the Lower, 12 for the Middle and 14 for the Upper Triassic), with 11 of them divided into a total number of 28 subzones.

New developments in the rest of the world The biostratigraphic studies of the Tethyan, Boreal and Germanic successions restarted in the 1960s and were driven by the stimulating results achieved in North America. During the 1970s, the ‘new wave’ of research was also fueled by the very successful IGCP (International Geological Correlation Programme) 4 ‘Triassic of the Tethys Realm’ and by the beginning of the activities of the Subcommission on Triassic Stratigraphy (International Commission on Stratigraphy; see Tozer 1983). The new researchers had to face several difficulties such as the revival of the old schools and the difficulties in sampling old localities impoverished or even destroyed by tens of years of collecting without detailed stratigraphic provenance. Many contributors provided new data, and they will be mentioned in the following sections of this overview (see Zapfe 1983 for a summary of the achievements of IGCP 4). Among them, it is worth introducing H. Rieber, L. Krystyn, and A. S. Dagys, whose contributions have been very significant. The difficulties of restarting research on old sites are well exemplified by the research and career of H. Rieber (Zu¨rich). He started his activities on Triassic ammonoids during the late 1960s (Rieber 1967), with a very accurate bed-by-bed study of the ammonoids faunas of the Besano Formation (Grentzbitumenzone sensu Swiss authors; Mt. S. Giorgio area, Southern Alps, southern Switzerland), erroneously included by Mojsisovics in the Ceratites trinodosus Zone (lithofacies c: ‘westlombardischen dolomite:’ Fig. 3; Rieber 1973). Afterwards, he extended his investigations to Lombardy and the Dolomites (Southern Alps, Italy), where he studied several classic and new sites in the Buchenstein Formation, recording faunas of the Trachyceras reitzi Zone sensu Mojsisovics (1882b), renamed as the Protrachyceras curionii

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Zone in Mojsisovics et al. (1895). It took him about 30 years to complete the revision of uppermost Anisian –Lower Ladinian ammonoid biostratigraphy of the Southern Alps (Brack & Rieber 1986, 1993), which allowed him to demonstrate that the true faunal succession consists of three ammonoid zones and that Reitziites reitzi and Eoprotrachyceras curionii are index ammonoids of different zones (Fig. 5). Rieber’s new high resolution biostratigraphic correlation chart provided a crucial contribution to the definition of the GSSP of the Ladinian Stage at the top of the ‘Chiesense groove’ at the Bagolino section B, at the FAD of the first species of Eoprotrachyceras, namely E. curionii (Brack et al. 2005). The contribution of L. Krystyn (Vienna) was innovative in a different way. He restarted investigations of the Upper Triassic condensed facies of the Northern Alps studied by Mojsisovics (Krystyn & Schlager 1971; Krystyn et al. 1971a, b; Krystyn 1973, 1978, 1980). He also worked in the Himalayas (Krystyn 1982) and extended his investigations down to the Lower and Middle Triassic (e.g. Krystyn 1974, 1983). Krystyn’s contribution to the revision of Triassic biostratigraphy is both unique and important in that it integrates the use of conodonts and ammonoids. His approach to stratigraphic research consists of documenting the simultaneous occurrences of ammonoids and conodonts, but keeping the ammonoid and conodont biostratigraphic scales separate (e.g. Krystyn 1983). This type of integrated approach was applied much later in the 1990s in North America (e.g. Orchard & Tozer 1997a, b). No discussion of prominent contemporary Triassic ammonoid workers is complete without mentioning the eclectic career of A. S. Dagys (1933–2000), whose reputation for an open-minded approach to national and international cooperation is well known. Although of Lithuanian nationality, Dagys spent the bulk of his career based in Novosibirsk, Russia, where his initial research in the early 1960s involved Mesozoic brachiopods and Triassic brachiopods, in particular, for which he became a world authority. In the late 1970s, he began working on Triassic ammonoids and he also eventually became one of the most prominent specialists in the world for this group. His numerous and varied works include studies of ammonoid systematics, palaeobiogeography, palaeobiology and ecology, biostratigraphy, intraspecific variation and the phylogeny of Boreal Triassic ammonoids. His friendly, cooperative manner helped him to establish an unusually large network of personal contacts with specialists from all over the world in spite of the practical difficulties of working within the ‘old Soviet’ system. His collaboration with E. T. Tozer and W. Weitschat was especially fruitful and

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culminated in highly precise Boreal realm correlation charts for Siberia, Arctic Canada and Svalbard (Dagys & Tozer 1989; Weitschat & Dagys 1989; Dagys & Weitschat 1993).

Modern biostratigraphy Triassic ammonoid biostratigraphy during the last 30 years has undergone the first major advance since Mojsisovics’ time. This progress has involved both taxonomy, which always provides the basis for any use of fossils as biostratigraphic tools, and the crucial contribution of ammonoids to integrated biostratigraphy and chronostratigraphy. This is well revealed by several papers and monographs describing new faunas and revising traditional ammonoid groups. First and foremost of these is the outstanding description and revision of all the Canadian Triassic ammonoid faunas by Tozer (1994), which represents the summary of one life dedicated to Triassic ammonoids. Also quite notable is the classic study by Silberling & Nichols (1982) in which they utilized intensive bed-by-bed sampling in the Humboldt Range of Nevada to establish a basic ammonoid zonation for the early and middle Anisian, and a precise biostratigraphic ammonoid succession for the late Anisian. The knowledge of Anisian ammonoid faunas in North America was also greatly improved by the impressive series of subsequent works by Bucher and his students (Bucher 1988, 1989, 1992a, b, 1994; Monnet & Bucher 2005b). These works have produced a much expanded and refined series of high resolution ammonoid successions for the early, middle and earliest late Anisian, which in total represents the most complete Anisian ammonoid record in the world (Monnet & Bucher 2005b). Other members of this group are now working on a complex revision of the Lower Triassic faunas of the eastern Tethys (Bru¨hwiler et al. 2008; Brayard & Bucher 2008). In Germany, Urlichs (1980, 1987, 1997, 2009) revised the ceratitids of the Upper Muschelkak, and in the Tethyan realm, apart from the above mentioned contributions by Rieber, the Anisian – Lower Ladinian ammonoid faunas of Bakony have been described by Vo¨ro¨s (1998, 2003). Ammonoid faunas from the Anisian to the Lower Carnian of the Southern Alps have been described from basinal facies (Balini 1992a, b, 1997; Balini et al. 2000b; Mietto et al. 2008) to carbonate platforms (Fantini Sestini 1994, 1996; Manfrin et al. 2005). Uppermost Olenekian to Lower Anisian faunas were also described from Albania and Greece (Germani 1997; Mertmann & Jacobshagen 2003). New faunas were described from Siberia (Dagys & Ermakova 1990; Dagys 2001) and from far

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Fig. 5. Integrated range chart showing the distribution of the short-ranging ammonoid genera and of the conodont species and subspecies at the Bagolino section (Southern Alps), where the GSSP of the Ladinian Stage has been defined (Brack et al. 2005). The long-ranging ‘leiostraca’ ammonoids Proarcestes and Flexoptyhites have not been plotted. The observed range of ammonoid genera is usually restricted to a part of an ammonoid zone and rarely covers more than one zone. The distribution of ammonoid species usually is restricted to one to four fossil-bearing levels. To provide an example, the distribution of the five species of Kellnerites is plotted. Key for abbreviations: K.hal., Kellnerites halilucensis; K. bosn., K. bosnensis; K. fiss., Kellnerites fissicostatus. Open circles: occurrence of ammonoids in correlated section of eastern Lombardy; the range of conodonts with a dotted line is from correlated sections from eastern Lombardy and Dolomites. The figure is redrawn and simplified from Brack et al. (2005, fig. 6).

eastern Russia (Zakharov 1978, 1997; Shigeta et al. 2009). The refinement of taxonomy in the context of bed-by-bed collections has confirmed the high power of stratigraphic resolution of ammonoids,

so it has enhanced their significance to chronostratigraphy. This field of stratigraphy has also experienced a major revolution in the last 20 years because of the evolution of the boundary stratotype concept (e.g. Hedberg 1976; Salvador 1994;

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Remane et al. 1996) and research aimed at the definition of chronostratigraphic units by GSSPs (Global Stratotype Section and Point), carried out within the Working Groups of the Subcommissions of the ICS (International Commission on Stratigraphy). The modern approach to the definition of the GSSP emphasizes the correlateability of (bio)events, and this is tested by a cross-comparison of bioevents of different fossil groups. This is accomplished by integrated stratigraphy, that is, by applying several biostratigraphic tools to the same sections and also by calibrating the correlations with other sections by means of magnetostratigraphy and stable isotope analysis. The role of ammonoid biostratigraphy in the definition of the GSSPs of the Stages of the Triassic has thus been confirmed and even enhanced. The comparison of ammonoid biostratigraphy to conodont, radiolarian, pelagic bivalve and palynomorph biostratigraphies demonstrates, without doubt, that the ammonoids are the taxonomic group with the highest power of temporal resolution within Triassic marine environments (Fig. 5). On the other hand, ammonoids are not as frequent as conodonts in marine successions, so the combination of the two tools represents the up-to-date biostratigraphic approach to modern chronostratigraphy (Fig. 5). Ammonoids are used to tie the conodont bioevents to the ammonoid bio-chronostratigraphic scale and to demonstrate their significance as time markers. Conodont bioevents are often preferred for the selection of the GSSP primary marker event, but this is not a strict rule. When there are no good, distinct or evolutionarily significant conodont bioevents, few complain about the choice of an ammonoid bioevent. It is not an accident that two out of three GSSPs ratified for the Triassic System

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in the Standard Global Chronostratigraphic Scale are defined by means of ammonoid bioevents. The FAD (First Appearance Datum) of Eoprotrachyceras curionii (Mojsisovics) is the marker event for the base of the Ladinian (Brack et al. 2005), and the FO (First Occurrence) of Daxatina canadensis is the marker event for the base of the Carnian (Mietto et al. 2007a, b; Gaetani 2009).

Triassic Ammonoid Record Discoveries of Triassic ammonoids in the early 1800s revealed an essentially global record early during their history of study (Tozer 1984). The most important records for bio-chronostratigraphic scale construction have been those in: (1) the Canadian Arctic Islands (especially Ellesmere and Axel Heiberg islands); (2) British Columbia and (3) Idaho, Utah and Nevada in the western American Cordillera; (4) Germanic basin; (5) Western Tethys (from the Alps to Turkey); (6) Transcaucasia (current border region of Azerbaijan and Iran); (7) Salt Range in Pakistan; (8) Himalayas; (9) southern China (Guangxi and Guizhou provinces); and (10) eastern Siberia (Fig. 6): 1. Canadian Arctic Islands: strata of the Blind Fjord Formation deposited in the Sverdrup basin yield the Lower Triassic ammonoid assemblages upon which Tozer (1965, 1967) based his Griesbachian, Dienerian, Smithian and Spathian stages. This is one of the most extensive Lower Triassic ammonoid records, and its significance is all the more remarkable given that all stage definitions are based on ammonoids occurring in similar facies within a single, largely undisturbed formation in the same depositional basin (Tozer 1994).

Fig. 6. Triassic world (artwork by Matt Celeskey) showing important Triassic ammonoid collecting areas.

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Overlying strata in the Sverdrup basin yield a less dense succession of Middle and Upper Triassic ammonoid assemblages (Tozer 1994). 2. British Columbia: in the western Canadian Cordillera, important successions of Lower, Middle and Upper Triassic ammonoid assemblages are known from the Canadian Rocky Mountains and adjacent foothills, particularly from the Toad and Sulphur Mountain formations (Lower and Middle Triassic) and Ludington and Pardonet formations (Upper Triassic) (Tozer 1994). In some midpalaeolatitude northeastern British Columbia localities, many boreal and palaeoequatorial Tethyan ammonoids occur together, thus facilitating correlation between the Arctic and the Tethys (Tozer 1994). Aside from the Griesbachian, the Lower Triassic ammonoid record nicely complements that of the Arctic Sverdrup basin (Tozer 1994). The exquisitely well-preserved Upper Triassic ammonoid fauna of the Pardonet Formation contributed immensely to deciphering the correct sequence of Upper Carnian/Lower and Middle Norian ammonoid zones (Tozer 1994). Indeed, the section of the Pardonet Formation at Black Bear Ridge is one of the leading GSSP candidates for the Carnian/Norian boundary (Orchard 2007). On the British Columbia coast, important assemblages of uppermost Triassic and lowermost Jurassic ammonoids are present on Vancouver Island and in the Queen Charlotte Islands (e.g. Tozer 1994; Tipper et al. 1994; Longridge et al. 2007). 3. Idaho, Utah and Nevada: strata of the Thaynes Group in Idaho and Utah yield important Lower Triassic ammonoid successions, including the Smithian Meekoceras and Anasibirites assemblages as well as the Spathian Columbites and Prohungarites assemblages, which were first developed by Smith (1932) and Kummel (1969), and more recently studied by Guex et al. (2005a, b), Jenks (2007) and Lucas et al. (2007a), among others. Additionally, a well-preserved early Dienerian ammonoid assemblage is known from the Candelaria Formation of southwestern Nevada (Muller & Ferguson 1939; Silberling & Tozer 1968). Preliminary results from ongoing work on multiple Smithian ammonoid levels contained in two west-central Utah sections indicate a much expanded Smithian ammonoid succession than the other more-or-less condensed sections of the western states (Stephen et al. 2008; Bylund et al. 2009). Recent work by Guex and collaborators reveals that the Spathian ammonoid succession of Idaho and Nevada is nearly complete and contains significantly more faunal levels than previously recognized. The Nevada record of Triassic ammonoids is most extensive for the Anisian and to a lesser extent the Ladinian, particularly from the Prida and Favret formations in northwestern Nevada (e.g. Silberling

& Nichols 1982; Bucher 1988, 1989, 1992a, b, 1994; Monnet & Bucher 2005b). Indeed, the zonation of the Nevada Anisian in all probability represents the most refined Triassic ammonoid succession known today (Monnet & Bucher 2005b). Nevada also includes significant Upper Triassic ammonoid assemblages from diverse localities (e.g. Silberling 1956, 1959; Silberling & Tozer 1968; Balini 2008), including an important record of uppermost Triassic and lowermost Jurassic ammonoids in the New York Canyon area in westcentral Nevada (e.g. Taylor et al. 1983; Guex et al. 2004; Lucas et al. 2007b). Other well-known Upper Triassic ammonoid assemblages from Nevada include the Lower Carnian Trachyceras desatoyense succession from the New Pass Range (Johnston 1941; Balini et al. 2007; Balini & Jenks 2007) and the Upper Carnian –Lower Norian Klamathites–Guembelites successions from West Union Canyon in the Shoshone Range (Silberling 1959). Important Lower and Upper Triassic ammonoid records are also known from California and include the Union Wash (Inyo Mountains) Smithian/Spathian assemblages (Smith 1932) and the Shasta County Upper Carnian Tropites welleri and T. dilleri assemblages, which dominated Smith’s 1927 Upper Triassic monograph (Silberling & Tozer 1968). 4. Germanic basin: the cradle of the Triassic yields ammonoids of Anisian to possibly earliest Carnian age that in the nineteenth century were crucial for the identification of marine Triassic sediments all over the world. The fossiliferous successions of the Lower and Upper Muschelkalk are exposed in southern Poland (Silesia) and centralsouthern Germany (Thuringia, Baden–Wu¨rttemberg). They were deposited in an epicontinental basin with limited connections to the Tethys ocean, in which the peculiar conditions lead to the development of endemic lineages of ammonoids (Urlichs 1980, 1987, 1997, 2009; Klug et al. 2005). Occasional migration (or drifting?) of Tethyan elements into the basin (Urlichs 1978; Urlichs & Mundlos 1985; Kaim & Niedzwiedzki 1999; Klug et al. 2005) provides elements to calibrate the biostratigraphic scale of the Muschelkalk with the Tethyan scale (e.g. Brack et al. 1999). 5. Western Tethys: the extensive ammonoid assemblages from the Alpine regions (especially Austria, Italy, Hungary and Bosnia) provided the basis for definition of most of the Middle and Upper Triassic stages and substages (see above). The Lower Triassic in the Alpine region generally consists of facies less than ideal for ammonoid preservation (such as the Werfen Formation), with the exception of the classic Spathian ‘Tirolites fauna’ (Mojsisovics 1882b; Kittl 1903, Kummel 1969). Spathian faunas are also documented in Hallstatt

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facies of Albania and Chios island in the Aegean Sea (Arthaber 1911; Renz & Renz 1948; Kummel 1969; Germani 1997; Mertmann & Jacobshagen 2003). Ammonoid assemblages from the Aegean Islands (particularly Chios) and coastal Turkey provided the basis for subdivision of the lower part of the Anisian into Aegean and Bithynian substages by Assereto (1974). In the Alpine region, carbonate facies (both limestone and dolomite) with extensive reef development yield most of the Middle and Upper Triassic ammonoid assemblages, including the Hallstatt facies, which contains exquisitely preserved ammonoids but suffers from problems of condensed fossil assemblages. Particularly significant ammonoid successions include those from basinal facies of Lombardy, the Dolomites and the Julian Alps in northern Italy (e.g. Brack & Rieber 1993; Mietto & Manfrin 1995a, b; Brack et al. 2005; Mietto et al. 2007a), the Northern Calcareous Alps of Austria and the Balaton highland of Hungary (Voro¨s 1987, 1998, 2003). The GSSPs of the Ladinian and Carnian stages are defined by ammonoid faunas at Bagolino (Lombardy) and Prati di Stuores/Stuores Wiesen (Dolomites). Other important GSSP candidate sections are Pizzo Mondello (Sicily, Italy) and Steinbergkogel in the Northern Alps (Austria). Pizzo Mondello is located in western Sicily, famous for the Upper Carnian– Lower Norian ammonoid faunas illustrated by Gemmellaro (1904), and is one of two candidate sections for the definition of the GSSP of the base of the Norian (Nicora et al. 2007; Balini et al. 2008). Steinbergkogel, located in the world famous Hallstatt region (Northern Alps, Austria), is the only GSSP candidate section for the base of the Rhaetian (Krystyn et al. 2007b; Krystyn 2008). 6.Trans-Caucasia:the Permian–Triassicboundary in the Dzhulfa (Julfa) region of the Iran –Azerbaijan borderland yields an important succession of ammonoid assemblages first described by Ruzhentsev & Sarycheva (1965). Several ammonoid faunas ranging in age from the Olenekian to Rhaetian are also known from the Crimea, northern Caucasus, Mount Bodgo and western Kazakhstan (Shevyrev 1968). Among them, the Olenekian faunas of northern Caucasus and west Kazakhstan are of special significance (Shevyrev 1968; Balini et al. 2000a). 7. Salt Range: Waagen’s work concerned with the Lower Triassic ammonoid assemblages in the Salt Range of western Pakistan provided the basis for subdivision of the Early Triassic as early as Mojsisovics et al. (1895). Kummel & Teichert (1970) presented a comprehensive study of this section, and Guex (1978) divided the Lower Triassic into Griesbachian, Nammalian (¼ Dienerian þ Smithian) and Spathian stages based on the ammonoid assemblages of what have long been called

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the (ascending order) Lower Ceratite limestone, Ceratite marls, Ceratite sandstone and Upper Ceratite limestone (though Kummel & Teichert [1966] proposed a formal stratigraphic nomenclature for these strata). Recent work in the Salt Range has resulted in a preliminary, but much expanded and refined, Smithian ammonoid succession (Bru¨hwiler et al. 2007). This work, carried out by scientists from the University of Zu¨rich, is ongoing and has been expanded to include the Dienerian and Spathian ammonoid assemblages. 8. Himalayas: Triassic ammonoid assemblages occur in the Himalayas from Kashmir to Spiti and were published early on by Oppel (1865) and Diener (1895, 1897, 1906, 1907, 1908, 1909, 1912). Most extensive are Lower and Middle Triassic ammonoid assemblages, but Upper Triassic ammonoids are also present. The section in the Mud (¼Muth auctorum) Valley of Spiti is the most impressive of the area for its spectacular outcrops (see Albertiana 35, front cover: http://www3.bio. uu.nl/palaeo/Albertiana/Albertiana01.htm) and fossil abundance. The ammonoid record is so rich and important that in recent years the section was investigated for the definition of two different GSSPs, namely for the base of the Olenekian and for the base of the Carnian. Originally described by Diener (1897) and Krafft & Diener (1909), Mud provides the best low-mid palaeolatitude ammonoid record in the world for the base of the Olenekian Stage and it is also rich in conodonts (Krystyn et al. 2007a). At present the section has been selected by the Induan/Olenekian boundary Working Group as the best GSSP section, and the LO (Lowest Occurrence) of conodont Neospathodus Waggeni is under examination as possible marker event. Sections in the Mud Valley were also studied for the rich Upper Ladinian to lowermost Carnian ammonoid-daonellid-conodont record (Balini et al. 1998, 2001; Krystyn et al. 2004), for which it is one of the three best sites in the world. 9. Southern China: an important, and growing record of Lower Triassic ammonoid assemblages is known from southern China, especially from Guangxi and Guizhou provinces (e.g. Bru¨hwiler et al. 2008; Brayard & Bucher 2008; Brayard et al. 2008). 10. Siberia: the most significant Siberian ammonoid records are those related to definition and characterization of the Olenekian Stage, which are in the eastern Verkhoyansk region, the lower reaches of the Olenek River and the Taimyr Peninsula (e.g. Kiparisova & Popov 1956, 1964; Dagys & Ermakova 1988, 1990, 1993; Dagys 1994, 1995, 1999; Ermakova 1999; Shevyrev 2002). The Spathian portion of the ammonoid record is exceptionally complete and more refined than that of Arctic Canada (Dagys 1994). Perhaps less well

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known, but still significant, is the comprehensive Induan Boreal ammonoid record from the eastern Verkhoyansk region (Dagys & Ermakova 1996). The ammonoid succession from this area correlates well with those from the Griesbachian and Dienerian substages of Arctic Canada, and its zonation is slightly more refined (Dagys & Weitschat 1993; Dagys & Ermakova 1996). Another significant ammonoid record is that provided by southern Primorye, in eastern Siberia. This mid-palaeolatitude region yields well-preserved ammonoids from a relatively complete biostratigraphic sequence (Griesbachian to Spathian) and includes taxa common to Boreal, Tethyan and the eastern Panthalassic realms (Shigeta et al. 2009). 11. Other regions: significant Triassic ammonoid records are also known from other diverse localities (for example, Spain, Israel, Afghanistan, Tibet, Timor, Svalbard, Greenland, Alaska, New Guinea, Peru, Mexico), but have not played as large a role in timescale subdivision as the regions discussed above.

Early Triassic Stages Two Early Triassic stages are now recognized by the STS – Induan and Olenekian – originally defined (and redefined) by Kiparisova & Popov (1956, 1964) based on ammonoid assemblages in IndiaPakistan and northeastern Siberia (Fig. 7). As Tozer (1978) observed, about 20 chronostratigraphic terms have been proposed to refer to all or part of the Lower Triassic, and all are based on ammonoid biostratigraphy. In 1991, the STS decided to use the term Induan as the name of the oldest Triassic stage (Visscher 1992), so that the Induan is now used on the current ICS timescale as the oldest Triassic stage (Ogg 2004a, b). Its base has been formally defined as the FAD of the conodont Hindeodus parvus at the Meishan section in southern China (Yin et al. 1996, 2001). There is also broad agreement that the Induan is equivalent to the upper Griesbachian and the Dienerian, which are ammonoid-based substages. In 1991, the STS decided to adopt the Olenekian as the younger stage of the Lower Triassic (Visscher 1992). Recent lengthy debate about a GSSP definition of the Olenekian base (Zakharov 1994; Chinese Triassic Working Group 2007; Krystyn et al. 2007c) may be resolved by the choice to use the LO of the conodont Neospathodus waageni as the defining criterion, with a likely GSSP located in the Mud section in India (Krystyn et al. 2007a). There is broad agreement that the Olenekian is equivalent to the Smithian and Spathian, which are ammonoid-based substages.

Fig. 7. Ammonoid zonation of the Induan and Olenekian Stages, as originally proposed by Kiparisova & Popov (1956).

Ammonoid provinces and biotic events Various schemes of Early Triassic ammonoid provinces have been proposed (e.g. Diener 1916; Tozer 1981b; Dagys 1988), and changes in provinciality during the Early Triassic are obvious (Brayard et al. 2007). Thus, the Griesbachian–Dienerian was a time interval of ammonoid cosmopolitanism, followed by a shift to the more endemic and latitudinally-restricted ammonoid assemblages of the Smithian and Spathian (Brayard et al. 2007). This is readily seen in the degree of similarity of ammonoid zones of the Griesbachian–Dienerian and the divergence in zonation during the Smithian– Spathian (Fig. 8). The palaeolatitudinal differences in Early Triassic ammonoid assemblages are revealed by a generally lower diversity in more northern assemblages when compared to more palaeoequatorial assemblages, and this has been seen as due to colder water (lower diversity) vs. warmer water temperatures (higher diversity) (Kummel 1973; Tozer 1981b, 1994). Similarity of low and high palaeolatitude ammonoid assemblages of Griesbachian and Dienerian age has been taken to indicate a lack of steep diversity gradients shortly after the end-Permian mass extinction (Brayard et al. 2006, 2007).

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Fig. 8. Lower Triassic ammonoid zones (based on Shevyrev 2002; Galfetti et al. 2007; and Shigeta et al. 2009).

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Substage divisions of the Lower Triassic have long been based on a succession of ammonoid biotic events: 1. The base of the Griesbachian (a Late Permian datum) is marked by the Lowest Occurrence of Otoceras (Fig. 9), an ammonoid genus found in both Tethyan and boreal ammonoid assemblages. The LO of Otoceras was long considered to be the base of the Triassic (Griesbach 1880; Mojsisovics et al. 1895), but the definition of the Induan base using the LO of the conodont Hindeodus parvus places the LO of Otoceras in the Late Permian (Orchard & Tozer 1997a, b). Indeed, the base of the Triassic is now within the Griesbachian, close to the LO of Ophiceras (Fig. 9), which is marked by a significant ammonoid event at the base of the upper Griesbachian (appearance of abundant ophiceratids). Thus, the base of the Triassic as now defined can be recognized by a significant ammonoid biotic event. 2. The base of the Dienerian is marked by the appearance of abundant meekoceratids, a globally recognizable event (e.g. Tozer 1974; Shevyrev 2001; Bru¨hwiler et al. 2008). 3. The base of the Smithian is marked by the appearance of many new ammonoid taxa, such as Flemingites, Kashmirites and Hedenstroemia. This was accompanied by formation of a pronounced latitudinal diversity of ammonoids (Brayard et al. 2006, 2007). 4. The base of the Spathian is also marked by the appearance of many new ammonoid taxa, notably the dinaritines, tirolitines and columbitids. The disappearance of typical Smithian genera (Anasibirites, Wasatchites, etc.) also marks the base of the Spathian. In other words, a major ammonoid extinction occurred at the end of the Smithian/ beginning of the Spathian and was followed by a rather rapid major evolutionary radiation (Tozer 1982; Galfetti et al. 2007). This event corresponds to a major perturbation of the carbon cycle that has been interpreted as a change from warm and equable global climate (Smithian) to a latitudinallydifferentiated global climate (Spathian) (Galfetti et al. 2007).

Ammonoid zones Zonal schemes for Lower Triassic ammonoids (Fig. 8) have been constructed primarily for: (1) the Canadian Arctic (Figs 9– 10); (2) the western United States (Fig. 11); (3) the Himalayas; (4) southern China; and (5) Siberia (northeastern and southern Primorye). Recent and ongoing work in Oman, the Salt Range, southern China, the Himalayas (Spiti) and southern Primorye has resulted in much more comprehensive and, more importantly, correlateable ammonoid successions (Bru¨hwiler

et al. 2007; Brayard & Bucher 2008; Krystyn et al. 2007a; Shigeta et al. 2009). In addition, work underway in the western United States will ultimately result in a much more refined Smithian and Spathian ammonoid record for this eastern Panthalassic region (Guex et al. 2005a, b; Stephen et al. 2008; Bylund et al. 2009).

Middle Triassic Stages The Middle Triassic has been divided into two stages – the Anisian and Ladinian – since early in the twentieth century, and this practice was sanctioned by two votes by the STS in 1984 and in 1991. Unlike the Lower Triassic, this part of the Triassic chronostratigraphic scale has been fairly stable – only two stages have been advocated for the Middle Triassic, and few competing (overlapping) stage terms have been proposed. The Anisian Stage encompasses four substages (or, in North America, three: Lower, Middle and Upper), and the Ladinian stage is divided into two substages (Fig. 1). Ammonoid zonation of the Anisian (Fig. 12) is particularly well developed in the western Tethys and in Nevada, USA (e.g. Bucher 1988, 1989, 1992a, 1994; Silberling & Nichols 1982; Brack & Rieber 1993; De Zanche et al. 1993, 1995; De Zanche & Gianolla 1995; Manfrin & Mietto 1995; Mietto & Manfrin 1995a; Brack et al. 2005; Monnet & Bucher 2005a, b; Monnet et al. 2008), whereas the Ladinian zonation (Fig. 13) is more refined in southern Europe (Brack & Rieber 1993; Mietto & Manfrin 1995a, b; Mietto et al. 2008) and Canada (Tozer 1994).

Ammonoid provinces and biotic events Middle Triassic ammonoids show provinciality that follows palaeolatitude, much as during the Olenekian, but this is also influenced by the maximum extension of the epicontinental seas around the Tethys. In the Aegean and Bithynian (Early to early Middle Anisian), Tethyan and Boreal ammonoid assemblages are readily distinguished, the first dominated by ussuritids, japonitids, achrochordiceratids and balatonitids, and the second mostly consisting of longobarditids, parapopanoceratids and danubitids (Tozer 1981b; Dagys 1988). The two assemblages, however, show some common genera, such as Paradanubites, Paracrochordiceras (Fig. 14), Leiophyllites, Japonites and Stenopopanoceras. The Pelsonian (late Middle Anisian) and the Illyrian (Late Anisian) initiated the maximum palaeobiogeographical differentiation of the Triassic ammonoids, concomitant with the

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Fig. 9. Selected index ammonoids of the Lower Triassic (Griesbachian) from the Canadian Arctic. Otoceras boreale, GSC 28236, phragmocone with about one quarter whorl of body chamber (a) lateral view, (b) aperture view. Otoceras concavum, GSC 18883 [paratype] (c) lateral view, (d) ventral view. Ophiceras commune, GSC 28056 [hypotype], phragmocone with about one quarter whorl of body chamber (e) lateral view, (f) ventral view, (g) aperture view. Bukkenites nanus, GSC 28075 [holotype] (h) lateral view, (i) aperture view. a– i, from Tozer, 1994, a–b, pl. 1, fig. 1a– b, c–d, pl. 1, fig. 3a –b, e–g, pl. 4, fig. 1a –c, h–i, pl. 8, fig. 2a –b. Scale bars 20 mm. Upper scale bar applies to a–b, middle left scale bar applies to c –d, middle right scale bar applies to h –i, lower scale bar applies to e–g. Figures a–i are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Canadian Triassic ammonoid faunas, GSC Bulletin 467).

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Fig. 10. Selected index ammonoids of the Lower Triassic (Dienerian) from the Canadian Arctic. Vavilovites sverdrupi, GSC 14275 [paratype] (a) lateral view, (b) sectional view. Proptychites candidus, GSC 28092, last half whorl appears to be body chamber but septa are not visible on phragmocone (c) lateral view, (d) aperture view. a– d, from Tozer, 1994, a– b, pl. 15, fig. 1a– b; c– d, pl. 12, fig. 5a– b. Scale bars 20 mm. Upper scale bar applies to a, lower left scale bar applies to c–d, lower right scale bar applies to b. Figures a–d are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Canadian Triassic ammonoid faunas, GSC Bulletin 467).

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Fig. 11. Selected index ammonoids of the Lower Triassic from Nevada and Idaho. Meekoceras gracilitatis (a) lateral view, (b) ventral view. Arctoceras tuberculatum (c) lateral view, (d) ventral view. Anasibirites kingianus (e) lateral view, (f ) ventral view. Columbites parisianus (g) lateral view, (h) ventral view. Prohungarites mckelvei (i) lateral view, ( j) ventral view. a– j, from Jenks et al. 2007, a–b, pl. 2E–F, c– d, pl. 1C –D, e– f, pl. 4C–D, g –h, pl. 7G– H, i –j, pl. 9D–E. Scale bars for a–b and e –j 20 mm; for c–d 40 mm.

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Fig. 12. Middle Triassic ammonoid zones [based on Mietto & Manfrin 1995a (Tethyan Summary) and Monnet & Bucher 2005a, b (Western Nevada to Balaton Highlands)].

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Fig. 13. Upper Triassic ammonoid zones (based primarily on Tozer 1994; Mietto & Manfrin 1995a, b and Krystyn et al. 2007b).

eustatically-driven transgression of the epicontinental seas into the Germanic basin (Germanic province) and on the northern continental shelf of Gondwana (Sephardic province). The most common form of the Pelsonian was Balatonites (Figs 14, 15), which colonized all the provinces except for the Boreal realm. During the Illyrian

(Late Anisian) and Fassanian (Early Ladinian), the Germanic and Sephardic faunas became quickly differentiated by the development of endemic forms and lineages consisting of divergent subgenera of ceratitids in the Muschelkalk (e.g. Urlichs & Mundlos 1985) and of Paraceratitoides-Gevanites in southern Israel (Parnes 1975, 1986).

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Fig. 14. Selected index ammonoids of the Tethyan Anisian. Paracrochordiceras asseretoi, paratype (a) lateral view. Nicomedites osmani syntype (b) lateral view, (c) drawing of sutural pattern. Aghdarbandites ismidicus, holotype (d) lateral view, (e) ventral view, (f ) drawing of sutural pattern. Balatonites balatonicus, holotype (g) lateral view, (h) ventral view. Paraceratites trinodosus lectotype (i) lateral view, (j) ventral view. Reitziites reitzi, specimen from Bagolino (k) lateral view. Reitziites reitzi syntype from Bakony (l) lateral view, (m) ventral view. a, from Fantini Sestini 1981, pl. 4, fig. 1; b –c, from Toula 1896, pl. 22, fig. 6a –b; d –f, from Arthaber 1914, pl. 12, fig. 6a– c; g– h, from Mojsisovics 1882b, pl. 4, fig. 2a –b; i– j, from Mojsisovics 1882b, pl. 8, fig 6a– b; k, from Gaetani (ed.) 1993 pl. 5, fig. 2. Scale bars 20 mm. Upper left scale bar applies to a, upper right scale bar applies to b –f, middle right scale bar applies to l– m, lower left scale bar applies to k, lower right scale bar applies to g –j. Figure a is reproduced with the permission of the Rivista Italiana di Paleontologia e Stratigrafia.

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Fig. 15. Selected index ammonoids of the Middle Triassic from Nevada. Acrochordiceras carolinae (a) lateral view, (b) ventral view. Nevadisculites taylori (c) lateral view, (d) ventral view. Eogymnotoceras thompsoni (e) lateral view, (f ) aperture view. Balatonites shoshonensis (g) lateral view, (h) ventral view. Rieppelites shevyrevi (i) lateral view, ( j) ventral view. Silberlingitoides cricki (k) lateral view, (l) ventral view. a–l, from Jenks et al. 2007, a–b, pl. 18G–H, c–d, pl. 19C–D, e–f, pl. 18K– L, g– h, 16H–I, i– j, pl. 23G–H, k– l, pl. 22E –F. Scale bars 20 mm.

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During the Illyrian, Tethyan faunas were dominated by paraceratitids together with beyrichitids and showed similarities with North America (Eastern Tethyan to ?Pacific province: see Tozer 1981b; Dagys 1988) at the subfamily level, but with several differences at the generic level. During the late Illyrian the paraceratitids were replaced by nevaditids (Fig. 16) in both North America and Tethys, leading to a significant increase of faunal similarity. Boreal assemblages lacked paraceratitids, but included some beyrichitids accompanying typical longobarditids, danubitids and parapopanoceratids (Dagys 1988). At the genus level, Frechites, Gymnotoceras, Longobardites and Ptychites are common elements shared by North America and the Boreal realm. The Ladinian is characterized by an increase in ammonoid cosmopolitanism at low-mid palaeolatitudes. From the Early to Late Ladinian (Fassanian –Longobardian), the faunal similarity between western Tethys and North America is impressive, with assemblages showing the same kind of structure and genera. The Boreal assemblages are poor and dominated by ‘leiostraca’ long-ranging ammonoids in the Early Ladinian (Tozer 1981b; Dagys 1988), whereas the Late Ladinian shows the development of extremely similar faunas from the Canadian Arctic to eastern Siberia, wholly consisting of nathorstitids (Nathorstites and Stolleyites) with some ‘leiostraca’ long-ranging forms (Tozer 1981b, 1994; Konstantinov 2008). Correlations with Tethyan successions are allowed by Nathorstites, which occurs also at mid-palaeolatitudes in northeastern British Columbia (Tozer 1981b, 1994). The substage subdivision of the Anisian and Ladinian was proposed based on ammonoid zones that sometimes were interpreted in different ways in the literature. As a consequence, the position and correlateability of the main faunal changes are occasionally a matter of discussion. 1. The base of the Aegean and of the Anisian is marked by a major ammonoid turnover. Most Spathian genera disappear at the beginning of the Anisian, and a variety of taxa (Paracrochordiceras, Japonites, Gymnites as well as danubitids, longobarditids and cladiscitids, among others) first appear during the latest Spathian or at the Anisian base (Tozer 1981a, b, 1984; Brayard et al. 2006). The proposed GSSP for the base of the Anisian (LO of the conodont Chiosella timorensis) is approximately equivalent to the LOs of the ammonoids Paracrochordiceras, Japonites and Aegeiceras (Gradinaru 2000; Gradinaru & Sobolev 2006; Orchard et al. 2007). The Anisian was a time of great ammonoid differentiation, witnessing the appearance of some important long-ranging ‘leiostraca’ ammonoid families such as Gymnitidae, Isculitidae, Ptychitidae and Arcestidae, some of which extend up into

the earliest Ladinian. It was also characterized by important radiations of members of the superfamily Ceratitaceae. 2. The Bithynian substage is characterized by a major radiation of Beyrichitinae (e.g. Nicomedites, Hollandites, Anagymnotoceras; Fig. 14), Ptychitidae, and Acrochordiceratidae (Assereto 1974; Bucher 1988, 1992a). 3. The base of the Pelsonian substage is characterized by the appearance of Balatonitidae and the LO of its most typical representative, Balatonites (Figs 14, 15), in association with Acrochordiceratidae. Balatonites is a remarkable guide fossil for the substage, because it occurs in the Muschelkalk, Tethys, Sephardic and Pacific palaeobioprovinces. However, the definition of the base of the Pelsonian is still a matter of debate. 4. The base of the Illyrian is marked by the radiation of Paraceratitinae (e.g. Paraceratites; Fig. 14) and the disappearance of Acrochordiceratidae and Balatonitidae. 5. Ammonoid diversity notably decreased at the beginning of the Fassanian due to the extinction of several short-ranging groups of the superfamily Ceratitaceae. However, the base of the Ladinian has recently been formalized with emphasis on the onset of the family Trachyceratidae (LO of the ammonoid Eoprotrachyceras curionii: Brack et al. 2005; Figs 17 –18), the most important shortranging ammonoid group of the Ladinian and Early Carnian. 6. The base of the Longobardian, although still a matter of debate, is characterized by an increasing diversity of Arpaditidae (e.g. appearance of Meginoceras and Silenticeras; Fig. 18) and Protrachyceratinae (e.g. different species of Protrachyceras), later followed by the Trachyceratinae (e.g. Maclearnoceras) that dominate the upper part of the substage.

Ammonoid zones Monnet & Bucher (2005b) provided a review of North American Anisian ammonoid biostratigraphy, arguing that it represents the most complete succession of Triassic ammonoid assemblages in the world. They thus recognize 13 zones, eight of which are divided into subzones (Fig. 12). Monnet et al. (2008) and Stiller & Bucher (2008) suggested a biostratigraphic correlation of this scale with some classic Tethyan successions, which are interpreted in a slightly different way by Mietto & Manfrin (1995a), who mostly studied the Southern Alps. Mietto & Manfrin’s scale subdivides the interval between the base of the Anisian and the LO of Eoprotrachyceras (¼ Ladinian sensu Brack & Rieber 1993) into 15 interval zones and subzones (Fig. 12).

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Fig. 16. Selected index ammonoids of the Middle and Upper Triassic from Nevada. Nevadites hyatti (a) lateral view, (b) ventral view. Paranevadites furlongi (c) lateral view, (d) ventral view. Eoprotrachyceras subasperum (e) lateral view, (f ) ventral view. Daxatina gr. desatoyense (g) lateral view, (h) ventral view. Joannites jacobus (i) lateral view, ( j) ventral view. a– j, from Jenks et al. 2007, a–b, pl. 20H–I, c–d, pl. 21A–B, e –f, pl. 24I–J, g –h, pl 26E–F, i– j, pl. 25F–G. Scale bars 20 mm.

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Fig. 17. Selected index ammonoids of the Tethyan Ladinian (Fassanian– Longobardian). Eoprotrachyceras curionii holotype (a –b) ventral and aperture views of the holotype, (c) specimen from Bagolino in lateral view. Protrachyceras margariosum, syntype (d) lateral view, (e) ventral view, (f ) aperture view. Protrachyceras gredleri, lectotype (g) lateral view, (h) ventral view. Protrachyceras longobardicum, syntype (i) lateral view, ( j) aperture view. Protrachyceras neumayri, syntype (k) lateral view. Frankites regoledanus, syntype (l) lateral view. a –b, from Mojsisovics 1882b, pl. 14, fig. 4a; c, from Gaetani (ed.) 1993 pl. 8, fig. 7; d– f, from Mojsisovics 1882b, pl. 82, fig. 1; g– h, from Mojsisovics 1882b, pl. 17; i –j, from Mojsisovics 1882b, pl. 22, fig. 5a –b; k, from Mojsisovics 1882b, pl. 14, fig 1; l, from Mojsisovics 1882b, pl. 29, fig. 7. Scale bars 20 mm. Upper left scale bar applies to a–b, upper right scale bar applies to c, middle left scale bar applies to d– f, middle right scale bar applies to g– h, lower left scale bar applies to i– j, lower middle scale bar applies to k, lower right scale bar applies to l.

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Fig. 18. Selected index ammonoids of the Ladinian (Fassanian–Longobardian) from British Columbia. Eoprotrachyceras matutinum, holotype (a) lateral view, (b) ventral view. Tuchodiceras poseidon holotype (c) lateral view, (d) aperture view, (e) ventral view. Meginoceras meginae holotype (f ) lateral view, (g) aperture view. Maclearnoceras maclearni holotype (h) lateral view, (i) ventral view. Frankites sutherlandi holotype ( j) lateral view. a–b, from Tozer 1980, pl. 17.1, figs. 4 –5; c–g, j, from Tozer 1967, pl. 8, c–e, fig. 4, f–g, fig. 6, j, fig. 12; h–i, from Tozer 1994, pl. 77, fig. 1a– b. Scale bars 20 mm. Figures a– b, c–g, and j, h –i are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (New genera of Triassic Ammonoidea. Current Research, Part A, GSC Paper 80-1A; Standard for Triassic Time, GSC Bulletin 156; Canadian Triassic ammonoid faunas, GSC Bulletin 467).

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Ladinian scales have been recently provided by Tozer (1994) for North America and by Brack et al. (2005) and Mietto & Manfrin (1995a, b) for the western Tethys. Brack et al. subdivide the Ladinian into four zones, whereas the same interval is divided by Mietto & Manfrin into six (sub)zones (Fig. 13). In North America, the Ladinian consists of five zones (Fig. 13), among which the last three are subdivided into a total number of eight subzones.

Late Triassic Stages At present, the STS recognizes three stages – Carnian, Norian and Rhaetian – for the Upper Triassic. The LO of the ammonoid Daxatina canadensis at the Prati di Stuores/Stuores Wiesen section in northern Italy (Mietto et al. 2007a, b) has recently been ratified as the GSSP for the base of the Carnian stage (Gaetani 2009). Most workers divide the Carnian into two substages, which in the Tethyan realm are referred to as Julian and Tuvalian (Fig. 1). The Norian stage is divided into three substages, which are simply distinguished as Lower, Middle and Upper in North America, whereas in the Tethys they are named Lacian, Alaunian and Sevatian. All of the Carnian and Norian substages are based on ammonoid biostratigraphy. The STS recognizes the Rhaetian as the youngest Triassic stage. The recently agreed on redefinition of the Rhaetian is of a ‘long’ Rhaetian that includes two (or three) ammonoid zones with its base very close to or at the base of the Paracochloceras suessi zone. However, the favoured definition of the Rhaetian base is the LO of the conodont Misikella posthernsteini (Krystyn et al. 2007b).

Ammonoid provinces and biotic events The Late Triassic ammonoid record is dominated by diverse Tethyan assemblages that are readily distinguished from much less diverse and somewhat endemic Boreal ammonoid assemblages. Furthermore, the faunal diversity ratio between low and high palaeolatitudes is about 10:1 (Dagys 1988) during both the Carnian and the Norian. The boreal ammonoid assemblages from the Carnian to the Middle Norian are mostly composed of Sirenitinae and rarer Ussuritidae, Arcestidae (Arcestes) and Gymnitidae (Placites). The Late Norian record consists of very few Arcestes, Placites and Rhacophyllites (Dagys 1988), together with the bivalves Eomonotis and Monotis (e.g. Konstantinov et al. 2003). The major subdivisions of the Late Triassic directly reflect the appearance and radiation of several important families and subfamilies, as well

as extinctions. The substages are subdivided in the following manner: 1. The base of the Lower Carnian and of the Julian substage in its emended meaning (see Tozer 1984) is defined by the Lowest Occurrence of Daxatina canadensis at level SW4 of Prati di Stuores/ Stuores Wiesen (Dolomites, Italy). On the whole, the Julian is dominated by Trachyceratinae, in particular Trachyceras and Austrotrachyceras, and by Sirenitinae (Figs 19, 21– 22). 2. The base of the Tuvalian is marked by one of the major changes in the evolution of Triassic ammonoids, namely the crisis of Trachyceratinae, whose only survivor in the Upper Carnian is Trachysagenites, as well as the radiation of Tropitidae (e.g. Tropites (Figs 19, 22) and closely allied forms) and to a lesser extent Arpaditinae. 3. The base of the Norian and of the Lacian is characterized by the nearly complete disappearance of Tropitidae and the appearance of new members of Juvavitinae, such as Guembelites and Dimorphites, and of the Thisbitidae such as Stikinoceras (Figs 20 –22). 4. The base of the Alaunian is marked by the appearance of new genera of Cyrtopleuritidae (Drepanites (Fig. 20) and Cyrtopleurites). Members of this family (e.g. Himavatites, Mesohimavatites, Neohimavatites; Figs 20, 22) characterize the substage together with certain Haloritinae such as Halorites and Thisbitidae (e.g. Phormedites). 5. The base of the Sevatian is characterized by a decrease in ammonoid diversity and the first occurrence of Rhabdoceras, the first heteromorphic ammonoid. Common forms (Figs 23 –24) are Haloritinae (Gnomohalorites and Catenohalorites) and Sagenitidae (Sagenites gr. quinquepunctatus). 6. The base of the Rhaetian eventually will be defined by conodonts, and the best marker candidate at present is the Lowest Occurrence of Misikella posthernsteini (Krystyn et al. 2007b; Krystyn 2008). In terms of ammonoids, the Norian/Rhaetian boundary is characterized by the appearance of Sagenites reticulatus and the heteromorphic Cochloceras and Paracochloceras (Krystyn et al. 2007b; Krystyn 2008; Fig. 24).

Ammonoid zones Most workers in southern Europe use the Carnian ammonoid zonation of Krystyn (1978; Krystyn in Zapfe 1983) that recognizes five zones divided into nine subzones (Fig. 13). The definition of the base of the Carnian as the LO of Daxatina canadensis (Mietto et al. 2007b; Gaetani 2009) is one of two ammonoid datums used to define Triassic stage bases. The other well-established Carnian ammonoid biozonation is from British Columbia and was most recently reviewed by Tozer (1994) as

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Fig. 19. Selected index ammonoids of the Carnian (Julian–Tuvalian) from British Columbia. Austrotrachyceras obesum holotype (a) lateral view, (b) ventral view. Sirenites nanseni holotype (c) lateral view, (d) ventral view. e–g, Tropites welleri (e) lateral view, (f ) aperture view, (g) drawing of sutural pattern. a– b, from Tozer 1967, pl. 9, fig. 1; c–d, from Tozer 1961, pl. 23, fig. 8; e– g, from Smith 1927, pl. 78, fig. 5– 7. Scale bars 20 mm. Upper left scale bar applies to a– d, lower left scale bar applies to g, lower right scale bar applies to e– f. Figures a– b, c–d are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Standard for Triassic Time, GSC Bulletin 156; Triassic stratigraphy and faunas, Queen Elizabeth Islands, Artic Archipelago. GSC Memoir 316).

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Fig. 20. Selected index ammonoids of the Carnian (Tuvalian) and Norian (Lacian– Alaunian) from British Columbia. Klamathites macrolobatus holotype (a) lateral view, (b) aperture view, (c) drawing of sutural patern. Stikinoceras kerri holotype (d) lateral view, (e) ventral view. Malayites dawsoni holotype (f ) lateral view, (g) ventral view. Drepanites ruthefordi (h) lateral view, (i) ventral view. Himavatites columbianus holotype ( j) lateral view, (k) ventral view. Himavatites columbianus (l) lateral view, (m) ventral view. a –c, from Silberling 1959, pl. 1, figs. 20, 22; d –e, from Tozer 1962, pl. 9, fig. 1; f– g, from Tozer 1967, pl. 10, fig. 2; h– i, from Tozer 1967, pl. 10, fig 4; j –k, from Tozer 1962, pl. 11, fig. 7; l –m, from Tozer 1962, pl. 10, fig. 9. Scale bars 20 mm. Upper scale bar applies to a– c, upper middle scale bar applies to d–g, lower middle scale applies to h– i, lower left scale bar applies to j –k, lower right scale bar applies to l– m. Figures d –e, f–g, h –i, j– k are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Illustrations of Canadian fossils: Triassic of western and Arctic Canada, GSC Paper 62-19; Standard for Triassic Time, GSC Bulletin 156).

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Fig. 21. Selected index ammonoids of the Carnian (Julian) of the Tethys and British Columbia. Daxatina canadensis (a) lateral view, (b) ventral view, (c) sutural pattern (26.25 mm wide). (d–i) Trachyceras aon (d) lectotype in lateral view, (e–f ) original of Klipstein 1843 (e) ventral view, (f) lateral view, (g– i) original specimen of Mu¨nster in Wissmann & Mu¨nster 1841 (g) lateral view, (h) ventral view, (i) aperture view. Trachyceras aonoides, typus in ( j) lateral view, (k) aperture view, (l) sutural pattern (78.5 mm wide). a–c, from Tozer 1994, fig. 68b, pl. 85, fig. 9ab; d– i, from Urlichs 1994 pl. 1, d, fig. 1, e–f, fig. 2 and g –i, fig. 3; j – k, from Mojsisovics 1893, pl. 191, fig. 3. Scale bars 20 mm. Upper left scale bar applies to a– b, upper right scale bar applies to c, middle scale bar applies to d –i, lower scale bar applies to j – k, l not to scale. Figures a–c are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Canadian Triassic ammonoid faunas, GSC Bulletin 467).

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Fig. 22. Selected index ammonoids of the Tethyan Carnian (Julian– Tuvalian) and Norian (Alaunian). Austrotrachyceras austriacum, holotype (a) lateral view, (b) ventral view. Tropites dilleri holotype (c) lateral view, (d) aperture view. Tropites subbullatus holotype (e) lateral view, (f ) aperture view, (g) cross-section view. Anatropites spinosus, holotype (h) lateral view, (i) aperture view. Guembelites jandianus holotype ( j) lateral view, (k) aperture view. Juvavites magnus, (l) lateral view, (m) ventral view, (n) aperture view. Himavatites hogarti holotype (o) lateral view, (p) ventral view. a–b, from Mojsisovics 1893, pl. 184, fig. 1; c –d, from Smith 1927, pl. 68, figs. 1–2; e– g, from Mojsisovics 1893, pl. 106, fig. 7; h –i, from Mojsisovics 1893, pl. 110, fig. 2; j– k, from Mojsisovics 1896, pl. 10, fig. 1. l –n, from Tozer 1962, pl. 9, fig. 5; o –p, from Diener 1906, pl. 9, fig. 1. Scale bars 20 mm. Upper left scale bar applies to a– b, e–i, upper right scale bar applies to c– d, middle scale bar applies to j – k, lower left scale bar applies to l– n, lower right scale bar applies to o– p. Figures l –n are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Illustrations of Canadian fossils: Triassic of western and Arctic Canada, GSC Paper, 62-19).

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Fig. 23. Selected index ammonoids of the Norian (Sevatian) and Rhaetian from British Columbia. Gnomohalorites cordilleranus, holotype (a) lateral view, (b) ventral view, (c) aperture view. Parachocloceras amoenum holotype (d –e) lateral views, (f ) axial view. Choristoceras crikmayi, holotype (g–i), lateral views, (h) ventral view. a– c, from Tozer 1994, pl. 145, fig. 4; d–f, from Mojsisovics 1893, pl .137, fig. 26; g– i, from Tozer 1994, pl. 147 fig. 22. Scale bars 20 mm. Upper left scale bar applies to a–c, upper right scale bar applies to d– f, lower scale bar applies to g–i. Figures a –c, g– i are reproduced with the permission of Natural Resources Canada 2009, courtesy of the Geological Survey of Canada (Canadian Triassic ammonoid faunas, GSC Bulletin 467).

consisting of six zones that are readily correlated to the European zonation (Fig. 13). The cosmopolitanism of Tropites dilleri at the base of the Tuvalian is a key element in this correlation. Norian ammonoid zonations are also generally twofold, one for Tethyan successions and a second for British Columbia (Fig. 10), although local zonations have been proposed, for example in Siberia and Svalbard (Spitzbergen) (Weitschat & Dagys 1989; Dagys & Weitschat 1993). Again, the cosmopolitanism of one ammonoid (Juvavites magnus, Fig. 22) aids a fairly non-controversial correlation of Norian ammonoid zones with each other.

The recent agreed-on definition of the Rhaetian (based on a conodont datum: Krystyn et al. 2007b) essentially equates the base of the stage with the Paracochloceras amoenum zone (Figs 13, 23, 24). This produces a so-called ‘long’ Rhaetian comprised of two or three ammonoid zones. The youngest substage of the Norian, the Sevatian, is thereby reduced to one ammonoid zone. Recent voting by the Subcommission on Jurassic Stratigraphy has chosen the LO of the ammonoid Psiloceras spelae tirolicum (cf. Lucas et al. 2007b; Hillebrandt et al. 2007; Hillebrandt & Krystyn 2009) as the base of the Hettangian stage (¼ base

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Fig. 24. Selected index ammonoids of the Tethyan Norian (Alaunian –Sevatian) and Rhaetian. Halorites macer holotype (a) lateral view, (b) aperture view. Sagenites quinquepunctatus holotype (c) lateral view, (d) aperture view. Sagenites reticulatus, (e) lateral view, (f) aperture view. Choristoceras marshi (g) lateral view. a– b, from Mojsisovics 1893, pl. 75, fig. 2a –b; c–d, from Mojsisovics 1893, pl. 125, fig. 3; e– f, from Mojsisovics 1893, pl. 100, fig. 1; g, Mojsisovics 1893, pl. 135, fig. 13. Scale bars 20 mm. Upper left scale bar applies to a– b, e–f, upper right scale bar applies to c–d, g.

of the Jurassic). The Upper Triassic thus is composed of at least 17 ammonoid zones (Fig. 13).

Some remarks on modern ammonoid biostratigraphy The present day legacy of the very long history of ammonoid biostratigraphy is complex, and the evolution of theory and practice, as well as the personalities of some palaeontologists, have produced some diversity of opinions and different approaches. Here, we briefly address four points of special significance.

The influence of taxonomy over biostratigraphy As already mentioned in previous sections, during the Triassic the ammonoids underwent a very rapid major evolutionary radiation as well as a generally wide palaeogeographical dispersal. These features make this group ideal as a bio-chronostratigraphic tool, but quite frequently their high potential for stratigraphy is impaired by a wide variety of taxonomic problems that often take up most of the specialist’s research time. Some of these problems were emphasized by Tozer in 1971, namely, the typological or populationistic approach to species

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definition and the different approaches to phylogeny reconstruction. Here, we complement his overview by also offering a Tethyan point of view on these issues and by taking into account the last 20 years of experience with ammonoid-based GSSP definitions. The first requirement for the utilization of fossils for bio-chronostratigraphy is a clear and agreed-on definition of species and genera. However, for Triassic ammonoids such a basic condition often is not established. For several groups of ammonoids, the 150-year-old taxonomy does not represent a strong foundation for building a shared bio-chronostratigraphic scale, but instead produces a more-or-less long chain of opinions, differing interpretations and uncertainties. A significant number (1500–2000) of ammonoid species were described in the nineteenth century on the basis of specimens with no accurate position of collection. Some of these species were described based on incomplete or poorly preserved or juvenile specimens and, consequently, they amount to no more than empty names. Other species either were not very well described or they were not illustrated properly (or both) and, as a result, their separation from similar species has been a matter of debate for decades. When these uncertainties involve species that are of great bio-chronostratigraphic value, the consequences for the accuracy of biozones and correlation are dramatic. A good example of this type of problem is the controversy surrounding the separation of Ceratites binodosus (Hauer) and Ceratites trinodosus Mojsisovics (Fig. 14), which in the first half of the twentieth century called into question the separation of the two corresponding bio-chronozones (Fig. 3; for more details see Assereto 1971). Ceratites binodosus (Hauer) was revised by Mojsisovics in 1882, in the same monograph in which he described C. trinodosus. He also used these two species as indexes of the two most complex bio-chronozones of his scale (Fig. 3). The faunal lists of the lithofacies of these two zones were not totally different, and some important genera, such as Balatonites and Judicarites (known at the time as the group of Balatonites euryomphalus), were documented from both zones. A further source of confusion was that the two faunas actually did not occur in succession at any field locality, but instead their succession was inferred from faunal comparision and stratigraphic correlations. These two serious shortcomings aptly demonstrate the importance of taxonomic separation of the two index species in justifying the separation of the two ammonoid zones. At the beginning of the twentieth century certain authors suggested that the two index species were conspecific, and for several decades only one of the two zones was included in chronostratigraphic

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scales; otherwise, they were replaced by a brachiopod (Rhynconella decurtata zone in place of C. binodosus zone: Arthaber 1905) or daonellid and algae zones (Daonella sturi and Diplopora annulatissima zone in place of C. trinodosus zone: Pia 1930). Then, in 1971, Assereto demonstrated the faunal succession of the two zones, but the generic attribution of Ceratites binodosus has not yet been solved. Mietto & Manfrin (1995a), Monnet et al. (2008) and Stiller & Bucher (2008) refer the species to Schreyerites Tatzreiter & Balini and to the Pelsonian substage of the Anisian. Assereto, in a former time, attributed C. binodosus to Paraceratites, a very typical Illyrian genus, and then he suggested that the Pelsionian/ Illyrian boundary should be placed at the base of the P. binodosus-bearing level. Another example of the influence taxonomy may have on biostratigraphy is the recent history of the generic assignment of Ceratites reitzi Bo¨ckh. This species was based on five specimens from Felso´´o¨rs (Bakony), characterized by radial ribs bearing ventrolateral nodes on the shoulder and ending on the ventral side with a second row of more prominent nodes (Fig. 14). The furrow-like middle part of the venter is in part similar to the furrow of the earlier trachyceratids and this is the reason Mojsisovics (1882) assigned the species to Trachyceras and then later to Protrachyceras (Mojsisovics et al. 1895). The species remained within Protrachyceras (e.g. Arthaber 1905; Jacobshagen 1967) until its generic status became one of the key points for the definition of the GSSP of the Ladinian stage. In 1986, Brack & Rieber described a group of Nevadites specimens showing similarities with C. reitzi, from the Buchenstein Formation of eastern Lombardy, some levels below the level yielding the first narrowly-furrowed true trachyceratid Eoprotrachyceras (Fig. 17). Brack & Rieber’s attribution of Ceratites reitzi to Nevadites Hyatt suggested the correlation of the 100-yearold Tethyan ‘Trachyceras’ reitzi zone sensu Mojsisovics 1882 with the North American Frechites occidentalis zone of Nevada (Silberling & Nichols 1982), characterized by the occurrence of Nevadites, as well as with the Nevadites zone of Epidauros (Krystyn 1983). However, this new attribution of C. reitzi was not fully satisfactory, because Nevadites usually exhibits a rather broad and depressed flat venter (Fig. 16) instead of the narrow, and shouldered venter of C. reitzi. A short time later the taxonomic position of C. reitzi was again revised. Vo¨ro¨s & Palfy (1989) returned the species to Wang’s (1983) earlier attribution of Xenoprotrachyceras, but the final solution was provided again by Brack & Rieber (1993), who discovered in the Buchenstein Formation at Bagolino (eastern Lombardy) several layers below the layers yielding Nevadites, about

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40 specimens (Fig. 14) fitting very well with the types of C. reitzi. This large number of specimens helped Rieber to understand the ontogeny and variability of the group, which he recognized as quite different from the ontogeny and variability of Nevadites. For these reasons Brack & Rieber erected the new genus Reitziites to accommodate C. reitzi, and they then described the specimens formerly attributed to the group of Nevadites reitzi as the new species Nevadites secedendis. Thus, the beds yielding Nevadites secedensis were confirmed as coeval with the Occidentalis zone of North America on the basis of the common occurrence of Nevadites, and R. reitzi could not be used to correlate the Tethyan successions with North America. A final consequence of this long-lasting taxonomic revision was that the proposal to use the Lowest Occurrence of R. reitzi to define the base of the Ladinian stage (Vo¨ro¨s et al. 2003) could not get enough support to be selected as marker event for the GSSP of the Ladinian stage. The third and last example of intimate interdependence of ammonoid taxonomy with Triassic biochronostratigraphy is from the Ladinian/Carnian boundary interval. A large part of the work and discussions within the Ladinian/Carnian boundary Working Group of the STS was focused on the comparison of ammonoid events with conodont (LO of Metapolognathus polygnathiformis) and bivalve (LO of Halobia) events. Two possible ammonoid markers were examined, that is, the Lowest Occurrence of Daxatina canadensis and the Lowest Occurrence of Trachyceras s.s. The genus Trachyceras s.s. has always been considered as a global marker for the Carnian stage, while Daxatina has been always regarded as Late Ladinian in North America. The discovery of Daxatina below Trachyceras, but also together with the earliest Trachyceras in the Dolomites (Mietto & Manfrin 1995b), together with the calibration of the Lowest Occurrence of the very typical Carnian conodont M. polygnathiformis slightly below the LO of Daxatina in the Himalayas (Balini et al. 1998; Krystyn et al. 2004), has redirected the attention from Trachyceras to Daxatina, which was eventually selected as the GSSP marker, partly on the basis of practical problems with the taxonomic separation from Trachyceras. Daxatina and Trachyceras (Fig. 21) are almost identical in shell shape and ornamentation, especially because each has a row of double nodes on each side of the ventral furrow, but they differ in their suture line, which is ceratitic in the former (Fig. 21c) and ammonitic (Fig. 21l) in the latter. Given the lack of better marker candidates, the basal Carnian defined at the LO of Daxatina cannot be confused with the latest Ladinian; either the suture line is visible or not, but the same cannot be said for Trachyceras. A single specimen

with double nodes on each side of the furrow, but with no exposed suture, would be of questionable latest Ladinian or earliest Carnian age.

Biostratigraphy and Biochronology By tradition, ammonoids have been always regarded as the best guide fossils for the Triassic System. The link between ammonoids, their distribution and their time significance has been so strict in the past that often it is difficult to discriminate whether an author referred to an ammonoid fauna from a biostratigraphic point of view or from a chronostratigraphic point of view. Oppel zones well represent this dual significance. At the present in the Triassic literature there are two different but not conflicting ways to use ammonoids (as is true for other fossils) to detect/recognize time: biostratigraphy and biochronology. Biostratigraphy works very well on fossiliferous successions with a very good and continuous fossil record, but this approach is only an intermediate step towards chronostratigraphy. Biochronology (sensu Tozer 1971; Bucher 1988, 1989) is the only good solution for the case of a discontinuous stratigraphic distribution of fossils. In this case, the occurrence/assemblage of the species (i.e. the content of the fossil-bearing beds) is much more important than the first and last occurrence of the single taxa, which are facies controlled. The North American standard scale (Tozer 1967; Silberling & Tozer 1968) was presented as a biostratigraphic scale, but actually it was conceived as a biochronologic scale (see Tozer 1971: p. 990), that is, as a succession of ammonoid assemblages separated by bodies of rocks without fossils. This approach, followed also by Bucher (1988, 1989 and following papers), unfortunately does not conform with the requirements for the formalization of the chronostratigraphic units of the Standard Global Chronostratigraphic Scale. These units are organized in a continuous succession, so that they must be formalized through the definition of their boundaries and not by their content. This requirement implies working on fossiliferous successions spanning the boundary, instead of on successions with discrete distribution of fossils. For this reason, the biostratigraphic approach is the most suitable for the definition of the GSSPs.

Environmental control on ammonoid distribution and on biozones Triassic ammonoid distribution was mostly a function of time, but was also influenced by palaeogeography and environment. However, our

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understanding of these kinds of controls is still incomplete. The general outline of ammonoid palaeobiogeographical distribution, consisting of five palaeobioprovinces, has been well established for decades (e.g. Tozer 1981a, b; Urlichs & Mundlos 1985; Dagys 1988), but recent papers notably improve the detail of the reconstructions, especially for the Lower Triassic (Brayard et al. 2006, 2007, 2008). On the contrary, palaeoenvironmental control of the distribution of ammonoids is generally underestimated. In a broad sense it has been well known since the nineteenth century that red limestones of the ‘Ammonitico Rosso’ or Hallstatt/Han Bulog facies are dominated by smooth and relatively longranging leiostracan ammonoids, while the ornamented and short-ranging trachyostracan forms are more common in uncondensed successions (e.g. Jacobshagen 1967; Krystyn et al. 1971a; Tozer 1971; Krystyn 1991). However, palaeoenvironmental controls not only influence Hallstatt facies with respect to the remainder of marine successions, but they also affect basinal and carbonate platform assemblages. Faunal variability can be so important that it leads to mistakes in the correlation of faunas and, as a consequence, in the definition of chronozones. This was the case with the above-mentioned Dinarites avisianus Zone, introduced by Mojsisovics (in Mojsisovics et al. 1895) to separate a fauna from Latemar and Forno (Dolomites) previously included in the Trachyceras archelaus Zone as a part of lithofacies b, ‘thonarmen Kalksteine’ (Mojsisovics 1882b; Fig. 3). The ‘avisianus fauna’ was found in limestone debris from carbonate platforms and was placed by Mojsisovics (Fig. 4) in an intermediate position between the Protrachyceras curionii and the Protrachyceras archelaus zones on the basis of supposed stratigraphic position, being the ‘avisianus fauna’ dominated by trachyostraca ammonoids not documented from basinal successions. Unfortunately, Mojsisovics’ stratigraphic correlation between basinal facies, where the P. curionii and P. archelaus zones were defined, and the carbonate platform (Latemar Limestone) was incorrect, and it took almost one century to restore the correct stratigraphic successions of zones (Assereto 1969) and to calibrate the ammonoid biostratigraphy of carbonate platforms with basinal successions (Manfrin et al. 2005). The Hallstatt and the carbonate platform ammonoid faunas are not the only examples of palaeoenvironmental control over faunal composition, because sometimes faunal composition might change also in different parts of the same basin, as shown by Vo¨ro¨s (2002) for the latest Anisian of the Balaton Highlands (Hungary). Faunal diversity

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and composition in terms of genera change notably from basinal palaeoenvironments relatively close to carbonate platforms to more distal environments, over a distance of some tens of kilometres.

Type of zones One of the main difficulties when dealing with papers on ammonoid stratigraphy is understanding the type of zones used by the authors. The meaning, boundaries, composition and name of the zones have usually changed through time, and this is especially true of the old zones coined in the nineteenth century. Unfortunately, even in the recent literature the authors do not always pay much attention to explaining these crucial details, and they have to be deduced from their text and figures. Taking into account the possible dual bio- and chronostratigraphic meaning of the ammonoid zones, and that a wide variety of stratigraphers inexperienced with respect to ammonoid problems are actually using or referring to ammonoid zones, these uncertainties might have some severe consequences. It is almost impossible to describe all the problems in the nomenclature and definition of ammonoid zones, but we would like to warn stratigraphers about them. Generally speaking, in the relatively recent biochronologic ammonoid scales the most frequent change consists of finding new faunas within a barren interval between two ammonoid zones. In the palaeobiogeographical domains with an older tradition, such as the Tethyan realm, the main problem is the ‘translation’ of the nineteenth century Oppel zones, no longer accepted by the International Stratigraphic Guide (Salvador 1994), into one or more biozones fulfilling the requirements of the ISG (range, interval, assemblage, abundance and lineage zone; for example, see Fig. 5).

The future of ammonoid biostratigraphy The main uncertainty for the future of Triassic ammonoid biostratigraphy is not the decline of ammonoids as a tool for high resolution dating and correlation of Triassic sedimentary sequences, but rather it is in the dramatic decrease in the number of specialists, due to the lack of replacement of experienced palaeontologists who started their activity in the 1950s and 1960s. Undoubtedly such a problem is common in several fields of palaeontology, but its consequence will be very serious for the Triassic scientific community. In the last 10 years the number of active ammonoid specialists has been reduced by about 40% because N. J. Silberling (Denver, USA), E. T. Tozer (Vancouver, Canada), H. Rieber (Zu¨rich, Switzerland), W. Weitschat (Hamburg, Germany) and M. Urlichs (Stuttgart,

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Germany) retired and A. S. Dagys (Lithuania), A. A. Shevyrev (Moscow, Russia) and N. Sestini Fantini (Milano, Italy) passed away. The problem is most severe in North America, where there is not a single full time palaeontologist working on Triassic ammonoids. However, the situation in the rest of the world is not much better, because the number of full time positions is decreasing and, consequently, several very active, young PhD palaeontologists will not have as many opportunities to continue their research. Very warm thanks are extended for careful and timely reviews by M. Gaetani (Milano, Italy), J. G. Ogg (Purdue University) and N. J. Silberling (Lakewood, CO). Their comments and suggestions have been very helpful in improving the manuscript. Special thanks to G. Chiodi (Milano University) for taking several tens of pictures from plates of Middle and Upper Triassic ammonoids. We also warmly thank J. Tremblay (GSC, Ottawa) for the authorization to reproduce photographs of specimens from the Geological Survey of Canada publications and C. Lombardo (Milano University) for the permission to reproduce photograph from Rivista Italiana di Paleontologia e Stratigrafia. We warmly thank GSL production editor Helen Floyd-Walker for her invaluable help and remarkable patience during the final preparation of this paper.

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Triassic palynology of central and northwestern Europe: a review of palynofloral diversity patterns and biostratigraphic subdivisions ¨ RSCHNER* & G. F. WALDEMAAR HERNGREEN WOLFRAM M. KU Laboratory of Palaeobotany and Palynology, Palaeoecology, Utrecht University, Budapestlaan 4, 3584 CD Utrecht, The Netherlands *Corresponding author (e-mail: [email protected]) Abstract: We document palynofloral trends through the Triassic in the Germanic and Alpine facies with an emphasis on diversity trends and possibly related palaeoenvironmental changes. As a first order approximation of palynofloral diversity, we used the range through method of the software package PAST based on a range chart compiled from several Triassic palynological studies and reviews. Our analysis suggests that during the entire Triassic the diversity of plants producing spores was largely controlled by the availability of water, while diversity among gymnosperms was also affected by other environmental and biotic factors. In general, palynofloral diversity declines by some 50% between the early Carnian and the Norian, mainly as a result of a decrease in the number of pollen species. This is the second most severe loss in pollen species after the Permian– Triassic biotic crisis. In comparison to the marked palynofloral turnover at the Permian–Triassic transition and the end-Carnian decrease in palynofloral diversity, the endTriassic biotic crisis appears to have little affected palynofloral species diversity in Europe. A study of the palynostratigraphy of NW Europe recognizes nine zones (and nine subzones) that encompass the Triassic, most of which have their boundaries based on the first occurrences of marker species. The palynostratigraphic zones and subzones in Europe are correlated to the marine Triassic stages based on various data, including numerous palynological records in marine Alpine Triassic strata.

The Trias in Central Europe is developed in two major facies domains, the predominantly continental facies of the Germanic realm and the mainly marine facies in the Alpine realm. It has been about 175 years since Friedrich August von Alberti established the Triassic System based on his observations in the Germanic Basin (Alberti 1834). He combined into one ‘formation’ a succession of three lithostratigraphic units that comprises the Bunter Sandstein (redbeds); the Muschelkalk (marine limestones); and the Keuper (continental shales, sandstones and evaporites). About the same time the British geologists Roderick Impey Murchison and Adam Sedgwick explored the marine limestones of the Alps and published a geological map of the eastern Alps showing the extent of the Alpine limestone (Sedgwick & Murchison 1835). However, they misinterpreted the age of these rocks, which are now known to be Triassic, assigning them to the Jurassic. The German geologist Christian Leopold von Buch was among the first to describe the fossil content of the Alpine Triassic (von Buch 1845). Since the introduction of palaeopalynology as a biostratigraphic discipline, the study of pollen and spores has become an important tool for dating the continental successions of the Germanic Basin, which are often poor in stratigraphically useful

fossils. While the initial work on (in-situ) Triassic miospores were already done during the early 20th century (Sollas 1901; Wills 1910), the first palynostratigraphic studies were carried out by one of the pioneers of pre-Quaternary palynology, Friedrich Thiergart (1949), followed by, among many others, Leschik (1955), Schulz (1962, 1964, 1965, 1966, 1967), Ma¨dler (1964a, b, 1968), Clarke (1965a), Visscher (1966), Warrington (1967, 1970), Chaloner (1969) and Scheuring (1970). As pollen and spores can be dispersed by wind and water over wide distances, they can be deposited not only in continental but also in marine sedimentary environments. Consequently, pollen and spores, in contrast to virtually all other fossils, have the potential to provide biostratigraphic correlations between sediments deposited under different environmental conditions. The first palynological studies in the Alpine realm were carried out by Klaus (1960, 1963, 1964) and Venkatachala & Go´cza´n (1964). Since then palynological studies (e.g. Morbey 1975; Visscher & Krystyn 1978; Van der Eem 1983; Roghi & Dalla Vecchia 1997; Roghi 2004) in marine successions with ammonoids or other biostratigraphic markers have resulted in a palynological characterization of Triassic stages and sub-stages in Europe. Unfortunately, some of the ammonoid-bearing lithologies in the classical areas of the Alpine facies in the

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 263– 283. DOI: 10.1144/SP334.11 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Austrian and Italian Alps, such as the Norian limestone successions, appear to be unfavourable for palynological investigations. However, several palynological studies (Brugman 1986; Visscher & Brugman 1981; Visscher et al. 1994) have demonstrated their importance for the correlation of units in the Germanic Triassic facies with those in the Alpine realm. The aim of this review is twofold. The first part deals with documentation of palynofloral trends through the Triassic in the Germanic and Alpine facies with an emphasis on diversity trends and possibly related palaeoenvironmental changes. The second part is a review of the palynostratigraphy of this region. As a first order approximation of palynofloral diversity we used the range through method of the software package PAST (Hammer et al. 2001) based on a range chart compiled from several Triassic palynological studies and reviews (e.g. Brugman 1983a; Visscher & Brugman 1981; Orłowska-Zwolin´ska 1983, 1985, 1988; Fijałkowska 1992, 1994b; Warrington 1996b; Batten & Koppelhus 1996; Herngreen 2005). This approach has been successfully applied in Cenozoic palynological studies (Jaramillo et al. 2006) and may serve as a proxy for estimating the standing diversity (defined in the original sense of number of species) of the pollen and spore producing plants (Weng et al. 2007). In Figure 1 diversity (expressed as the number of pollen and spore taxablack dots), the number of originations (diamonds) and the number of extinctions (squares) per Triassic sub-stage is shown. A European Triassic palynological database, which contains records of c. 300 pollen and spore morphotaxa can be obtained on request from the first author.

General palynofloral trends, diversity patterns and related palaeo-environmental changes Latest Permian palynomorph assemblages recovered from Zechstein deposits are characterized by highly diverse bisaccate pollen associations with c. 80 sporomorph taxa (Fig. 1a). Monolete or alete, multi- or non-taeniate forms, produced by gymnosperms (Schaarschmidt 1963; Fijałkowska 1992), with such morphogenera as Jugasporites, Limitisporites, Lueckisporites, Lunatisporites, Protohaploxypinus and Strotersporites, as well as nonsaccate taeniate forms such as Vittatina and the large prepollen of the Late Permian conifer species (Poort et al. 1997) known as Nuskoisporites dulhuntyi, typify the latest Permian pollen assemblages. Late Permian conifer macro-remains described from the Zechstein of Western Europe

(e.g. Stoneley 1958; Schweitzer 1960) and the Val Gardena Formation of the Southern Alps (e.g. Clement-Westerhof 1984, 1987) possess xeromorphic characters. This, and the presence of evaporitic deposits, is indicative of a warm and dry climate, which is also supported by climate models (e.g. Kutzbach & Ziegler 1994). The Permian– Triassic transition is characterized by a marked palynofloral turnover whereby assemblages dominated by bisaccate pollen are frequently replaced by ones dominated by spores (Visscher 1971; Visscher & Brugman 1988). The decrease in pollen diversity during the earliest Triassic was dramatic, with almost 70% of Late Permian pollen taxa disappearing (Fig. 1b). However, spore diversity seems to have been relatively unaffected by the endPermian biotic crisis (Fig. 1c). The oldest Triassic assemblages, of the Lundbladispora obsoleta–Protohaploxypinus pantii Zone, are from the Lower Buntsandstein. They are characterized by the appearance of spore genera (Densoisporites, Endosporites, Lundbladispora, Kraeuselisporites) together with abundant taeniate bisaccate pollen (mainly Protohaploxypinus pantii with subordinate Lunatisporites, Strotersporites and Striatopodocarpites). Throughout nearly the entire Triassic, central Europe was located within the subtropical climate belt (Kent & Tauxe 2005). General circulation models indicate that the climate was influenced by a strong monsoonal circulation pattern (Kutzbach & Gallimore 1989; Parrish 1993; Sellwood & Valdes 2007). During the course of the Triassic, changes in the strength of the monsoon circulation and palaeogeographic changes, such as the marine inundation during the Mid Triassic, may have resulted in the long-term alternation of wetter and drier periods. A warm, semi-arid climate during the first half of the Early Triassic is suggested by the predominance of bisaccate pollen, particularly taeniate forms (Visscher & Van der Zwan 1981; Yaroshenko 1997). After the demise of the highly diverse endPermian palynofloras, those from the Early Triassic (Lower Buntsandstein) became progressively dominated by spores, mainly Densoisporites nejburgii, produced by the lycopsid Pleuromeia, a succulent quillwort (Grauvogel-Stamm 1999). Other lycopsid spores include Aratrisporites, Densoisporites playfordii, Endosporites pallidus, Kraeuselisporites spp. and Lundbladispora spp., which were produced by Isoetales and Selaginellales. These lycopsid spores reflect a herbaceous pioneer vegetation that is believed to have populated relatively dry, welldrained floodplain environments (Van der Zwan & Spaak 1992). As a result of the continuing impoverishment of the end Permian, palynofloral diversity declined to a minimum of c. 20 taxa during the

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

Number of Sporetaxa[n]

(c)

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60

40

20

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80

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100 80 60 40 20 0 Ch. I. Ol. Anisian

250

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240

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Rh. Het.

200

Fig. 1. Changes in palynofloral diversity from the Late Permian (Changhsingian) to the early Hettangian (Lias a1) calculated for mean ages of sub-stages using TSCreator version 4.0 visualization program at www.tscreator.org, June 2009. (Ogg, J. G. and ICS Subcommission for Stratigraphic Information) for: (a) pollen and spores; (b) pollen only; and (c) spores only. The line with dots indicates standing diversity (number of pollen and spore taxa), the line with diamonds the number of originations and the line with squares the number of extinctions.

early Induan (Fig. 1a). These monotonous palynofloras persist into the late Induan though spores such as Cycloverrutriletes presselensis and Verrucosiporites are common locally. The Densoisporites nejburgii acme, which corresponds to the Middle Buntsandstein in Germany (Schulz 1964), is also recognized in time equivalent intervals in Poland

(Orłowska-Zwolin´ska 1984; Fijałkowska 1994b) and the Dolomites (Visscher 1974). The predominance of D. nejburgii and related lycopsid forms, in association with numerous acritarchs, coincides with the Olenekian transgression. It has been suggested that the wide spread of lycopsid spores in both high and low latitudes may have been

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related to an increase in humidity during the second half of the Early Triassic (Yaroshenko 1997). Renewal of the gymnosperm flora, which resulted in an increase in pollen diversity (Fig. 1b), started in the late Olenekian (in the Hardegsen and Solling formations in NW Europe). Pollen such as Angustisulcites, Triadispora crassa and Voltziaceaesporites heteromorphus enter the record. Spore species diversity, however, remained relatively stable until the Ladinian (Fig. 1c). The first wave of reoccupation by woody plants and the concomitant stepwise retreat of the lycophyte-dominated pioneer vegetation may have been sustained by an amelioration of environmental conditions (Galfetti et al. 2007). However, the ultimate proliferation of woody vegetation indicated by a progressive increase in pollen species diversity (Fig. 1b) did not occur until the Early/Mid Triassic transition, at the Solling-Ro¨t Formation transition in NW Europe (Looy et al. 1999). The appearance of taxa such as Illinites spp., Protodiploxypinus doubingeri, P. fastidioides, Stellapollenites thiergartii, Tsugaepollenites oriens and Triadispora plicata reflects the final reforestation by Triassic gymnosperms such as Voltziales and pteridosperms (seedferns). A duration of 4 to 5 million years for the process of repopulation is inferred from radiometric dating. A first maximum in pollen diversity was reached during the early Anisian (Aegean), at the Solling-Ro¨t Formation transition in the Germanic basin (Szurlies 2007), but with pollen species diversity still only about half that of the Late Permian palynofloras from the Zechstein sequence (Fig. 1b). The Mid Triassic was, after the Permian– Triassic transition, the most dynamic time in the entire Triassic in terms of microfloral turnover. Many new pollen and spore types that become important during the later Triassic and even during the later course of the Mesozoic appeared at this time. After a brief decline during the Anisian, species diversity increased strongly throughout the Mid Triassic and reached a maximum in the late Ladinian (Longobardian) (Fig. 1a). The early Mid Triassic sediments (Lower Muschelkalk) were deposited in a marine environment, and their sporomorph assemblages indicate a dry palaeoclimate. Assemblages from the Middle Muschelkalk are composed almost exclusively of xeromorphic taxa (e.g. Triadispora) indicative of a very dry climate. The early Ladinian, represented largely by the Upper Muschelkalk, was characterized by semi-arid conditions, which resulted in spore species diversity remaining low while that of pollen steadily increased. More humid conditions during the deposition of the Lettenkeuper (Late Ladinian) are indicated by a dominance of Aratrisporites and Todisporites. It appears that, over the entire Triassic, the diversity of plants producing spores (Fig. 1c)

was largely controlled by the availability of water, while diversity among gymnosperms was also affected by other environmental and biotic factors. This can be inferred from pollen diversity (Fig. 1b), which remained high even during a subsequent arid phase in the early Carnian (Lower Gipskeuper). Some pollen with striking morphological characters that entered the European Triassic palynological record during the Anisian have been recorded from the Permian. These include the girdling monosaccate star-shaped pollen Stellapollenites (Visscher & Brugman 1981) and the multisaccate form Dyupetalum (Brugman 1983b) recorded from Permian deposits in India (Lele & Maithy 1963; Venkatachala et al. 1995), and the western Urals (Dyupina 1974), respectively. At present it is uncertain whether these morphological types reappeared independently in the Mid Triassic, or whether they evolved from Permian taxa that migrated into the European realm. The number of extinctions (red line in Fig. 1) increased with the last appearances of Dyupetalum vicentinense, Stellapollenites thiergartii, Strotersporites sp., Voltziaceaesporites heteromorphus, Illinites chitonoides and Concentricisporites spp. during the late Mid Triassic. The initial increase in microfloral diversity during the Late Anisian is caused by the entries of sporomorph taxa such as Apiculatisporites plicatus, Concentricisporites spp., Institisporites sp. and Perotrilites minor during the Late Anisian. The most pronounced increase in palynofloral diversity in the Triassic is, however, seen in the Late Ladinian (Longobardian), with high-diversity palynofloras extracted from fluviatile and lacustrine sediments of the Lettenkeuper in the Germanic Basin. The maximum in palynomorph diversity is mainly the result of a significant increase in the proportion of spore taxa, which originated from the first extensive wetland floras formed after the retreat of the Muschelkalk sea which resulted in a large deltaic system prograding from northeast to southwest (Kozur1976).Thesedimentaryenvironmentchanged from shallow marine, via brackish lagoonal to deltaic fluviatile facies (Hauschke & Heunisch 1990). A study by Brugman et al. (1994) on the Lettenkeuper revealed a palynofloral succession beginning with assemblages dominated by alete bisaccoid conifer pollen representing hinterland plant associations accompanied by Protodiploxipinus gracilis and Podosporites amicus and representing xerophytic coastal pioneers, together with Aratrisporites spp. representing mangrove vegetation. The influence of the hinterland vegetation decreases upwards, while mangrove, marsh and swamp elements, like the equisetophyte spore Calamospora become more abundant. However, the increase in hygrophytic elements may have been controlled

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by local environmental factors, rather than palaeoclimatic changes (Brugman et al. 1994). A decrease in salinity, from normal marine to brackish conditions, with a temporarily enhanced freshwater influx during the Muschelkalk regression, is supported by 87Sr/86Sr isotope records from the Germanic basin that deviate significantly from those in the Alpine realm (Korte et al. 2003). The Germanic basin Sr-isotope record suggests that a basin-wide increase in runoff and an input of radiogenic Sr resulted from greater monsoonal activity, rather than from marine regression without any contemporaneous changes in the hydrology. A change to hypersaline conditions during deposition of the upper Lettenkeuper is indicated by the presence of evaporitic tidal flats and inland sabkha deposits (Aigner & Bachmann 1989), which mark the end of the wet phase. During the Lettenkeuper (Ladinian) the number of spore taxa increased with the appearance of Anapiculatisporites, Asseretosporites gyrate, Aulisporites astigmosus, Camarazonosporites rudis, Echinitosporites iliacoides, Heliosaccus dimorphus, Keuperisporites baculatus, Lycopodiacidites and Porcellispora longdonensis. In addition to these spore taxa, Ovalipollis, a common element of Late Triassic microfloral assemblages, has its first occurrence at the base of the Ladinian. Another noteworthy development is that of the Circumpolles group with early forms such as Partitisporites spp. and Praecirculina granifer originating during the Ladinian. These new pollen types document the dawn of the evolving Cheirolepidaceae that later became an important conifer group that produced the morphogenus Classopollis (which includes also Corollina and Circulina) in the Late Triassic and most of the succeeding Mesozoic. With regard to the systematic nomenclature of the Classopollis-Corollina-Circulina complex, Traverse (2004, 2008) proposed according to the International Code of Botanical Nomenclature (Greuter et al. 2000) to conserve the morphogeneric name Classopollis against Corollina and Circulina, which are commonly used for the same fossil pollen types. The International Botanical Congress accepted the proposal in 2005. Palynofloral diversity, mainly that of spores, declined temporarily during the early Carnian as the result of an arid phase that is represented by evaporitic sediments of the Lower Gipskeuper, the assemblages from which are dominated by xeromorphic elements including taeniate bisaccates and Triadispora. Several characteristic sporomorphs, including Echinitosporites iliacoides, Heliosaccus dimorphus, Podosporites amicus, Striatoabieites balmei, Triadispora crassa, and T. plicata, have their last occurrence at this level while taxa such as Camerosporites secatus,

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Enzonalasporites vigens, Triadispora verrucata and Vallasporites ignacii enter the record. A second maximum in species diversity occurred later in the Carnian, similar to that in the Lettenkeuper, and was associated with a fluviatile/ lacustrine facies (Schilfsandstein). Wetlands alongside rivers and lakes supported a high-diversity, spore-producing plant community that yielded assemblages that are slightly dominated by halophytes (Aratrisporites spp.) and hygromorphics (A. astigmosus). It has been suggested that the climate was rather humid at this time (e.g. Simms & Ruffell 1989), but Visscher et al. (1994) emphasized that widespread humid environmental conditions could be ascribed to local high groundwater tables in a fluvial depositional setting, rather than to a climatically induced ‘pluvial event,’ and existence of a humid phase at this time remains controversial. In eastern North America there are deposits that may have formed under slightly more humid conditions but their exact age is uncertain (Olsen & Kent 2000). During the late Carnian (Upper Gipskeuper in NW Europe) palynofloral diversity starts to decline. Of note is the further diversification of the cheirolepidacean pollen (Circumpolles group) with the first appearances of forms such as Classopollis meyeriana, C. torosus, C. zwolinskae and Granuloperculatipollis rudis. The Norian is, at about 15 million years, the longest Triassic stage, but the pollen and spore assemblages are insufficiently known due to the scattered nature of the palynomorph record. It is probable that Peromonolites and Ricciisporites (R. tuberculatus, R. umbonatus) appear during this interval. Data from the Upper Gypsum and the Lower Rhaetic (sensu polonico) beds in Poland suggest that assemblages from the latest Carnian to earliest Norian, and to a lesser extent from the younger substages, are dominated by xeromorphic elements, particularly the circumpolles Classopollis and Granuloperculatipollis. Because of the incomplete Norian records, considerable uncertainty exists about the ranges of many so-called ‘typical’ Rhaetian (in part also earliest Liassic) sporomorphs, such as Cornutisporites spp., Limbosporites lundbladii, Perinosporites thuringiacus, Rhaetipollis germanicus, Semiretisporis spp., Triancoraesporites spp., and Zebrasporites spp., for which a late Norian appearance cannot be excluded. In general, palynofloral diversity (Fig. 1a) declines by some 50% between the early Carnian and the Norian, mainly as a result of a decrease in the numbers of pollen species (Fig. 1b). This is the second most severe loss in pollen species after the Permian-Triassic biotic crisis and occurred at a similar time of turnover among terrestrial reptile faunas (Tucker & Benton 1982; Benton 1991, 1993; see also review by Lucas & Tanner 2008).

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More diverse assemblages, with many new species and greater numbers of spores, reappear in Rhaetian deposits (Fig. 1c), but the pollen diversity remains rather stable (Fig. 1b). The palynomorph associations are dominated by spores of lycophytes and ferns and testify to an initially fairly dry but later increasingly humid palaeoclimate. At the Triassic –Jurassic transition, palynofloral diversity declines by about 20%, mainly as a result of a decrease in the numbers of spore species (Fig. 1c). The early Hettangian pollen and spore species richness is similar to that of the preRhaetian. In comparison to the marked palynofloral turnover at the Permian–Triassic transition and the end-Carnian decrease in palynofloral diversity, the end-Triassic biotic crisis appears to have hardly affected palynofloral species diversity in Europe. A recent detailed study of the Triassic –Jurassic key sections in the western Tethys realm even suggests an increase in diversity during the latest Rhaetian (Bonis et al. 2009a, b), which is in agreement with previous, but less detailed studies in Europe (e.g. Morbey 1975; Lund 1977; Achilles 1981; Batten & Koppelhus 1996). Some characteristic Late Triassic forms such as Ricciisporites tuberculatus and Zebrasporites interscriptus range into the Hettangian. Other taxa (e.g. Ovalipollis pseudoalatus, Limbosporites lundbladii, Rhaetipollis germanicus and Triancoraesporites spp.) disappear during the late Rhaetian. A recent study from the Germanic basin claimed evidence for a major floral turnover across the Triassic-Jurassic transition (van Schootbrugge et al. 2009). The high abundance of spores, such as Polypodiisporites polymicroforatus, Concavisporites spp., Osmundacidites, in the Upper Rhaetian Triletes beds have been interpreted as evidence for a fern proliferation after the dieback of the Triassic woody vegetation. The presented data are, however, ambiguous as previous studies (Brenner 1986; Heunisch 1996) showed that the quantitative composition of the palynological record is biased by marked facies changes in the Exter Fm. (Rha¨tkeuper). The sedimentary sequence starts with shallow marine Contorta beds, followed by fluvial-lacustrine Triletes beds, which are followed by shallow marine sediments (pre-planorbis beds and Psilonoten beds) in the lower Liassic, each of them bounded by dis-/unconformities (Nitsch 2005; Nitsch et al. 2005). The presence of typical Triassic pollen species in the Triletes beds (e.g. Schulz 1967; Achilles 1981) is in disagreement with a major extinction scenario. Nevertheless, a minor decline in palynomorph diversity towards the early Hettangian is also shown in Figure 1. No major palynofloral changes have been found that correlate precisely with the ratified TriassicJurassic boundary as defined by the first occurrence

of the ammonite Psiloceras spelae (von Hillebrandt et al. 2007). The only miospore with a FAD close to the base of the Early Jurassic may be Cerebropollenites (C. thiergartii) (Ku¨rschner et al. 2007; Bonis et al. 2009a, b). C. thiergartii occurs within the turn to more negative d13C values in the lower part of the main carbon-isotope excursion, well above the extinction level of Triassic biota but significantly below the lowest occurrence of the first Jurassic ammonite. Fisher & Dunay (1981) reported from the British Rhaeto-Liassic the entry of C. thiergartii in the Watchet area at the top of Richardson’s (1911) Paper Shale in the lower part of the Blue Lias. Moreover, C. thiergartii has been reported from lowermost Liassic sediments in the high latitudes, Greenland (Pederson & Lund 1980) and the Sverdrup Basin (Suneby & Hills 1988) and in the eastern Tethys realm, the Alborz Mountains in Iran (Achilles et al. 1984). After the Triassic –Jurassic transition, the sporomorph assemblages initially have a fairly ‘neutral’ aspect, particularly those from the Early Jurassic, which are characterized and dominated by long-ranging taxa such as Chasmatosporites, Classopollis meyeriana and C. turosus, Peromonolites, Quadraeculina anellaeformis, Retitriletes, scabrate-, echinate- and baculate triletes of the Acanthotriletes/Anemiidites/Osmundacidites (Baculatisporites) comaumensis wellmanii/Trachysporites-group, and some alete, non-taeniate bisaccates that first appear in the Late Triassic. The Early Jurassic shows a remarkable development in variety and numbers of what is here called the Stereisporites-plexus, a complex of species belonging to Annulispora, Dicyclosporis, Distcyclosporis, Polycingulatisporites, Rogalskaisporites and Stereisporites. When the palynological trends from the Germanic Basin are compared to those in the Alpine realm one may note the general resemblance of the composition of palynological assemblages. Only Late Anisian and Ladinian records may show some significant differences but these need further verification.

Palynostratigraphy Several biostratigraphic schemes based on palynomorphs have been proposed for the Germanic Triassic. These include the (sub) phases of Brugman (1986) and zones of Orłowska-Zwolin´ska (1983, 1985, 1988), Schuurman (1977), Reitz (1985, 1988) and Schulz (1994), and the informal subdivisions of Brugman (1983a), Van der Zwan & Spaak (1992), Heunisch (1999) and Schulz & Heunisch (2005). Visscher & Brugman (1981) reviewed palynofloras from independently-dated Tethyan sequences. Individual treatments of the Early

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

Triassic appear in, for example, Visscher & Brugman (1988), for the Middle Triassic in Van der Eem (1983), Brugman (1986), Go´cza´n & Oravecz-Scheffer (1993), Kustatscher & Roghi (2006) and Kustatscher et al. (2006), and for the Late Triassic in Klaus (1960), Mostler & Scheuring (1974), Dunay & Fisher (1978), Van der Eem (1983), Blendinger (1988), Visscher & Krystyn (1978), Schuurman (1979), Hochuli & Frank (2000), Cirilli & Roghi (1999), Roghi (2004), Ku¨rschner et al. (2007) and Bonis et al. (2009a, b). The following part of this review is based on a study of the palynostratigraphy in NW Europe (Herngreen 2005). Zones are, with the exception of the lowest Triassic palynozone, based on first occurrences of marker species. A correlation of the palynostratigraphic zones and subzones in Europe is presented in Figure 2. A range chart with all major FAD’s and LOD’s of palynomorphs is shown in Figures 3a and 3b for the Early and Mid Triassic as well as the Late Triassic, respectively. The following palynomorph zones can be distinguished:

Lueckisporites virkkiae Zone Fijałkowska (1994a) The Permian palynofloras of NW Europe belong to the Euramerian province; this region is now mainly represented by Europe and North America (Warrington 1996a). All western European assemblages are assigned to the Lueckisporites virkkiae Zone. The eponymous zonal index species is a typical component and was thought to be restricted to Late Permian associations from the Zechstein succession in western Europe and time-equivalent strata in the Alps (e.g. Klaus 1963; Clarke 1965a, b; Visscher 1971). However, more recent studies (e.g. Warrington & Scrivener 1990; Edwards et al. 1997) have shown records of L. virkkiae from beds that are now known to pre-date the Illawara Reversal in the Middle Permian. In Russia it appears in rocks of similar age (see discussion in Warrington 2008). The zone includes, in ascending order (Fijałkowska 1994a), assemblage I with L. virkkiae Ab, assemblage II with L. virkkiae Ac and assemblage III with L. virkkiae Bc. Bisaccate pollen are dominant; L. virkkiae is the most striking and characteristic component and occurs with Lunatisporites, Protohaploxypinus, Striatoabieites and rare Vittatina. Non-taeniate bisaccates include representatives of the AlisporitesFalcisporites plexus, Jugasporites, Limitisporites and Klausipollenites. Monosaccates, including Nuskoisporites and Potonieisporites, are scarce, and spores, monosulcates (Cycadopites) and Ephedripites are usually rare.

269

This zone correlates with phase LT-1 of Brugman (1983a), the Densoisporites playfordii–Endosporites papillatus Zone of Reitz (1985, 1988) and the GTr2 zone of Heunisch (1999). The age of the zone is Late Permian.

Lundbladispora obsolete – Protohaploxypinus pantii Zone Orłowska-Zwolin´ska (1985) The Lower Buntsandstein interval (Calvo¨rde and Bernburg formations) preceding the Densoisporites nejburgii Zone is poor in sporomorphs but is characterized by the absence of Jugasporites delasaucei, Lueckisporites virkkiae, Nuskoisporites dulhuntyi and Vittatina, which are Permian markers. Some taeniate bisaccate pollen, such as Lunatisporites, Protohaploxypinus and Strotersporites, and Cycadopites spp., range from the Permian into the Triassic. Characteristic of this zone, although not restricted to it, are the spores Densoisporites playfordii, Endosporites papillatus and Lundbladispora willmottii, but few assemblages have been observed. The zone is not based on first appearance(s) of species, and, though the nominate species have a wider range (Reitz 1985, 1988), the zonal name as proposed by Orłowska-Zwolin´ska (1985) is used here. This zone correlates with phase LT-1 of Brugman (1983a), the Densoisporites playfordii – Endosporites papillatus Zone of Reitz (1985, 1988) and the GTr2 zone of Heunisch (1999). The age of this interval falls within the Induan.

Densoisporites nejburgii Zone Orłowska-Zwolin´ska (1977, 1984, 1985) Middle Buntsandstein assemblages are usually characterized by high numbers of Densoisporites nejburgii, the nominate species of this zone. The age of this zone is probably Late(?st) Dienerian (Late Induan) and Smithian (Early Olenekian). Orłowska-Zwolin´ska (1977, 1984, 1985) described three subzones, in ascending order: the Densoisporites nejburgii–Acritarcha Subzone, the Densoisporites nejburgii Subzone and the Cycloverrutriletes presselensis Subzone. Densoisporites nejburgii–Acritarcha Subzone. The base of this subzone is defined by the FAD of the nominate species D. nejburgii and low numbers of spores (D. playfordii and Endosporites papillatus) in contrast to relatively common, sometimes abundant, spinose acritarchs assigned to Micrhystridium and Veryhachium and to a lesser extent Leiofusa. This assemblage is found mainly in the Volpriehausen Formation, but may occur in the lowermost part of the succeeding Detfurth Formation. This subzone correlates with phase LT-2 of Brugman (1983a), the

(pars)

This paper / Herngreen 2005 Heun. Pollen zone P. subzone 1999 Lias α1

Keuper

Arnstadt Fm.

L. lundbladii

C. meyeriana 15

Weser Fm. Stuttgart Fm.

C. secatus

A. astigmosus Tr. verrucata

PERMIAN LATE

Changhsingian

Muschelkalk Buntsandstein Zechstein

MIDDLE EARLY

TRIASSIC

Erfurt Formation Ladinian

Induan

R. tuberculatus

16

Grabfeld Fm.

Olenekian

20 19 18 17

G. rudis

Middle (Gipskeuper)

Carnian

Anisian

Brugman (1983a) Brugman et al. (1994)

C. thiergartii R. germanicus

Exter Fm.

OrlowskaZwolinska (1984)

H. dimorphus

Warburg Fm.

10

Upper

Thiergartiivicentinensis

P. doubingeri

8

Perotrilites minor

thiergartii-oriens

P. leschikii

7 6

P. fastidioides Voltziaceaesporites heteromorphus

C. presselensis

5

C. presselensis

LT - 4

D. nejburgii

4

D. nejburgii

LT - 3

D. nejburgii-Ac.

3

Acritarcha D. nejburgii

LT - 2

2

L. obsoleta Pr. pantii

LT - 1

T. crassa Verrucosisp.

Middle

Bernburg

Bröckelschiefer

Transition phase

D. nejburgii

Volpriehausen

Calvörde

gracilis-perforatus

crassa-thiergartii

Hardegsen

Lower

perforatus-dimorphus

Tsugaepollenites oriens

Röt

Detfurth

dimorphus-iliacoides Heliosaccus dimorphus

9

S. thiergartii

Solling

A. astigmosus P. longdonensis

Institisporites

Anhydrit Folge Jena Formation

14 13 12 11

L. obsoleta P. pantii L. virkkiae

Fig. 2. Comparison of previous Triassic palynostratigraphic subdivisions in Central and NW Europe.

1

¨ RSCHNER & G. F. WALDEMAAR HERNGREEN W. M. KU

LATE

Rhaetian

Norian

Palynostratigraphy

Germanic Basin Lias

JURASSIC Early Hettangian

270

Lithostratigraphy Chronostratigraphy

(a) Age

Period

Standard Chronostratigraphy Epoch Stage

Spores & Pollen Alpine realm & Germanic basin

Substage

Enzonalasporites vigens Duplicisporites spp. (common) Partitisporites spp. (common)

Grabfeld Fm. (Lower Gypsum Keuper) without ‘Estheria’ Beds

230

W. magnus

E. iliacoides (T) P. fastidioides

Erfurt Fm. H. dimorphus

Ladinian

Partitisporites novimundanus, Ovalipollis pseudoalatus Staurosaccites quadrifidus Duplicisporites granulatus Echinitosporites iliacoides

235

Fassanian

I. chitonoides (T)

Warburg Fm. Heliosaccus dimorphus St. thiergartii, Ts. oriens

Middle

D. vicentinensis (T)

Meißner Fm.

Kuglerina meieri? Camarozonosporites rudis

Illyrian 240

Triassic

Institisporites Trochitenkalk Fm. Institisporites sp. Podosporites amicus

Heilbronn Fm.

Anisian

Schaumkalk Mbr. Pelsonian

S. thiergartii P. doubingeri

Jena Fm

I. chitonoides (T)

Bithynian 245 Aegean

Röt Fm. Solling Fm.

Spathian Olenekian Early

250

Induan

P. leschikii T. crassa -Verruc. C. presselensis

Hardegsen Fm. Detfurth Fm.

D. nejburgii D. nejburgii

Smithian Dienerian u. Griesbachian (Gangetian)

Dyupetalum vicentinensis (T)

Volpriehausen Fm.

D. nejburgii acritarch acme

Tsugaepollenites oriens P. doubingeri (common)

I. kosankei,

Triadispora spp. acme

P. leschikii, R. jenensis, D. nejburgii

I. kosankei, P. fastidioides, J. conmilvinus Angustisulcites spp., I. chitonoides, Stellapollenites thiergartii D. nejburgii (common), Triadispora crassa, J. conmilvinus D. playfordii, C. presselensis Cycloverrutriletes presselensis Platysaccus leschikii Densoisporites nejburgii

Bernburg Fm. Calvörde Fm.

L. obsoleta P. pantii L. virkkiae

L. virkkiae, Jugasporites delasaucei, Nuskoisporites dulhuntyi, Vittatina

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

Infernopollenites spp. Aulisporites astigmosus

Longobardian

Fig. 3. Range chart for the Early and Mid Triassic (a) and Late Triassic (b) with the main FAD’s & LOD’s of palynomorphs as discussed in this review in the Germanic basin and Alpine realm. Where ranges of palynomorphs in the Alpine realm differ significantly from those in the Germanic basin they are indicated with ‘(T)’ ¼ Tethys. The timescale has been made with TSCreator version 4.0 visualization program at www.tscreator.org, June 2009. (Ogg, J. G. and ICS Subcommission for Stratigraphic Information). 271

272

(b) Standard Chronostratigraphy Age

Period

Epoch

Stage

Jurassic

Early

Hettangian

Substage

Lithostratigraphy Germanic Triassic (from Bachmann & Kozur 2005)

Pollen Zonation (this paper)

Spores & Pollen Alpine realm & Germanic basin Cerebropollenites thiergartii, Ischyosporites levigatus L. lundbladii, C. rhaeticus, O. pseudoalatus, R. germanicus

200

Exter Fm.

Rhaetian

L. lundbladii

Triletes beds R. germanicus

Contorta beds Stubensandstein

Cingulizonates rhaeticus, Limbosporites lundbladii

G. rudis Enzonalosporites vigens, V. ignacii, Patinasporites densus

lower & middle Postera beds

205 Sevatian

G. rudis Alaunian 210

Loewenstein Fm., pars (1.-3. Stubensdst. + 1.-3. Hangendletten, uppermost Obere Bunte Mergel)

Norian

Triassic

Perinopollenites elatoides, M. fuscus, Heliosporites reissingeri

Arnstadt Fm. with common Steinmergel

Lacian

Late

215 Classopollis zwolinskae, Chasmatosporites spp., G. rudis (common)

Mainhard Fm.

Duplicisporites granulatus, Camerosporites secatus, T. verrucata

Heldburggips Classopollis spp.

Hassberge Fm.

220

Steigerwald Fm.

Infernopollenites spp.

Tuvalian

Lehrb. beds Rote Wand

Weser Fm. (Upper Gypsum Keuper)

R. tuberculatus

C. secatus

Carnian Schilfsandstein Fm. (Stuttgart Fm.)

A. astigmosus

225

Partitisporites novimundanus, Aulisporites astigmosus, Lagenella martinii, Praecirculina granifer, Staurosaccites quadrifidus Partitisporites quadruplices

Infernopollenites spp.

Julian Kuglerina meieri (T)

upper Grabfeld Fm. (”Estheria” beds)

Fig. 3. Continued.

Tr. verrucata

Patinasporites densus, Camarozonosporites rudis, Vallasporites ignacii T. verrucata Camerosporites secatus

E. iliacoides, A. klausii Heliosaccus dimorphus

¨ RSCHNER & G. F. WALDEMAAR HERNGREEN W. M. KU

Q. anelliformis, Rhaetipollis germanicus

Postera Sdst.

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

Micrhystridium –Veryhachium Zone of Reitz (1985, 1988) and the GTr3 zone of Heunisch (1999). Densoisporites nejburgii Subzone. This subzone is defined by the abundance and general dominance of the nominate species. Densoisporites playfordii is common in this subzone, which also includes the FADs of Platysaccus leschikii and Punctatisporites triassicus. The assemblages occur from the upper part of the lower Detfurth Formation to the middle of the Hardegsen Formation. This subzone correlates with phase LT-3 of Brugman (1983a), the nejburgii–heteromorphus phase (Reitz 1985, 1988; Brugman 1986) and the GTr4 zone of Heunisch (1999). Cycloverrutriletes presselensis Subzone. The base of this zone is defined by the FAD of the nominate species, which occurs regularly and in high numbers; Densoisporites and Lundbladispora remain common. The assemblages are restricted to the upper part of the Hardegsen Formation. This subzone correlates with the Cycloverrutriletes presselensis zone of Reitz (1985, 1988), the GTr5 zone of Heunisch (1999), phase LT-4 of Brugman (1983), and the lower part of the heteromorphusconmilvinus phase of Brugman (1986).

Triadispora crassa – Verrucosisporites Zone Herngreen (2005) New elements in assemblages from the Solling Formation are Triadispora crassa and relatively common Verrucosisporites. The zone is defined as the interval between the FAD of Triadispora crassa and those of Illinites chitonoides, I. kosankei, Protodiploxypinus doubingeri, P. fastidioides and Stellapollenites thiergartii. The FAD of Angustisulcites klausii and the incoming of common Voltziaceaesporites heteromorphus occur in this zone. The zone is the lower part of the Voltziaceaesporites heteromorpha Zone of Orłowska-Zwolin´ska (1984, 1985, 1988), which was defined as the interval between the FAD of the nominate species in the Ro¨t Formation, as it was known at that time, and the informal unit with common Perotrilites minor (Orłowska-Zwolin´ska 1984). This part was formalized as the Perotrilites minor Zone (Orłowska-Zwolin´ska 1985, 1988), at the base of the Muschelkalk. Later it was shown that the FAD of V. heteromorphus is in the Detfurth Formation and that this taxon becomes very abundant in the Solling Formation (Brugman 1983a, 1986). This is the transition from LT-4 to the crassa-thiergartii phase of Brugman (1983a) and the upper part of his heteromorphus– conmilvinus and the conmilvinus –crassa phases (Brugman 1986). A three-fold division of the Solling interval proposed by Reitz

273

(1985, 1988) on the basis of different ratios of Verrucosisporites spp. is not applicable on a regional scale, nor are the two subzones of Schulz (1994). The age of the zone, which is restricted to the Solling Formation, is Spathian (Late Olenekian).

Stellapollenites thiergartii Zone Herngreen (2005) The zone is defined by the range of Stellapollenites thiergartii. At or near the base of this zone, which correlates with the base of the Ro¨t, are the FADs of Angustisulcites grandis, Apiculatasporites plicatus, Concentricisporites, Illinites chitonoides, I. kosankei, Protodiploxypinus doubingeri, and P. fastidioides. Tsugaepollenites oriens appears higher in the Ro¨t, and in the uppermost part (Ro¨t 4) Protodiploxypinus becomes common and Triadispora dominant. Densoisporites nejburgii disappears in this zone and is replaced by various species of the lycopod spore Aratrisporites. The LOD of Tsugaepollenites oriens is almost at the top of the zone and coincides with the last definite occurrence of S. thiergartii, of which younger records are questionable. The zone is found from the Ro¨t up to the middle part of the Upper Muschelkalk and is correlated with the Anisian. It correlates with the Perotrilites minor and Tsugaepollenites oriens zones of Orłowska-Zwolin´ska (1985) and includes the informal zones GTr uppermost 6 and 7 –9 of Heunisch (1999).The following subdivisions of this zone have been proposed (Herngreen 2005): Platysaccus leschikii Subzone. The top of this subzone is marked by the LOD of Platysaccus leschikii. Protodiploxypinus doubingeri and P. fastidioides occur consistently in low numbers, and Triadispora, mainly T. crassa, becomes common and is generally dominant. The subzone occurs in the Ro¨t and is correlated with the Aegean (earliest Anisian). It correlates with the P. fastidioides Subzone of the Voltziaceaesporites heteromorpha Zone of Orłowska-Zwolin´ska (1985), the Triadispora spp.–Stellapollenites thiergartii (Salinarro¨t) and Concentricisporites nevesi-Triadispora plicata subzones (postsaline Ro¨t 1–3) of Reitz (1985, 1988), and the P. doubingeri–V. heteromorphus interval of Schulz (1994). Protodiploxypinus doubingeri Subzone. The base of the subzone is at the FAD of Tsugaepollenites oriens and is marked by the appearance of common Protodiploxypinus doubingeri. This subzone spans the uppermost Ro¨t to lower Middle Muschelkalk succession and is latest Aegean to earliest Illyrian in age. It correlates with the Perotrilites minor Zone of Orłowska-Zwolin´ska (1985, 1988),

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¨ RSCHNER & G. F. WALDEMAAR HERNGREEN W. M. KU

who distinguished the Tsugaepollenites oriens Zone on the basis of the association of the nominate species with Perotrilites minor and common Angustisulcites, Protodiploxypinus and Triadispora, and an absence of acritarchs. Reitz (1985) introduced the Protodiploxypinus sittleri–Illinites chitonoides (Ro¨t 4) and Osmundacidites senectus –Perotrilites minor and Aratrisporites fimbriatus –A. tenuispinosus zones (Lower Muschelkalk). Brugman (1986) proposed the informal phases Ro¨t– 3 and Mu –1 for this subzone, which also correlates with the lower part of the thiergartii–oriens phase of Brugman et al. (1988). The Protodiploxypinus doubingeri Subzone correlates with zone GTr 8 of Heunisch (1999). In the Southern Alps the thiergartii– vicentinensis phase was proposed for the Pelsonian to Illyrian (Brugman 1983a). Institisporites Subzone. The base of this subzone is defined by the FADs of Institisporites sp. and Podosporites amicus, and the LOD of the eponymous taxon defines its top. The first occurrence of Doubingerispora filamentosa is at or near the base of this subzone, and the last common occurrence of Protodiploxypinus doubingeri is in its basal part. The LOD of Tsugaepollenites oriens is near the top of the subzone. It is found in the upper Middle to middle Upper Muschelkalk and is correlated with the Illyrian. The subzone correlates with the Uvaesporites gadensis-Doubingerisporites filamentosa Zone of Reitz (1985), Mu –2 of Brugman (1986), and GTr 9 of Heunisch (1999). It should be emphasized that the middle part of the Upper Muschelkalk (mo2) is difficult to characterise palynologically. In this sequence the last irregular and rare occurrences of Stellapollenites thiergartii and the first spasmodic and scarce occurrences of Heliosaccus dimorphus are found, but the S. thiergartii–H. dimorphus zone boundary cannot be placed with certainty. Herngreen (2005) favoured the presence of Institisporites sp., the LOD of which coincides with that of S. thiergartii, as a definition of this zone. In the Alpine Triassic this interval correlates with the upper part of the thiergartii– vicentinensis phase (Brugman 1983a, 1986).

Heliosaccus dimorphus Zone Orłowska-Zwolin´ska (1983, 1985, 1988), emended Herngreen (2005) The base of the zone is marked by the FAD of consistent Heliosaccus dimorphus, which probably also coincides with that of Retisulcites perforatus. At the base of the zone the eponymous species occurs only sporadically. It is accompanied by Anapiculatisporites telephorus, Camarozonosporites rudis and Institisporites sp. Slightly higher are the FADs of

Echinitosporites iliacoides, Ovalipollis pseudoalatus, the Partitisporites novimundanus morphon, and probably Cordaitina minor, which all appear at the base of the upper part of the Upper Muschelkalk (mo3), followed by the FAD of Anapiculatisporites spiniger in the topmost Upper Muschelkalk and those of Aulisporites astigmosus, Keuperisporites baculatus and Lycopodiacidites at the base of the Lettenkeuper. The LOD of H. dimorphus coincides with the transition from associations with the last common Aratrisporites spp. to those with the first common Porcellispora longdonensis. Just below the LOD of H. dimorphus are the FADs of Camerosporites secatus, Enzonalasporites vigens and Vallasporites ignacii. This zone occurs from the upper Upper Muschelkalk to the transition from the Lettenkeuper to the lowermost Lower Gipskeuper and is correlated with the Ladinian. It corresponds with zone I Heliosaccus dimorphus and the lowermost part of IIa Echinitosporites iliacoides Subzone of the Conbaculatisporites longdonensis Zone of Orłowska-Zwolin´ska (1983, 1985, 1988). Reitz (1985) used the terms Aratrisporites spp. – Minutosaccus spp., Leschikisporites aduncus – Heliosaccus dimorphus and Echinatosporites iliacoides –Nevesisporites lubricatus Zone for this interval. The zone includes the uppermost Muschelkalk (Protodiploxypinus) gracilis- (Retisulcites) perforatus phase, the Lettenkeuper (Retisulcites) perforatus –(Heliosaccus) dimorphus phase, and the (Heliosaccus) dimorphus –(Echinitosporites) iliacoides phase of Brugman et al. (1994) in the Lower Myophorienschichten. In Heunisch (1999), this interval includes the zones GTr10-11.

Camerosporites secatus Zone Herngreen (2005) The base of this zone is defined by the FAD of Camerosporites secatus, which coincides with those of Enzonalasporites vigens, Triadispora verrucata and Vallasporites ignacii, and the first common occurrence of Ovalipollis pseudoalatus. The zone corresponds with the range of the eponymous species. Slightly above its base, in the upper part of the Lower Gipskeuper, are the LODs of Retisulcites perforatus and Striatoabieites balmei and the FAD of regular (and often common) Triadispora verrucata, which ranges into the middle of the Upper Gipskeuper. At the top of the zone, C. secatus may occur with rare G. rudis and Classopollis. The zone is found in the Middle Keuper, specifically the Lower Gipskeuper to lower Steinmergelkeuper, and is correlated with the Carnian. It corresponds with zones II Conbaculatisporites longdonensis, III Aulisporites astigmosus and

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

subzone IVa, the lower part of the Corollina meyeriana Zone of Orłowska-Zwolin´ska (1983, 1985, 1988); this was the Porcel(l)ispora longdonensis Zone in her 1985 paper. Reitz (1985) used the terms Retisulcites perforatus-Eucommiidites microgranulatus and Partitisporites quadruplicis-Patinasporites densus zones in the Lower Gipskeuper. The Camerosporites secatus Zone also corresponds with zones GTr 12–15 of Heunisch (1999). The Porcellispora longdonensis and Aulisporites astigmosus zones of Orłowska-Zwolin´ska (1983, 1985, 1988) have been emended by Herngreen (2005) and assigned subzonal status, as follows: Triadispora verrucata Subzone. The base of this subzone is defined by the FAD of the nominate species Triadispora verrucata and extends upwards to include its rapid increase. OrłowskaZwolin´ska (1983, 1985, 1988) restricted her subzone to the acme of the species. However, the level, within the higher part of the Lower Gipskeuper, at which this occurs is not precisely known and for this reason the unit has been extended downwards to the first appearance of T. verrucata in the Grenzdolomit. In practice, the subzone is found in the Lower Gipskeuper. It correlates with the Conbaculatisporites longdonensis, later Porcellispora longdonensis, Zone of Orłowska-Zwolin´ska (1983, 1985, 1988) and zones GTr 12–13 of Heunisch (1999). Aulisporites astigmosus Subzone. This subzone corresponds with the acme of Aulisporites astigmosus and includes the LOD of Podosporites amicus. Throughout the Germanic facies the eponymous species is very common in the Schilfsandstein; the subzone is correlated with the Julian. OrłowskaZwolin´ska (1983, 1985, 1988) gave assemblages with common Aulisporites astigmosus zonal rank. However, because it may constitute up to 10% in assemblages from the Lettenkeuper it is preferred to give the interval with the real acme (.30 up to 75%) in the Schilfsandstein the rank of subzone. The subzone correlates with Gibeosporites lativerrucosus-Aulisporites astigmosis Zone of Reitz (1985) and zone GTr 14 of Heunisch (1999).

Granuloperculatipollis rudis Zone Herngreen (2005) The base of this zone is placed at the FAD of common Granuloperculatipollis rudis. In this zone the circumpolles Classopollis meyeriana, C. zwolinskae and Granuloperculatipollis rudis, become common. Riccisporites tuberculatus and R. umbonatus probably first appear in the upper half of the zone, which is found in the middle and

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upper Steinmergelkeuper and is correlated with the Norian. Due to a paucity of sequences with good palynological assemblages from this interval, the FAD of many species that occur in the Rhaetian remains uncertain. The Granuloperculatipollis rudis Zone probably corresponds with the Corollina Porcellispora Subzone, the lower part of the Corollina Enzonalasporites Zone of Lund (1977), zones GTr 16 –17 of Heunisch (1999) and Subzone IVb of the Corollina meyeriana Zone (Orłowska-Zwolin´ska, 1983, 1985). Subzone IVb is thought to have its base in the Drawno Member (Orłowska-Zwolin´ska 1983) or Jarkowo Member (Orłowska-Zwolin´ska 1985), and the upper boundary is at the transition from the lower to the upper part of the Zbaszyn Member, the Lower Rhaetic sensu polonico. The IVb subzone assemblages are similar to those of the Steinmergelkeuper (Schulz 1967, 1996; Orłowska-Zwolin´ska 1985).

Rhaetipollis germanicus Zone Herngreen (2005), emended This zone corresponds to the range of Rhaetipollis germanicus. The base of this zone is defined by the FAD of the nominate species. In the Germanic Basin R. germanicus has been observed in the upper Postera beds (Lund 1977, 2003; Schulz & Heunisch 2005 and literature cited herein), which are assigned to the Lower Rhaetian. But, in Britain R. germanicus has its lowest occurrence in the Twyning Mudstone Fm. of the St. Audrie’s Bay section which was previously assigned to the Norian (Hounslow et al. 2004), but correlates according to a recently revised palaeomagnetic and cyclostratigraphic age assessment to the Early Rhaetian (Hu¨sing et al. 2010) R. germanicus has not yet been reported from the underlying Steinmergelkeuper. In the Alpine realm several previous palynological studies on Rhaetian reference sections, such as Kendelbach and Weissloferbach (e.g. Morbey 1975; Morbey & Neves 1974; Schuurman 1979), have suggested that the presence of R. germanicus may be a valuable biostratigraphic marker to characterize the Rhaetian. However, the FAD of R. germanicus is still uncertain here because of the predominant carbonate sediments (e.g. Plattenkalk, Po¨tschenkalk) underlying the Rhaetian deposits (Ko¨ssen or Zlambach beds). Unfortunately, palynological samples of the Steinbergkogel section near Hallstatt, which is the GSSP candidate for the base of the Rhaetian, are barren (Krystyn et al. 2007). However, R. germanicus has been recorded (Ku¨rschner, unpublished data) from the basal part of the Zlambach Formation and the Pedata-Kalk, which is lower Rhaetian (Sevatian 2, Misikella

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posthernsteini/Paracochloceras niveau). This finding is in agreement with the range that has been also proposed by observations from the Germanic Basin (e.g. Lund 1977, 2003; Schulze & Heunisch 2005), but the exact FAD of R. germanicus from independently-dated marine Norian-Rhaetian sediments has still to be documented. Assemblages from this zone are very diverse and include many new sporomorphs, for example Cingulizonates rhaeticus, Cornutisporites rugulatus, C. seebergensis, some Densosporites species, Heliosporites reissingeri, Limbosporites lundbladii, Quadraeculina anellaeformis, Semiretisporis gothae, S. wielichoviensis, Triancoraesporites ancorae, T. reticulatus and Zebrasporites laevigatus. Assemblages from the upper part of the zone are often dominated by Classopollis meyeriana, Classopollis spp. and Peromonolites elatoides. The zone is restricted to the Ra¨tkeuper and equivalent NW European Rhaetian strata, such as the Westbury Formation and Cotham Member in the UK (Orbell 1973; Morbey 1975; Warrington et al. 1995), the ‘middle and upper Rhaetic’ sensu germanico in the former German Democratic Republic (Schulz 1967) and Denmark (Lund 1977), the Upper Rhaetic sensu polonico (OrłowskaZwolin´ska 1983, 1985; Fijałkowska-Mader 1999), and the Ko¨ssener Schichten and the lower part of the Pre-planorbis Beds in the Alpine facies (Morbey 1975). For the lower part of this interval, Orłowska-Zwolin´ska (1983, 1985) introduced Subzone IVc (defined by the appearance of R. germanicus) of the Corollina meyeriana Zone and the Ricciisporites tuberculatus Zone (dominated by the eponymous species together with taxa such as C. seebergensis, L. lundbladii, S. gothae and T. ancorae) for the microfloras of, respectively, the upper part of the Zbaszyn Member and the Wielichowo Member in Poland. The Rhaetipollis germanicus zone correlates with the Corollina Conbaculatisporites Zone (Lower Rhaetian sensu germanico), the Rhaetipollis Limbosporites Zone (Contorta Schichten) and the Ricciisporites Polypodiisporites Zone (Triletes Schichten) of Lund (1977, 2003), and the GTr 18–20 of Heunisch (1999). Heunisch (1999) defined the top of zone 18 by the LOD of Ricciisporites umbonatus, the top of zone 19 by the LOD of C. zwolinskae, and zone 20 is characterized by Densosporites and Polycingulatisporites. The age of the R. germanicus zone is Rhaetian. Limbosporites lundbladii Subzone, new. The base of this subzone is defined by the FAD of Limbosporites lundbladii and the LOD of C. zwolinskae. The FADs of Cingulizonates rhaeticus and Semiretisporis spp. and the LODs of Enzonalasporites spp. and Granuloperculatipollis rudis are at or near the base of this

subzone, which is restricted to the middle part of the German Rhaetian, the ‘Middel-Rha¨t’. The subzone correlates with the Rhaetipollis Limbosporites Zone of Lund (1977) and zone GTr 19 of Heunisch (1999). In the Alpine realm this subzone correlates with the MI miospore subzone of Morbey (1975), with Phase 3 of Schuurman (1979) and the Rhaetipollis Limbosporites Zone of Ku¨rschner et al. (2007). The age of this zone is middle and late Rhaetian. Although no further prevalent miospore first appearance dates from uppermost Rhaetian deposits are known, the NW European palynological records show marked quantitative changes, which have been used to establish palynological assemblage zones with a significant finer stratigraphic resolution. It has been shown that they are useful at least for regional correlation within the Alpine realm (Schuurman 1979; Ku¨rschner et al. 2007; Bonis et al. 2009). The Rhaetipollis–Limbosporites Zone of Ku¨rschner et al. (2007) is followed by the Rhaetipollis–Porcellispora zone (RPo zone), which is characterized by a marked increase in spore diversity and abundance, particularly those of Porcellispora longdonensis, Polypodiisporites polymicroforatus and Calamospora tener show an acme. There is a brief decline in Classopollis and an increase in Vitreisporites in the lower part of the zone. Rhaetipollis germanicus disappears in the upper part of this zone. The Rhaetipollis– Porcellispora zone (RPo zone) is followed by the Trachysporites –Porcellispora zone (TPo zone): This zone is characterized by a continuous decline in Classopollis. Spore assemblages show a decrease in C. tener, Deltoidospora and P. polymicroforatus, accompanied by an increase of Carnisporites, Concavisporites, P. longdonensis and Trachysporites fuscus. The RPo and TPo assemblage zones together correspond to Phase 4 of Schuurman (1979) and the Ricciisporites –Polypodiisporites zones of Lund (1977).

Cerebropollenites thiergartii Zone, new zone The base of this zone is defined by the FAD of the nominate species and Ischyosporites variegatus. The transitional interval from the top Rhaetian into the Liassic (pre-Planorbis beds) is characterized by marked changes in the quantitative composition of pollen and spore assemblages, but the only miospore with a FAD close to the base of the Hettangian as defined by the first appearance of Psiloceras spelae is Cerebropollenites thiergartii. As the FAD of C. thiergartii predates the entry level of the P. spelae by about 2 m at the GSSP stratotype Kuhjoch (Bonis et al. 2009), the lowermost part of the C. thiergartii Zone is still of latest Rhaetian age. Typical Rhaetian spore types such as Rhaetipollis germanicus and Ovalipollis pseudoalatus are

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE

absent. The pollen and spore assemblages of this zone are characterized by a marked increase of Heliosporites reissingeri and Trachysporites fuscus while Ricciisporites tuberculatus and Polypodiisporites polymicroforatus are present only in very low abundances. The C. thiergartii Zone can be further divided into two pollen assemblage zones, the Trachysporites–Heliosporites zone (TH zone) and the Trachysporites–Pinuspollenites zone (TPi zone) of Ku¨rschner et al. (2007). The C. thiergartii zone corresponds to Phase 5 of Schuurman (1979) and the Pinuspollenites Trachysporites Zone of Lund (1977). The age of this zone is latest Rhaetian and early Hettangian (Lias a1).

Correlation of the Germanic Triassic with the Tethyan scale and palynostratigraphy of the Triassic stage boundaries (GSSP’s) in northwest Europe The base of the Triassic in the predominantly continental Germanic Triassic has, in the absence of the primary boundary indicator, a conodont, been identified by other stratigraphic markers, such as conchostracans (Kozur 1993), carbon isotope stratigraphy (Korte & Kozur 2005) and magnetostratigraphy (Szurlies et al. 2003; Szurlies 2004, 2007). The Permian–Triassic boundary is placed within the Calvo¨rde Formation, at the top of the Graubankzone in North Germany (Kozur 1999a). The very rare and poor pollen and spore assemblages from the uppermost Sandy Claystone Member of the Calvo¨rde Formation to the lower part of the upper Bernburg Formation belong to the Lundbladispora willmottii–Lunatisporites hexagonalis Zone (Kozur 1999a), which is assigned to the Induan. The base of these very poor associations, which occur also in the uppermost Bernburg Formation and in the Volpriehausen Formation, is provisionally correlated with the Induan–Olenekian transition. Assemblages dominated by the eponymous species of the Densoisporites nejburgii Subzone (upper lower –upper Detfurth and Hardegsen 1 and 2) also occur in Hungary, in the Csopak Marl Member (Brugman 1986) and the upper part of the Hidegku´t Formation (Go´cza´n et al. 1986). Go´cza´n et al. (1986) placed the Induan–Olenekian boundary at the base of their nejburgii–reductum dominance zone. The Csopak Marl contains Tirolites cassianus, which indicates a late Olenekian (earliest Spathian) age (Kozur 1999b). The Hardegsen 3 and 4 assemblages are very similar but also contain Cycloverrutriletes presselensis; a direct correlation is not possible, but an early Spathian age seems likely. Assemblages from the overlying Solling Formation are characterized by regular and

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common Triadispora crassa and Voltziaceaesporites heteromorphus, which indicate a correlation with the uppermost Csopak Marl of latest Olenekian (late Spathian) age (Kozur 1999b). Bachmann & Kozur (2004) assigned an earliest Olenekian (early Smithian) age to the top of the Bernburg Formation. They also assigned an early Olenekian age to the overlying Volpriehausen Formation and dated the succeeding Detfurth, Hardegsen and Solling formations as Late Olenekian (Spathian). In recent years 4 out of 6 Triassic stages boundaries (base of the Anisian, Ladinian, Carnian, Rhaetian) have been defined (or proposed) in the Alpine realm as discussed in this chapter. The Des¸li Caira Hill section in the northern part of Dobrogea, a province in Romania, has been proposed as the GSSP for the base of the Anisian (Gra˘dinaru et al. 2006). However, no palynological data have been published from this section. A rich palynological record is unlikely because of the development of this section in Hallstatt-type facies. However, with regard to the correlation of the mid-Triassic of the German Triassic with the Tethyan chronology some important results have been obtained from palynological studies. The appearance of Stellapollenites thiergartii in the Ro¨t is correlatable with its FAD in the lower Aszo´fo¨ Dolomite Formation in Hungary (Brugman 1986; Go´cza´n et al. 1986) and indicates the base of the Anisian (earliest Aegean). The Myophoria Beds in the uppermost Ro¨t Formation were assigned an early Bithynian age by Kozur (1999b). In contrast, Brugman (1986) had assigned the upper Ro¨t to the Bithynian and early? Pelsonian, while Visscher et al. (1993) correlated the upper Ro¨t and entire Lower Muschelkalk to the Pelsonian. Conodont and ammonoid data from the Lower Muschelkalk indicate a Bithynian age for the lower Wellenkalk and a Pelsonian age for the middle and upper Wellenkalk (Kozur 1999b). An early Illyrian age for miospore assemblages from the Middle Muschelkalk (Visscher et al. 1993) is indicated by ammonoid and conodont faunas from the Schaumkalk Member (Kozur 1999b). The base of the Ladinian at the Bagolino section in the Southern Alps (Northern Italy) has been ratified by ICS and IUGS in 2005 (Brack et al. 2005). Unfortunately, no palynological data have been published except for an abstract (Hochuli & Roghi 2002). Direct correlation of the Anisian/Ladinian boundary with the Germanic Triassic by ammonoids and conodonts is not possible because of environmental restrictions (Kozur 1999a). However, the LOD of Stellapollenites thiergartii and the FAD of Ovalipollis pseudoalatus in both realms indicate that this boundary may occur in the middle Upper Muschelkalk (mo2), halfway

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between the Spiriferina and Cycloides Beds of the Meibner Formation (Kozur 1999b). The FODs of Cannanoropollis spp. and Kuglerina meieri characterize the Illyrian– Fassanian boundary in the Alpine realm (Brugman 1986; Go´cza´n & Oravecz-Scheffer 1993, 1996). Placement of the base of the Longobardian just above the Cycloides Bed (Kozur 1999b) is accepted here. The base of the Carnian at Prati di Stuores/ Stuores Wiesen section (Province of Belluno, Veneto Region, NE Italy) has been ratified in 2008 (Mietto et al. 2008). The palynology of this section has been described in detail by Cirilli & Roghi (1999) and Roghi (2004) and is summarized in Mietto et al. (2008) as follows: Typical upper Ladinian–lower Carnian sporomorphs such as Uvaesporites gadensis, Kuglerina meieri, Ovalipollis pseudoalatus, Todisporites spp., Aratrisporites spp., Reticulatisporites dolomiticus, Sellaspora rugoverrucata and the circumpolles form Partitisporites novimundanus are present throughout the section. In the upper part of the regoledanus Subzone, Concentricisporites cf. C. bianulatus, Enzonalasporites vigens, Kyrtomisporis ervii, Gordonispora fossulata and Duplicisporites granulatus have their first occurrence. ‘Lueckisporites’ cf. singhii first appears very close to the top of the regoledanus Subzone. Moreover, Nevesisporites vallatus, Todisporites marginales, Calamospora sp. A, Apiculatisporites parvispinosus and Densosporites cf. variomarginatus are restricted to the regoledanus Subzone. Nevesisporites vallatus has its last occurrence at the proposed boundary. Above the proposed boundary, in the lower part of the canadensis Subzone, Vallasporites ignacii, Patinasporites densus and Aulisporites cf. A. astigmosus, together with Duplicisporites verrucosus and Camerosporites secatus, have their first occurrence. In the samples from the uppermost part of the canadensis Subzone, the first occurrences of Weylandites magmus, Camerosporites pseudoverrucatus and Samaropollenites speciosus are found. No other significant bioevents are documented in the overlying aon Subzone. The palynological content of the Prati di Stuores/Stuores Wiesen section supports the proposal to place the boundary at the FAD of Daxatina canadensis. This ammonoid event occurs very closely to significant changes of the palynological association, consisting in the replacement of upper Ladinian sporomorphs by typical Carnian microflora. The first appearance of Patinasporites densus marks the base of the vigens–densus phase and is associated with the first appearance of Vallasporites ignacii (Van der Eem 1983); their common occurrence has been traditionally considered to be close to the base of the Carnian (Fisher 1972; Visscher & Brugman 1981; Van der Eem 1983; Fisher & Dunay 1984; Warrington

1996b; Hochuli & Frank 2000; Roghi 2004), although the age of this phase can be confirmed with new findings. Moreover, on the basis of the first occurrence of Concentricisporites cf. C. bianulatus within the regoledanus Subzone, the base of the Concentricisporites cf. C. bianulatus assemblage (Roghi 2004) is late Ladinian (Longobardian) in age. In agreement with the first option of Kozur (1999b), the Ladinian-Carnian boundary lies close to the Acrodus Bank at the base of the lower ‘Estheria Beds’ (upper Grabfeld Formation); the lower part of that formation, formerly the Lower Gipskeuper, is late Longobardian in age. Consequently, the Lettenkeuper is of Longobardian age. It is widely accepted that the base of the Schilfsandstein is of late early Carnian, Julian age and may in some parts of Europe reach into the early Tuvalian. The LADs of Palaeospongisporis europaeus, Striatoabieites balmei, Triadispora crassa and T. plicata support this age assignment. A hiatus between the upper Grabfeld Formation and the Schilfsandstein (Bachmann & Kozur 2004) encompasses the early Julian. The palynostratigraphy of the Norian is largly unknown in NW Europe. In the Germanic Basin, palynological records are rare and an exact age assignment remains uncertain because the Steinmergelkeuper, now Arnstadt and Lo¨wenstein formations, is bounded by hiatuses (Bachmann & Kozur 2004; Nitsch et al. 2005). In the Alpine realm, Norian palynological records are virtually unknown because of the paucity of sediments that yield palynomorphs. Nevertheless, according to Bachmann & Kozur (2004), the Norian- Rhaetian boundary is provisionally correlated with the base (Postera Beds) of the Exter Fm. (Ra¨tkeuper). In the Alpine realm the GSSP for the base of the Rhaetian has been proposed at the Steinbergkogel near Hallstatt (Austria) (Krystyn et al. 2007). Unfortunately, palynological samples from this section are barren because of the Hallstatt-type facies. As discussed above, the FAD of Rhaetipollis germanicus is likely a usefull event to characterize the base of the Rhaetian palynologically, at least in Central and NW Europe. Thoughtful comments by Henk Visscher, Alfred Traverse, Geoffrey Warrington and James Ogg improved this work and are gratefully acknowledged. WMK acknowledge funding from the high potential program of Utrecht University. This is NSG publication no. 20100204 of the Netherlands Research School of Sedimentary Geology.

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S IMMS , M. J. & R UFFELL , A. H. 1989. Synchroneity of climatic change and extinction in the Late Triassic. Geology, 17, 265 –268. S OLLAS , I. B. J. 1901. Fossil in the Oxford University Museum, V. On the structure and affinities of the Rhaetic plant Naiadita. Quarterly Journal of the Geological Society, 57, 307–312. S TONELEY , H. M. M. 1958. The Upper Permian flora of England. Bulletin of the British Museum Geology, 3, 293–337. S UNEBY , L. B. & H ILLS , L. V. 1988. Palynological zonation of the Heiberg Formation (Triassic– Jurassic) eastern Sverdrup Basin, Arctic Canada. Bulletin of Canadian Petrolium Geologist, 36, 347– 361. S ZURLIES , M. 2004. Magnetostratigraphy: the key to a global correlation of the classic Germanic Trias – case study Volpriehausen Formation (Middle Buntsandstein), Central Germany. Earth and Planetary Science Letters, 227, 395– 410. S ZURLIES , M. 2007. Latest Permian to Middle Triassic cyclo-magnetostratigraphy from the Central European Basin, Germany: Implications for the geomagnetic polarity timescale. Earth and Planetary Science Letters, 261, 602– 619. S ZURLIES , M., B ACHMANN , G. H., M ENNING , M., N OWACZYK , N. R. & K A¨ DING , K.-C. 2003. Magnetostratigraphy and high-resolution lithostratigraphy of the Permian– Triassic boundary interval in Central Germany. Earth and Planetary Science Letters, 212, 263–278. T HIERGART , F. 1949. Der stratigraphische Wert mesozoischer Pollen und Sporen. Palaeontographica B, 89, 1 –34. T RAVERSE , A. 2004. Proposal to conserve the fossil pollen morphogeneric name Classopollis against Corollina and Circulina. Taxon, 53, 847– 848. T RAVERSE , A. 2008. Paleopalynology. 2nd Series. Topics in Geobiology, 28. [2nd ed. 2007. Corr. 2nd printing, 2008, XVIII.] T UCKER , M. E. & B ENTON , M. J. 1982. Triassic environments, climates and reptile evolution. Palaeogeography, Palaeoclimatology, Palaeoecology, 40, 361–379. VAN DE S CHOOTBRUGGE , B., Q UAN , T. M. ET AL . 2009. Floral changes across the Triassic/Jurassic boundary linked to flood basalt volcanism. Nature Geoscience, 2, 589–594. V AN DER E EM , J. G. L. A. 1983. Aspects of Middle and Late Triassic palynology. 6. Palynological investigations in the Ladinian and Lower Karnian of the Western Dolomites, Italy. Review of Palaeobotany and Palynology, 39, 189–300. V AN DER Z WAN , C. J. & S PAAK , P. 1992. Lower to Middle Triassic sequence stratigraphy and climatology of the Netherlands, a model. Palaeogeography, Palaeoclimatology, Palaeoecology, 91, 277–290. V ENKATACHALA , B. S. & G O´ CZA´ N , F. 1964. The sporepollen flora of the Hungarian ‘Ko¨ssen Facies’. Acta Geologica, 8(1 –4), 203– 228. V ENKATACHALA , B. S., T IWARI , R. S. & V IJAYA , 1995. Diversification of spore – pollen ‘character states’ in the Indian Permian. Review of Palaeobotany and Palynology, 85, 319– 340.

TRIASSIC PALYNOLOGY OF CENTRAL AND NORTHWESTERN EUROPE V ISSCHER , H. 1966. Palaeobotany of the Mesophytic III. Plant microfossils from the Upper Bunter of Hengelo, The Netherlands. Acta Botanica Neerlandica, 15, 316–375. V ISSCHER , H. 1971. The Permian and Triassic of the Kingscourt outlier, Ireland. Geological Survey of Ireland, Special Paper, 1, 1–114. V ISSCHER , H. 1974. The impact of palynology on Permian and Triassic stratigraphy in Western Europe. Review of Palaeobotany and Palynology, 17(1– 2), 5– 19. V ISSCHER , H. & B RUGMAN , W. A. 1981. Ranges of selected palynomorphs in the Alpine Triassic of Europe. Review of Palaeobotany and Palynology, 34, 115–128. V ISSCHER , H. & B RUGMAN , W. A. 1988. The PermianTriassic boundary in the southern Alps: A palynological approach. Memories Societa Geologica Italiana, 34, 121–128. V ISSCHER , H. & K RYSTYN , L. 1978. Aspects of Late Triassic palynology. 4. A palynological assemblage from ammonoid-controlled Late Karnian (Tuvalian) sediments of Sicily. Review of Palaeobotany and Palynology, 26, 93–112. V ISSCHER , H. & V AN DER Z WAN , C. 1981. Palynology of the Circum-Mediterranean Triassic: phytogeographical and palaeoclimatological implications. Geologische Rundschau, 70, 625– 636. V ISSCHER , H., B RUGMAN , W. A. & V AN H OUTE , M. 1993. Chronostratigraphical and sequence stratigraphical interpretation of the palynomorph record from the Muschelkalk of the Obernsees well, South Germany. In: H AGDORN , H. & S EILACHER , A. (eds) Proceedings Muschelkalk Scho¨ntaler Symposium, Goldschneck-Verlag, Stuttgart, 145–152. V ISSCHER , H., V AN H OUTE , M., B RUGMAN , W. A. & P OORT , R. J. 1994. Rejection of a Carnian (Late Triassic) ‘pluvial event’ in Europe. Review of Palaeobotany and Palynology, 83, 217– 226. ¨ ber einige merkwu¨rdige VON B UCH , C. L. 1845. U Muschel-Reste des oberen Italiens. Monatsberichte der Ko¨niglichen Preussischen Akademie der Wissenschaften Berlin, 1845, 25–28. ¨ RSCHNER , VON H ILLEBRANDT , A., K RYSTYN , L. & K U W. M. 2007. A candidate GSSP for the base of the

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Jurassic in the Northern Calcareous Alps (Kuhjoch section; Karwendel Mountains, Tyrol, Austria). International Subcommission on Jurassic Stratigraphy Newsletter, 34(1), 2– 20. W ARRINGTON , G. 1967. The correlation of the Keuper by miospores. Nature, 214, 1323–1324. W ARRINGTON , G. 1970. The stratigraphy and palaeontology of the ‘Keuper’ Series of the Central Midlands of England. Quarterly Journal of the Geological Society, 126, 183– 223. W ARRINGTON , G. & S CRIVENER , R. C. 1990. The Permian of Devon, England. Review of Palaeobotany and Palynology, 66, 263–272. W ARRINGTON , G. 1996a. Permian spores and pollen. In: J ANSONIUS , J. & M C G REGOR , D. C. (eds) Palynology: principles and applications. American Association of Stratigraphic Palynologists Foundation, 2, 607– 619. W ARRINGTON , G. 1996b. Triassic spores and pollen. In: J ANSONIUS , J. & M C G REGOR , D. C. (eds) Palynology: principles and applications. American Association of Stratigraphic Palynologists Foundation, 2, 755– 766. W ARRINGTON , G. 2008. Palynology of the Permian succession in the Hilton Borehole, Vale of Eden, Cumbria, UK. Proceedings of the Yorkshire Geological Society, 57, 123– 130. W ARRINGTON , G., I VIMEY -C OOK , H. C., E DWARDS , R. A. & W HITTAKER , A. 1995. The Late Triassic –Early Jurassic succession at Selworthy, West Somerset, England. Proceedings of the Ussher Society, 8, 426– 432. W ENG , C., H OOGHIEMSTRA , H. & D UIVENVOORDEN , J. F. 2007. Response of pollen diversity to the climatedriven altitudinal shift of vegetation in the Colombian Andes. Philosophical Transactions of the Royal Society B, 362, 253– 262. W ILLS , L. J. 1910. On the fossiliferous Lower Keuper rocks of Worcestershire, with descriptions of some plants and animals discovered therein. Proceedings of the Geologists, Association, 21, 249–331. Y AROSHENKO , O. P. 1997. Palynology and phytogeography of the Early Triassic. Paleontological Journal, 31, 168– 177.

Upper Triassic– lowermost Jurassic palynology and palynostratigraphy: a review SIMONETTA CIRILLI Department of Earth Sciences, University of Perugia, Piazza Universita`, 06123 Perugia, Italy (e-mail: [email protected]) Abstract: This review advances understanding of the palynostratigraphy of the Late Triassic –Early Jurassic by correlating the established palynozonations for the northern and southern hemispheres. Previous palynological studies have contributed greatly to our understanding of the biostratigraphy, paleoclimatology and paleogeography of the Upper Triassic. In general, palynology is a good tool for interregional cross-correlation of marine and non-marine successions because palynomorphs, unlike most of other fossils, commonly are present in continental and marine environments. Currently, however, biostratigraphical resolution based on Upper Triassic palynomorph assemblages is rather low, primarily because of the rarity of successions that are independently dated (i.e. via ammonoids, conodonts, isotopes, paleomagnetism) to correlate the palynomorph assemblages, but also for other reasons, such as microfloristic provincialism, palaeoenvironmental conditions and differential preservation of palynomorph assemblages. During the last few decades many palynological studies have attempted to integrate and improve the biostratigraphical correlations and paleoclimatologic reconstructions across the Triassic –Jurassic boundary. Several authors have recognized specific microfloral assemblages with well-defined and recognizable suites of palynomorphs that enhance the importance of palynomorphs in the definition of Triassic– Jurassic stages. Comparison of the palynomorph assemblages from different biostratigraphical stages demonstrates that a change occurred in the palynofloral composition of the Tethyan domain between the Carnian and the earliest Hettangian that was gradual and without abrupt changes.

Upper Triassic palaeogeography and palaeoclimatology The Late Triassic, a geological time interval of about 30 million years (Gradstein et al. 2004), was a period of rapidly changing palaeogeography. During the Triassic, Pangaea, which extended from about 858N to 908S, progressively evolved from maximum continental assembly to the inititiation of breakup (Ziegler et al. 1983; Hesselbo 2000; Golonka 2004; Buratti & Cirilli 2007). Beginning in the Middle Triassic, Pangaea started to divide into the northern landmass Laurasia (including present day North America, Europe and Asia) and the southern landmass Gondwana (South America, Africa, Antarctica, India and Australia), separated by the narrow and shallow Tethys seaway. In the southern hemisphere, the Cimmerian blocks (present day Turkey, Iran and Tibet) separated from Australia and moved northward, originating the back-arc basin of Neotethys, and after the subduction of Paleotethys, collided against Eurasia at the end of the Triassic (Corsin & Stampfli 1977; Sengo¨r 1984; Marcoux et al. 1993; Muttoni et al. 1996, 2001; Besse et al. 1998; Gaetani et al. 1998, 2005; Dercourt et al. 2000; Moix et al. 2008). The rift between Africa, South America and North America was probably already active by the end

of the Triassic, as documented by the Karoo synrift deposits in southeastern Africa (Zerfass et al. 2004; Geiger et al. 2004). This rifting stage took place contemporaneously with the emplacement of the voluminous tholeiitic dikes, sills and flood basalts of the Central Atlantic Magmatic Province (CAMP) over a total surface area of about 10 million km2 in North and South America, Africa and Europe (Marzoli et al. 1999, 2004; Olsen et al. 2003; Verati et al. 2007). The global Late Triassic climate and thus the floral distribution were strongly controlled by the symmetrical distribution of the Pangaean landmasses around the equator and by the presumed absence of polar ice at high latitude (Frakes & Francis 1988; Parrish 1993; Buratti & Cirilli 2007). The latitudinal temperature gradient was appreciably lower than today, resulting in a more equable average global temperature with warm temperate belts expanded to higher latitudes. Several authors have posited a strongly seasonal climate for both hemispheres influenced by the monsoon circulation and enhanced by the shape of Pangaea (e.g. Hallam 1985; Kutzbach & Gallimore 1989; Simms & Ruffell 1990; Dubiel et al. 1991; Parrish 1993). Kent & Olsen (2000) hypothesized a zonal climate pattern for the Late Triassic, with a narrow equatorial humid zone and an arid belt centred

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 285– 314. DOI: 10.1144/SP334.12 0305-8719/10/$15.00 # The Geological Society of London 2010.

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around 308, passing to more temperate climates at higher latitude. Rapid negative excursions in d13C in both carbonate and sedimentary organic matter have been observed in marine Tr– J boundary sections close to the end-Triassic mass extinction, for example, from Hungary, England, USA, Canada, Spain, Italy, Austria (e.g. Pa`lfy et al. 2001; Hesselbo et al. 2002; Guex et al. 2004; Ward et al. 2004; Go´mez et al. 2007; Galli et al. 2007; Ku¨rschner et al. 2007; Williford et al. 2007; Van de Schootbrugge et al. 2008; Go¨tz et al. 2009; Ruhl et al. 2009). Additionally, an apparently synchronous decrease of the stomatal index of fossil leaves (McElwain et al. 1999) has suggested a disruption of the global carbon cycle, potentially involving some combination of global warming, productivity decline and methane hydrate release (Pa`lfy et al. 2001; Tanner et al. 2004; Ward et al. 2004; Lucas & Tanner 2008; Van de Schootbrugge et al. 2008; Ruckwied & Go¨tz 2009). It has been suggested by some authors that there was a global greenhouse warming, estimated at about 38 to 4 8C, and acidic atmospheric pollution resulted from a significant atmospheric loading by CO2 and SO2 related to eruptions of the Central Atlantic Magmatic Province (CAMP) (Marzoli et al. 1999, 2004; McElwain et al. 1999; Hesselbo et al. 2002; Guex et al. 2004; Tanner et al. 2004, 2007; Schaltegger et al. 2008; Van de Schootbrugge et al. 2008, 2009), although it is yet unresolved if the eruptions started before (Marzoli et al. 2004, 2008) or after the Tr –J boundary (Whiteside et al. 2007, 2008). Recent palynological studies (Marzoli et al. 2004; Cirilli et al. 2009), in concert with other data (Jourdan et al. 2009) document that CAMP volcanism started before the Tr–J boundary and therefore support the hypothesis that the CAMP eruptions had a causative role in the biotic crisis across the Tr–J boundary. Around the Tr–J boundary evidence of increasing humidity and seasonality have been recorded as testified by the large amount of enriched organic matter and shaly sediments deposited within low dysoxyc-anoxic basins of the western Tethys (e.g. Cirilli et al. 1999; Berra & Cirilli 1997; Bonis et al. 2010). As recorded in the Northern Calcareous Alps, the beginning of black shale deposition coincides with the onset of the initial negative d13C excursion (Ku¨rschner et al. 2007; Ruhl et al. 2009) just before the Tr–J boundary. The climate change and related anoxic conditions have been well documented on the basis of high resolution palynological and geochemical data (Bonis et al. 2010). The location and shifting positions of the paleoclimatic belts strongly controlled paleofloral composition and its geographical distribution. Dolby & Balme (1976) distinguished two different

microfloras during the Carnian, the Onslow microflora, a mixture of Gondwanan and European taxa, spread throughout western Europe and northwestern Australia (308 –358S) that is characteristic of continental margins and warm temperate rain-forests, and the Ipswich microflora developed in southern and eastern Australia (extending from 258S in Queensland to more than 758S in Antarctica), which is referable to cool temperate plant communities (Fig. 1). The European taxa of the Onslow microflora mainly include the genera Aulisporites, Camerosporites, Duplexisporites, Decussatisporites (¼ Weylandites), Enzonalasporites, Infernopollenites, Minutosaccus, Ovalipollis and Samaropollenites, all of which are absent in the Ipswich Microflora (Dolby & Balme 1976; Buratti & Cirilli 2007). In the last three decades, numerous additional phytogeographic data, such as the recovery in the northern hemisphere of the index species Samaropollenites speciosus, and other typical southern elements of Onslow microfloral affinity, revealed a broad pattern of Triassic palynofloral provincialism (e.g. Dolby & Balme 1976; de Jersey & McKeller 1981; Fisher & Dunay 1984; Helby et al. 1987; de Jersey & Raine 1990; Cirilli & Eshet 1991; Foster et al. 1994; Buratti & Cirilli 2007). The common occurrences of the circumMediterranean sporomorphs in the Onslow Microflora assemblages support the hypothesis that the distribution of the Onslow Microflora includes the Carnian miospore assemblages extending from the western edge of Tethys to its southeastern shorelines (e.g. Timor) (Buratti & Cirilli 2007). The strong palynofloral affinity between the Circum-Mediterranean and Onslow microfloras confirm the existence of a homogeneous parent plant community. This cosmopolitan community most likely grew under the influence of an equable climatic regime characterised by favourable temperatures and humid conditions controlled by monsoonal circulation. A decrease in macrofloral diversity is associated with a less pronounced microfloristic provincialism, which in the latest Triassic-earliest Jurassic coincided with the migration events of some plant communities (e.g. Cheirolediaceae) from the Tethyan region to higher latitudes. The diffusion of Cheirolepidiaceae (Circumpolles producers) probably occurred across the coastal migration pathways created during the Carnian –Norian plate reorganization. The recovery of Circumpolles taxa from the northern Australian margin (West Timor) (Martini et al. 2004; Buratti & Cirilli 2007) demonstrates that Cheirolepidiaceae reached the southern hemisphere during the Norian. Their dispersal is thus a very important floral event that started in the northern and southern hemispheres during the Norian, causing the gradual decline of Upper Triassic microfloral provincialism,

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Fig. 1. Distribution of Ipswich and Onslow microfloras (after Buratti & Cirilli 2007).

and, later, the global diffusion of a more homogeneous Lower Jurassic flora. The dispersion of the Classopollis group (Circulina –Classopollis– Corollina complex) may be considered as a postCarnian global microfloristic event, although it appeared at different times in the northern and southern hemispheres.

Palynostratigraphy: limits and applicability Palynology has great potential as a tool for interregional cross-correlation of marine and non-marine successions because spores and pollens are commonly present in both continental and marine environments. At present, however, the biostratigraphical resolution based on Upper Triassic palynomorph assemblages is rather low in comparison to their potential wide distribution. This low resolution partially stems from the rarity of successions, particularly in the Norian, dated independently via biostratigraphy (i.e. ammonoids or conodonts) or other geochronologic tools (e.g. radio-isotopic dating or paleomagnetic correlation). Currently, few marine successions in the world provide an ammonoid-integrated palynostratigraphy. Consequently, palynomorph assemblages suffer from a

lack of well-calibrated ranges of the most important taxa, which sometimes are diachronous on a regional scale (Batten & Koppelhus 2002). However, the real limitation is represented by the classical approach based only on the concept of LO (Last Occurrence) and FO (First Occurrence) instead of on quantitative methods (e.g. Brugmann et al. 1994; Bonis et al. 2009, among others). Other important factors that may limit the extension of palynological correlations include microfloral provincialism, diversity of palaeoenvironmental conditions and differing grades of preservation of palynomorph assemblages. Microfloral provincialism is directly linked to phytogeographic distribution, which is in turn strongly controlled by climate. Additionally, the striking differences in composition of coeval palynomorph assemblages could derive from diagenetic overprinting, which may destroy some sporomorphs, but preserve others. Soon after deposition, during the first early diagenetic stages as well as later during the deep burial diagenesis, the organic facies can undergo pronounced alterations that may drastically reduce and select the amounts and types of sedimented microflora, thereby modifing the original palynomorph assemblages (Tyson 1995; Batten 2002). Consequently, the use of palynostratigraphic schemes in global scale cross-correlations should

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estimate the possible influence of the above-cited limiting factors and take into consideration the compositional variation of the palynomorph assemblages (cf. concept of phase in Schuurman 1977) rather than the occurrence of single forms.

Carnian palynomorph assemblages of the northern hemisphere The Carnian palynological record is distinguished by the blooming of Circumpolloid genera such as Camerosporites secatus, Duplicisporites

Ladin. late

granulatus, Paracirculina scurrilis, Praecirculina granifer, and the monosaccate pollen Patinasporites densus (Fig. 2). Bisaccate pollen grains are scarce in comparison to the older Ladinian palynomorph assemblages, although some taxa continue to be significant for palaeofloristic reconstructions (i.e. Samaropollenites speciosus, and the protosaccate Ovalipollis pseudoalatus (Dolby & Balme 1976; Visscher & Krystyn 1978; Visscher & Van der Zwan 1981; Cirilli & Eshet 1991; Brugman et al. 1994; Hochuli & Frank 2000; Warrington 2002; Buratti & Carrillat 2002; Buratti & Cirilli 2007; Traverse 2008).

Carnian early late

Norian early late

Callialasporites dampieri Camerosporites secatus Cerebropollenites macroverrucosus Cerebropollenites thiergartii Classopollis meyerianus Classopollis murphyae Classopollis torosus Duplicisporites granulatus Enzonalasporites vigens Granuloperculatipollis rudis Heliosporites reissingeri "Lueckisporites" cf. L. singhii Lunatisporites rhaeticus Paracirculina quadruplicis Partitisporites novimundanus Patinasporites densus Pinuspollenites minimus Porcellispora longdonensis Pseudoenzonalasporites summus Retitriletes semimuris Rhaetipollis germanicus Ricciisporites tuberculatus Samaropollenites speciosus Trachysporites fuscus Tsugaepollenites pseudomassulae Vallasporites ignacii Fig. 2. Stratigraphic ranges of selected sporomorphs of the northern hemisphere.

Rhaetian

Hett. early

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In the European domain, the Carnian stage, including the Ladinian– Carnian boundary, was palynologically defined on the base of the occurrence of Ovalipollis pseudoalatus in assemblages with numerous distinctive species such as Camerosporites secatus, Duplicisporites granulatus, Ellipsovelatisporites plicatus, Enzonalasporites vigens, Infernopollenites spp. and Triadispora spp. (phase 1 of Schuurman, 1977, 1979) (Figs 2–4). In Central Europe (Poland), Orłowka-Zvolin´ska (1983; 1985) distinguished five ammonoidcontrolled palynomorph assemblage Zones covering the Upper Triassic. The oldest Heliosaccus dimorphus assemblage can be correlated with the late Ladinian phases of western Europe, succeeded stratigraphically by the Carnian Ovalipollis–Triadispora and Toroisporis– Camarozonosporites – Aulisporites assemblages and by the uppermost Carnian –Norian (which includes also the Rhaetian) Classopollis assemblage, marked by the presence of Riccisporites at the top. Roghi (2004) provided a palynological characterization of the Julian – Tuvalian boundary in an ammonoid-controlled section of the Julian Alps (Italy) based on the FAD of significant sporomorphs (e.g. Duplicisporites continuus, Pseudoenzonalasporites summus in the Austrotrachyceras austriacum Zone, Ricciisporites tuberculatus in the Tropites dilleri Zone and Granuloperculatipollis rudis in the Tropites subbullatus Zone), within the Duplicisporites continuus assemblage spanning the Julian –Tuvalian interval (Figs 2–4). The Carnian successions from the Canadian Arctic (Fisher & Bujak 1975; Fisher 1979) and the western Barents Sea (Hochuli et al. 1989; Mørk et al. 1992) are comparable with those from continental Europe, except for some differences (Warrington 2002). The Circumpolles (e.g. Classopollis group) are rarer in the Barents Sea than in Europe, and some taxa such as Camerosporites and Duplicisporites seem to have their LAD in the lower Carnian. Both differences could reflect a colder climate at higher latitudes or, alternatively, poorer independent stratigraphic control (Warrington 2002). In North America, palynologically-dated nonmarine Upper Triassic successions crop out in the southwestern and eastern USA, for example, the Chinle Formation in Arizona and New Mexico (Dunay & Fisher 1979; Litwin et al. 1991; Cornet 1993), the Dockum Group in Texas (Fisher & Dunay 1984; Cornet 1993) and the Chatham Group of North Carolina (Cornet & Olsen 1985; Litwin & Ash 1993). Litwin et al. (1991) proposed three miospore assemblage Zones for the Chinle Formation in the southwestern USA (Fig. 5) correlated with analogous assemblages of the Newark Supergroup in the eastern USA (Dunay & Fisher 1974; Cornet & Traverse 1975; Manspeizer &

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Cousminer 1988; Robbins et al. 1988). Zone I, marked by abundant Brodispora striata, Equisetosporites chinleanus, Lagenella martini, Minutosaccus crenulatus, Samaropollenites speciosus and the LAD of Lunatisporites aff. L. noviaulensis, was assigned to the late Carnian. Zone II contains the FADs of Camarozonosporites rudis, Enzonalasporites vigens, Heliosaccus dimorphus, Ovalipollis ovalis, Pseudoenzonalasporites summus and other taxa (e.g. Alisporites spp., Cycadopites stonei, Guthoerlisporites cancellosus), and the LADs of Brodispora striata, Camerosporites secatus, Equisetosporites chinleanus and Lagenella martini. This assemblage, which has close similarities with European ones, was dated as Tuvalian. Zone III, founded on the FADs of Camerosporites verrucosus and Kyrtomisporis spp., was dated as early Norian because of the absence of significant Carnian taxa (e.g. Brodispora striata and Camerosporites secatus) and by the presence of Pseudoenzonalasporites summus.

The Camerosporites secatus phase Starting from the definition of Schuurman’s phase I, Visscher & Krystyn (1978) introduced the ‘Camerosporites secatus phase’, characterized by a rapid diversification of the Circumpolloid genera (e.g. Camerosporites secatus, Duplicisporites granulatus and Praecirculina granifer), associated with a group of monosaccate spores (Enzonalasporites vigens, Patinasporites densus, Pseudoenzonalasporites summus and Vallasporites ignacii) and with the bisaccate Samaropollenites speciosus. This phase was introduced to provide a practical palynological tool for world-wide correlation because its characteristic assemblage was recorded in many palaeofloristic provinces: in the Alpine Triassic of Europe (Visscher & Brugman 1981), Italy (Van der Eem 1983; Broglio Loriga et al. 1999; Roghi 2004; Mietto et al. 2007), southern Albania (Cirilli & Montanari 1994), Israel (Cirilli & Eshet 1991), Africa (Bourmouche et al. 1996), Tunisia (Mehdi et al. 2009), USA (Fisher & Dunay 1984; Litwin & Ash 1993), Arctic Canada (Fisher & Bujak 1975), Australia (Dolby & Balme 1976) and Timor (Martini et al. 2000). Some of the above records come from independently dated (ammonoid and/or conodonts) sections, for example, in Sicily, within the Tropites subbullatus Zone or the ‘Anatropites-Bereich’, dated as Tuvalian (Visscher & Krystyn 1978); in the Western Dolomites (Italy) within the Daxatina canadensis Subzone (Van der Eem 1983; Broglio Loriga et al. 1999; Mietto et al. 2007) (Figs 3–4); in Austria from the Trachyceras aon Zone (Cordevolian), the Halobia rugosa/Carnites floridus Zone (Julian) and the Tropites subbullatus Zone (Tuvalian) (Dunay & Fisher 1978).

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Fig. 3. Schematic comparison of the main palynozonations across the Late Triassic and Early Jurassic of the northern hemisphere. Data from: (1) Ku¨rschner et al. 2007; (2) Krystyn et al. 2007a, b; (3) Broglio Loriga et al. 1999; Mietto et al. 2007; (4) Roghi 2004; (5) Van der Eem 1983; (6) Schuurman 1977, 1979; (7) Morbey 1975, 1978; (8) Lund 1977; (9) Orbell 1973. FO: first occurrence; LO: last occurrence. Z (Zone); Sz (Subzone); Palynozones from

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291

Fig. 4. Main palynozones and palynological assemblages across the Ladinian–Carnian boundary. Z, Zone; Sz, Subzone; FO, first occurrence; LO, last occurrence.

The Camerosporites secatus phase was originally considered as an exclusively Carnian palynological event and later extended into the Ladinian (Visscher & Brugman 1981; Van der Eem 1983; Besems 1981, 1983; Brugman et al. 1994). However, the long range distribution of this phase (upper Ladinian to Carnian) made it unsuitable for a detailed biostratigraphical subdivision as many authors introduced further subdivisions. As already suggested by Visscher & Krystyn (1978), an upper Ladinian trend may be discerned from the occurrence of Camerosporites secatus with Echinitosporites iliacoides, Retisulcites perforates, Lunatisporites spp., and, at least in the Alpine –Mediterranean part of Europe, with the ‘northern’ element Staurosaccites quadrifidus, while an exclusively Carnian age may be defined on the gradual diversification of circumsulcate form genera and monosaccates belonging to the Enzonalasporites–Patinasporites –Pseudoenzonalasporites–Vallasporites group (Warrington 2002). The FAD of Patinasporites densus marks the base

of the early Carnian (cf. the vigens –densus phase of Van der Eem 1983) (Figs 2–3), which is commonly associated with the first appearance of Vallasporites ignacii at the base of the Carnian (Visscher & Brugman 1981; Van der Eem 1983; Fisher & Dunay 1984; Blendinger 1988; Hochuli et al 1989; Broglio Loriga et al. 1999; Hochuli & Frank 2000; Warrington 2002; Roghi 2004; Mietto et al. 2007), while the appearance of Paracirculina quadruplicis and Samaropollenites speciosus differentiates the youngest Carnian (Tuvalian) Camerosporites secatus phase (Visscher & Krystyn 1978; Visscher et al. 1980; Cirilli & Eshet 1991; Cirilli & Montanari 1994; Warrington 2002).

The palynological Ladinian –Carnian boundary In Italy (Prati di Stuores section, western Dolomites), Van Der Eem (1983) established three

Fig. 3. (Continued) Ku¨rschner et al. 2007: RL, Rhaetipollis– Limbosporites; RPo, Rhaetipollis –Porcellispora; TPo, Trachysporites–Porcellispora. TH, Trachysporites– Heliosporites; TPi, Trachysporites–Pinuspollenites. Palynozones from Morbey 1975: TR, Classopollis torosus–Granuloperculatipollis rudis; TL, Riccisporites tuberculatus– Hystrichosphaeridium langi; LR, Limbosporites lundbladii– Cingulizonates rhaeticus; TK, Perinosporites thuringiacus– Comparodinium koessenium; RG, Rhaetogonyaulax rhaetica– Rhaetipollis germanicus; Rk, Rhombodella kendelbachia; LL: Carnisporites lecythus–Zebrasporites laevigatus; Me, Carnisporites megaspiniger; Mi, Convolutispora microrugulata; FG, Densosporites fissus–Lycopodiumsporites gracilis. Details in Figs 4, 6, 7.

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ammonoid-controlled palynomorph assemblages across the Ladinian–Carnian transition (Figs 3– 4), which in ascending order are: the secatus–dimorphus and secatus-vigens phases (late Ladinian, Longobardian), the vigens–densus phase (early Carnian) and the densus-maljawkinae phase (Julian). The same section (Prati di Stuores) is a candidate GSSP section for the Ladinian–Carnian boundary, herein based on the FAD of the ammonite Daxatina canadensis (Broglio Loriga 1999; Mietto et al. 2007). The miospore assemblages for this section, correlated with ammonites and other fossil groups, show a compositional variation across the boundary distinguished by the FAD of Concentricisporites cf. C. bianulatus, Duplicisporites granulatus, Enzonalasporites vigens, Gordonispora fossulata, Kyrtomisporis ervii, and ‘Lueckisporites’ cf. singhii in the upper part of the regoledanus Subzone, and by the LAD of Nevesisporites vallatus at the putative boundary, with additional elements having a wider Ladinian-early Carnian range (e.g. Reticulatisporites dolomiticus, Partitisporites novimundanus). Aulisporites cf. A. astigmosus, Camerosporites secatus, Duplicisporites verrucosus, Patinasporites densus and Vallasporites ignacii, have their first occurrence in the lower part of the canadensis Subzone, while Camerosporites pseudoverrucatus, Samaropollenites speciosus and Weylandites magmus first occur in the uppermost part of this Subzone. According to this new palynomorph zonation, the base of the vigens-densus phase in the Prati di Stuores/Stuores Wiesen section is lowered by at least 130 metres with respect to Van der Eem (1983).

The Norian –Rhaetian miospore assemblages in the northern hemisphere In the Boreal realm, the Norian and Rhaetian palynomorph assemblages are documented in several sections, some of which, especially those of Rhaetian age, are independently dated: the Svalbard Archipelago, Arctic Canada, North Sea and Barents Sea (Fisher & Bujak 1975; Lund 1977, 2003; Fisher 1979; Pedersen & Lund 1980; Guy-Ohlson 1981; Smith 1982; Hochuli et al. 1989; Dybkjær 1991; Batten & Koppelhus 2002; Koppelhus & Batten 2002; Lindstro¨m & Erlstro¨m 2006). They share common taxa such as Camarozonosporites spp., Chasmatosporites spp., Cingulizonates rhaeticus, Kyrtomisporis spp., Ovalipollis spp., Rhaetipollis germanicus, Triancoraesporites reticulatus, Zebrasporites spp., and Granuloperculatipollis rudis, Limbosporites lundbladii, Quadraeculina anellaeformis, Ricciisporites tuberculatus, the latter four taxa dominating the Rhaetian assemblages (Fig. 2). Lund (1977) established four

palynozones for the Danish Basin across the Late Triassic and Early Jurassic (Fig. 3), although there are some local variations in the relative abundances of taxa (Dybkjær 1988, 1991): (1) the Norian–early Rhaetian Classopollis –Enzonalasporites and the early Rhaetian Ricciisporites –Conbaculatisporites Zones, characterized by the presence of Ricciisporites tuberculatus and by the absence of older forms such as Enzonalasporites vigens; (2) the middle Rhaetian Rhaetipollis–Limbosporites Zone, marked by the occurrence of Limbosporites lundbladii and Rhaetipollis germanicus; (3) the late Rhaetian Ricciisporites –Polypodiisporites Zone, defined by the occurrence of Polypodiisporites polymicroforatus (¼ Convolutispora microrugulata) and Semiretisporis spp., rare Rhaetipollis germanicus and abundant trilete ornamented spores. This last Zone correlates with the Classopollis – Ricciisporites Zone of Dybkjær (1991), introduced as an alternative Zone because of the sporadic occurrence of Polypodiisporites polymicroforatus and Semiretisporis in the Danish Basin; and (4) the Hettangian Pinuspollenites–Trachysporites Zone is marked by the common presence of Pinuspollenites minimus, by trilete spores such as Conbaculatisporites spp., Trachysporites spp. and Uvaesporites reissingerii and by the sporadic presence of Cingulizonates rhaeticus and Limbosporites lundbladii. Koppelhus & Batten (2002) proposed a detailed palynozonation across the Late Triassic-Early Jurassic valid for northwestern Europe. This includes, in ascending order: (1) The Norian–early Rhaetian Classopollis –Enzonalasporites Zone subdivided into the Norian–early Rhaetian Classopollis – Porcellispora Subzone and the early Rhaetian Granuloperculatipollis rudis and Enzonalasporites– Conbaculatisporites Subzones; (2) the early to middle Rhaetian Ricciisporites –Conbaculatisporites Zone; (3) the middle Rhaetian Rhaetipollis– Limbosporites Zone; (4) the late Rhaetian Ricciisporites –Polypodiisporites (Lund 1977) and Classopollis–Ricciisporites Zones (Dybkjær 1991); and (5) the Hettangian Pinuspollenites –Trachysporites Zone, below the FAD of Cerebropollenites macroverrucosus, which is herein considered as Sinemurian –Pliensbachian (cf. Cerebropollenites macroverrucosus Zone of Dybkjær 1991). The middle to late Rhaetian assemblages from the Boreal realm are overall typified by the occurrence and occasional abundance of some significant taxa such as Cingulizonates rhaeticus, Classopollis spp., Kraeuselisporites (¼ Heliosporites) reissingeri, Limbosporites lundbladii, Lunatisporites rhaeticus, Ovalipollis spp., Rhaetipollis germanicus, Triancoraesporites reticulates and Semiretisporis gothae. Homogeneous Norian–Rhaetian miospore assemblages have been recorded in several regions

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of Europe, for example, England, Germanic and Alpine domains (e.g. Morbey & Dunay 1978; Schuurman 1979; Visscher et al. 1980; Visscher & Brugman 1981; Smith 1982; Brenner 1986; Batten & Koppelhus 2002; Warrington 2002; Hounslow et al. 2004), Circum-Mediterranean area (e.g. Brugman & Visscher 1988; Conway et al. 1990; Cirilli et al. 1993, 1994; Barro´n et al. 2006; Gomez et al. 2007) and Western Ciscaucasia (e.g. Yaroshenko 2007) and documented from independently dated sections in Europe (e.g. Northern Calcareous Alps, Austria: Morbey & Neves 1974; Morbey 1975; Mostler et al. 1978; Krystyn & Ku¨rschner 2005; Krystyn et al. 2007a, b; Ku¨rschner et al. 2007; Bonis et al. 2009). However, the Norian palynological assemblages still need to be independently dated because of the scarcity of Norian continental successions that might be correlated to coeval marine beds. At present, two candidate GSSP sections have been proposed for the base of the Norian: Black Bear Ridge (Williston Lake, B.C. Canada) (Orchard 2007) and Pizzo Mondello (Sicily, Italy) (Nicora et al. 2007) but both lack palynological assemblages to correlate with other proxies. The available palynozonations show that, at least in Europe, the Norian palynomorph assemblages reveal transitional features, characterised mainly by the progressive LADs of Carnian taxa and by the progressive appearance of species that become dominant in the Rhaetian assemblages. The most significant Carnian forms that still persist in the

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early Norian are Camerosporites spp., Duplicisporites spp., Ellipsovelatisporites rugosus, Enzonalasporites spp., Paracirculina quadruplicis spp., Patinasporites densus, Pseudoenzonalasporites summus and Vallasporites ignacii, together with some bisaccates such as Triadispora spp. and Infernopollenites parvus. Most of these taxa (e.g. Enzonalasporites vigens, Patinasporites densus, Vallasporites ignacii) seem to disappear in the upper part of the Norian or, locally, in the early Rhaetian (Morbey 1975; Schuurman 1977, 1979; Visscher et al. 1980; Fisher & Dunay 1981; Visscher & Brugman 1981) (Fig. 2). The gradual evolution from Norian to Rhaetian assemblages is underlined by the progressive dominance of genera belonging to the Circumpolles Group (Classopollis, Geopollis, Gliscopollis, Granuloperculatipollis) in association with Ovalipollis, Rhaetipollis and Ricciisporites. A suite of additional diversified forms that also may continue in the Rhaetian are Acanthotriletes varius, Camarozonosporites spp., Chasmatosporites spp., Densosporites fissus, Heliosporites reissingeri, Kyrtomisporis spp., Limbosporites lundbladii, Lycopodiacidites rugulatus, Perinosporites thuringiacus, Quadraeculina anellaeformis, Triancoraesporites spp., Uvaesporites argenteaeformis and Zebrasporites spp. (the TL, LR and TK Zones of Morbey 1975; phase III and IV of Schuurman 1977) (Figs 3, 6). The late Rhaetian assemblages are marked by the FAD of Retitriletes semimuris (phase IV of Schuurman 1977; Orbell 1973; Visscher et al. 1980; Visscher

Fig. 5. Miospore assemblage zones proposed for the Chinle Formation in the southwestern USA (modified by Litwin et al. 1991). FO, first occurrence; LO, last occurrence.

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& Brugman 1981) and by the LADs of a group of taxa, such as Granuloperculatipollis rudis, Lunatisporites rhaeticus, Ovalipollis pseudoalatus, Protohaploxypinus sp. cf. P. microcorpus, Tsugaepollenites pseudomassulae and of a group of trilete spores. Tsugaepollenites pseudomassulae is a common form in the European Rhaetian assemblages (Fig. 2), because it has been frequently recorded in England (Hounslow et al. 2004; Warrington et al. 2008), Austria (Morbey & Neves 1974; Hillebrandt et al. 2007; Ku¨rschner et al. 2007; Bonis et al. 2009a), southern France (Peybernes et al. 1988; Fre´chengues et al. 1993), Italy (Cirilli et al. 1994; Jadoul et al. 1994; Galli et al. 2007), Spain (Baudelot & Taugourdeau –Lantz 1986; Vachard et al. 1990; Calvet et al. 1993; Fre´chengues et al. 1993; Barro´n et al. 2006) and, more seldom, in Norian strata (e.g. Mallorca in Boutet et al. 1982). Recent palynological data from key Austrian sections, such as the Steinbergkogel section near Hallstatt, and the Kuhjoch section near Hinterriss in Tyrol, which have been proposed as candidate GSSPs for the Norian –Rhaetian boundary (Krystyn & Ku¨rschner 2005; Krystyn et al. 2007a, b) and the Rhaetian –Hettangian boundary (Ku¨rschner et al. 2007; Hillebrandt et al. 2007), respectively, fit well with the TL and LJ concurrent-range palynozones established by Morbey & Neves (1974) within the Karparthian facies of Ko¨ssen beds in the Kendelbach section (Figs 3, 6, 7). The integration of

palynological, ammonoid, conodont and isotope data from the candidate GSSP section (Krystyn & Ku¨rschner 2005; Krystyn et al. 2007a, b) confirms the transitional nature of the Norian–Rhaetian palynological boundary. The authors distinguished two palynological Zones, the lowermost containing typical older ‘Carnian’ elements (Ellipsovelatisporites rugosus, Enzonalasporites vigens, Partitisporites spp., Patinasporites toralis, Triadispora spp., Vallasporites ignacii), and the upper typified by the FADs of taxa such as Chasmatosporites sp., Limbosporites lundbladii and Quadraeculina anellaeformis. Additional taxa such as Classopollis meyerianus, Classopollis torosus, Granuloperculatipollis rudis, Ovalipollis pseudoalatus, Rhaetipollis germanicus, Ricciisporites tuberculatus, and Tsugaepollenites pseudomassulae, range throughout the two palynological Zones. An acme of dinoflagellate cysts (Heibergella, Noricysta, Rhaetogonyaulax) has been also recorded at the transition of the two palynozones, which also approximately coincide with the Cycloceltites –Vandaites ammonoid event (Figs 3 & 6). On the whole, the uppermost Triassic assemblages of the Boreal realm (e.g. Lund 1977; Pedersen & Lund 1980; Hochuli et al. 1989; Dybkjær 1991; Batten & Koppelhus 2002; Lindstro¨m & Erlstro¨m 2006) show affinity with those found in continental Europe with some exceptions, such as Limbosporites lundbladii, which is more common in the northern assemblages, Quadraeculina anellaeformis, which

Fig. 6. Main palynozones and palynological assemblages across the Norian–Rhaetian boundary. Z, Zone; Sz, Subzone; FO, first occurrence; LO, last occurrence. Palynozones from Morbey 1975: TR, Classopollis torosus– Granuloperculatipollis rudis; TL, Riccisporites tuberculatus –Hystrichosphaeridium langi; RG, Rhaetogonyaulax rhaetica– Rhaetipollis germanicus; Rk, Rhombodella kendelbachia; LL, Carnisporites lecythus– Zebrasporites laevigatus.

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Fig. 7. Schematic correlation of two GSSP candidate sections based on palynological and geochemical proxies across the Triassic– Jurassic boundary. LAD, Last appearance datum. Data from: Hounslow et al. 2004; Hillebrandt et al. 2007; Ku¨rschner et al. 2007; Warrington et al. 2008; FO, first occurrence; LO, last occurrence.

seems to appear earlier in the Boreal realm than in the European domain, and the Circumpolles group (Classopollis spp.), which has a wider distribution in central and southern Europe (Hochuli et al. 1989). Other taxa, such as Tsugaepollenites pseudomassulae, are not common in the Boreal realm, except for the British Rhaetian (Orbell 1973; Warrington et al. 2008). In contrast, the composition of Norian and Rhaetian miospore assemblages from the Middle East (e.g. east –central Iran) differs considerably from that of the representative European assemblages. The Norian microflora shows more affinity with the Gondwanan (e.g. Eastern Australia and New Zealand), while the Rhaetian assemblages, although maintaining a strong affinity with the microflora of the southern hemisphere, show a gradual enrichment in cosmopolitan species found in both the northern and southern hemispheres (Cirilli et al. 2005).

The Rhaetipollis germanicus assemblage zone A considerable number of palynological studies in the Rhaetian reference sections of the Boreal and European realms point to Rhaetipollis germanicus as a valuable biostratigraphical marker to typify the Rhaetian stage (Figs 2–3) (e.g. Rhaetipollis germanicus assemblage Zone of Orbell 1973;

Lund 1977; Schuurman 1977, 1979; Visscher et al. 1980; Koppelhus & Batten 2002; Warrington 2002; Barro´n et al. 2006; Yaroshenko 2007). However, the FAD of Rhaetipollis germanicus is yet uncertain, although the ammonoid-dated Norian palynomorph assemblages of Svalbard provided constraints for a Norian first appearance of this taxon (Smith 1982). Recent data from Austria documents its lowest entrance in the middle Rhaetian, although the entry level of Rhaetipollis germanicus remains unknown ‘because of the carbonaceous nature (facies) of the deposits. . .underlying the Rhaetian deposits (Ko¨ssen or Zlambach beds)’ (Krystyn et al. 2007a, p.193; Krystyn et al. 2007b, p.165). At the St. Audrie’s Bay section, Rhaetipollis germanicus makes its first occurrence in the Late Alaunian to Early Sevatian Twyning Mudstone Fm. (Hounslow et al. 2004); in the German Keuper its lowermost occurrence has been recorded in the lower Rhaetian (¼ Sevatian 2 of the Alpine domain) (Lund 1977, 2003; Schulz & Heunish 2005). The LAD of Rhaetipollis germanicus has been dated as late Rhaetian (Orbell 1973; Achilles 1981; Brenner 1986) as shown in Austria where the Rhaetipollis germanicus Assemblage Zone ranges into the upper part of the pre-planorbis beds (Morbey 1975; Ku¨rschner et al. 2007), although it was also rarely recorded in the oldest

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Jurassic strata (Warrington 2002; Traverse 2008) (Figs 2, 3, 6). Phase III of Schuurman (1979), regarded as transitional between phase II (Norian) and phase IV (late Rhaetian), registers the rapid decrease and disappearance of Rhaetipollis germanicus. In the Southern Alps (Italy), Rhaetipollis germanicus disappears just before the blooming of Krauselisporites reissingeri (Galli et al. 2007). Also, in Northern Europe (North Sea basin, East Greenland and Scania) the last occurrence of Rhaetipollis germanicus has been dated as ‘middle Rhaetian’ (Lund 1977; Pedersen & Lund 1980; Guy-Ohlson 1981). The diachronous first and last occurrences of this taxon, as well as those of other taxa, may be related to the climate-controlled provincialsm of the parent flora (cf. palynological section in Krystyn et al. 2007a, b).

The palynological Triassic – Jurassic boundary One of the main obstacles to describing and constraining the age of the palynomorph assemblages at the Tr–J boundary and at the base of Hettangian is the lack, until now, of a formally accepted GSSP section for this time interval. Many sections have been proposed during the last decade, some of them including palynostratigraphic zonations; examples include the St. Audrie’s Bay –Doniford Bay section, England (Warrington et al. 1994, 2008) and the Kuhijoch section, in the Tyrol of Austria (Hillebrandt et al. 2007). Recently, members of the International Subcommission on Triassic Stratigraphy (ISTS) and the International Subcommission on Jurassic Stratigraphy (ISJS) voted to recommend the Kuhjoch section, Karwendel Mountains, northern Calcareous Alps (Austria) (Hillebrandt et al. 2007) as the GSSP section and the Ferguson Hill Section, New York Canyon, Nevada (USA) as the ASSP (Guex et al. 2006; Lucas et al. 2007). The two proposals were formally approved by the ISJS in August 2008 (during the 33rd International Geological Meeting in Oslo) and are presently waiting for the approval of the International Commission on Stratigraphy (ICS). In the proposed GSSP section the Tr–J boundary lies 5.8 metres above the top of the Ko¨ssen Formation, within the level bearing the FO of Psiloceras cf. P. spelae, which Guex considered as the primary marker defining this boundary. The microfloral record across the boundary, which is chronologically constrained by carbon-isotope stratigraphy and marine (ammonoid and conodont) biostratigraphy, is distinguished by a few notable palynostratigraphic events that correlate with those of the Tiefengraben section in the eastern part of the Eiberg basin (Ku¨rschner et al. 2007) (Figs 2, 3, 6). Four palynomorph assemblage Zones have

been proposed. These are, in ascending order: (1) the Rhaetipollis –Limbosporites Zone (RL Zone), from the uppermost part of the Ko¨ssen Formation, is dominated by Classopollis meyerianus and minor Classopollis torosus, with Limbosporites lundbladii, Ovalipollis pseudoalatus, Rhaetipollis germanicus and Ricciisporites tuberculatus. Marine palynomorphs such as dinoflagellate cysts (Dapcodinium priscum, Rhaetogonyaulax rhaetica), and acritarchs (e.g. Cymatiosphaera polypartita, Micrhystridium sp.) are also present; (2) The Rhaetipollis–Porcellispora Zone (RPo Zone) from the Schattwald Beds is distinguished by the increase in Calamospora tener, Classopollis torosus, Classopollis meyerianus, Convolutispora microrugulata and Deltoidospora spp. in the lower –middle part, by the decrease in Classopollis spp. and the disappearance of Rhaetipollis germanicus in the upper part (lower part of the Schattwald Beds); (3) The Trachysporites-Porcellispora (TPo Zone) occurs between the Schattwald Beds and the beds containing Psiloceras cf. P. spelae, where Ovalipollis pseudoalatus has its highest appearance, Classopollis torosus temporarily disappears, and Heliosporites reissingeri, Porcellispora longdonensis and Trachysporites fuscus are abundant; (4) The Trachysporites –Heliosporites (TH Zone) lies above the FO of Psiloceras cf. P. spelae, where Classopollis torosus reappears, Cerebropollenites thiergartii makes its first occurrence, and Heliosporites reissingeri and Porcellispora longdonensis become dominant in the lower-middle part and decline in the upper part; and (5) The Trachysporites – Pinuspollenites Zone (TPi Zone), Jurassic in age, is identified by the appearance of Pinuspollenites minimus associated with abundant Classopollis spp., Heliosporites reissingeri and Trachysporites fuscus. It is noteworthy that the first occurrence of Cerebropollenites thiergartii is close to the Psiloceras cf. P. spelae horizon, within the lower part of the main negative d13C excursion (Ku¨rschner et al. 2007; Ruhl et al. 2009). On the basis of these constraints and by correlation with other previously established palynological Zones, Cerebropollenites thiergartii, at least in the European domain, seems to be the first and the only morphologically distinct, post-Triassic species suitable for interregional correlations between terrestrial and marine sections (Ku¨rschner et al. 2007). Recent data from additional key sections of the Eiberg Basin provided a detailed palynological zonation (Bonis et al. 2009) which is equivalent, except for minor differences, with the palynomorph assemblage zones from the Tiefengraben. The zonation scheme for the Eiberg Basin correlates well with those of Morbey (1975) and Schuurman (1977, 1979) for the Alpine realm (Figs 3, 6). The Rhaetipollis–Limbosporites Zone correlates to the

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MI Subzone of Morbey (1975) and to phase III of Schuurman (1977, 1979), the assemblages of which contain Classopollis spp., Cingulizonates rhaeticus, Limbosporites lundbladii, Ovalipollis pseudoalatus, Rhaetipollis germanicus and Ricciisporites tuberculatus. The Rhaetipollis –Porcellispora and the Trachysporites–Porcellispora Zones are the equivalent of the FG Subzone of Morbey, including the transition from the Rhaetian to the Hettangian Stage, defined by the LAD of Densosporites fissus, Retitriletes gracilis, Triancoraesporites ancorae, and by abundant Classopollis torosus, Heliosporites reissingeri and Ricciisporites tuberculatus. The Rhaetipollis –Porcellispora and the Trachysporites– Porcellispora Zones correspond to phase IV of Schuurman (1977, 1979), which is distinguished by the rapid disappearance of Ovalipollis, Rhaetipollis and taeniate bisaccate pollen and by abundant Heliosporites reissingeri. The Trachysporites–Heliosporites and the Trachysporites–Pinuspollenites Zones fit Jurassic phase V of Schuurman (1979), dominated by Classopollis spp. and distinguished by the lack of typical Rhaetian forms except for some taxa (e.g. Heliosporites, Retitriletes). The established carbon-isotope curve, as well as palynostratigraphic and other biostratigraphical events of the Austrian section, also correlate with the section at St Audrie’s Bay on the Somerset coast of England (Fig. 7), previously a GSSP candidate section (Warrington et al. 1994, 2008) and others nearby in Somerset and South Wales (Orbell 1973; Warrington 1974; Hounslow et al. 2004), where the Tr–J boundary has been conventionally placed at the lowest occurrence of the Hettangian ammonite Psiloceras planorbis. Palynological data integrated with other fossils and proxies such as magnetostratigraphy and chemostratigraphy (Warrington 2002; Hounslow et al. 2004; Warrington et al. 2008), show miospore assemblages from the Triassic Penarth Group dominated by Circumpolles (Classopollis, Geopollis, Gliscopollis, Granuloperculatipollis), Ricciisporites tuberculatus, and minor Ovalipollis pseudoalatus, Rhaetipollis germanicus, which have their LADs at the top most of the Penarth Group, below the FAD of Psiloceras. Ricciisporites tuberculatus ranges into the basal part of the Lias Group, and Kraeuselisporites reissingeri and Quadraeculina anellaeformis range from the Penarth Group into beds that contain Hettangian ammonites, while the Triassic-type palynomorphs rapidly disappear. There are no other significant first occurrences that can be specifically correlated with the Tr–J boundary (Warrington 1974, 2002), except for Cerebropollenites thiergartii and Cerebropollenites macroverrucosus, whose occurrences in the Somerset succession are in the Hettangian Psiloceras planorbis beds, as in other European

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sections (Clement-Westerhof et al. 1974; Morbey 1975, 1978; Van Erve 1977; Visscher et al. 1980; Fisher & Dunay 1981; Ku¨rschner et al. 2007). Additionally, the palynological Zones of the Austrian section match those of central and southern European sections where similar assemblages have been recorded across the Tr –J boundary in Spain (Baudelot & Taugourdeau-Lantz 1986; Vachard et al. 1990; Calvet et al. 1993; Fre´chengues et al. 1993; Barro´n et al. 2006; Gomez et al. 2007), in the Lusitanian Basin and France (Adloff et al. 1974; Adloff & Doubinguer 1982; Grignac & Taugourdeau-Lantz 1982; Peybernes et al. 1988; Fre´chengues et al. 1993), in Italy (Cirilli et al. 1993, 1994; Galli et al. 2007), and in the Western Carpathians and Hungary (Go¨tz et al. 2009; Ruckwied & Go¨tz 2009). The Austrian Zones also correlate well with the palynozonations of Northern Europe and Scandinavia (Lund 1977, 2003) (Fig. 3). The Triassic Rhaetipollis–Limbosporites and Rhaetipollis– Porcellispora Zones compare with the Rhaetipollis –Limbosporites Zone of Lund (1977); the Trachysporites –Porcellispora Zone shows a relationship with the Ricciisporites–Polypodiisporites Zone of Lund (1977), primarily in the absence of Rhaetipollis and the abundance of Convolutispora microrugulata (¼ Polypodiisporites polymicroforatus in Lund 1977). The Early Jurassic Trachysporites–Heliosporites and Trachysporites –Pinuspollenites Zones compare with the Pinuspollenites–Trachysporites Zone of Lund (1977). In most of the European domain, the Tr–J boundary seems to be characterized by only a minor extinction/turnover of the terrestrial macroflora and microflora against a background of more gradual change (e.g. Orbell 1973; Schuurman 1979; Pedersen & Lund 1980; Fisher & Dunay 1981; Knoll 1984; Ash 1986; Hallam & Wignall 1997; Hallam 2002; Tanner et al. 2004; Lucas & Tanner 2004, 2007, 2008; Ku¨rschner et al. 2007). A floral change has been documented at the Tr–J boundary in northwestern Europe, for example, Greenland (McElwain et al. 1999, 2007; McElwain & Punyasena 2007), Germany and Sweden (Van de Schootbrugge et al. 2009). In the latter two sites Van de Schootbrugge et al. (2009) record a floral change at the base of the Triletes Beds, where typical Rhaetian pollen, such as Ricciisporites tuberculatus and Rhaetipollis germanicus, decrease until they disappear. They are replaced in the earliest Hettangian by assemblages dominated by pollen from Cheirolepidiaceae and Taxodiaceae conifers (e.g. Classopollis spp., Perinopollenites elatoides, Pinuspollenites minimus) and bisaccates from corystosperm seed ferns (e.g. Alisporites spp.). accompanied by an acme of lycopodiophyte spores (e.g. Kraeuselisporites reissingeri).

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In the Danish part of the Danish Basin, the Tr–J transition is associated with an abundance of the cheirolepidacean conifer pollen Classopollis spp., and follows a mass-occurrence of dinoflagellate cysts (Dybkjær 1988, 1991; Lindstro¨m & Erlstro¨m 2006). Variable abundances of Classopollis-type pollen, commonly associated with hot and/or arid climate conditions (Vakhrameev 1991; McElwain et al. 2007), are documented in numerous sections spanning the Tr –J transition, such as: Great Britain (Warrington & Ivimey-Cook 1990; Hounslow et al. 2004), central and southern Europe (see Ku¨rschner et al. 2007 for updated references), north Africa (Marzoli et al. 2004), and also in the southern hemisphere, such as in Australia (Backhouse & Balme 2002; Grice et al. 2005), Timor, Indonesia (Martini et al. 2000, 2004) and Antarctica (Shu et al. 2000). In the Eastern North America basins (i.e. Newark Supergroup basins) the palynological Tr–J boundary has been placed since the 1970s in the strata immediately beneath the lowest CAMP basalt flow (e.g. the Orange Mt. Basalt in the Newark basin), to coincide with a palynological turnover event that is characterized by a low diversity palynomorph assemblage recording a significant loss of Late Triassic taxa (Fowell & Olsen 1993; Fowell et al. 1994; Fowell & Traverse 1995; Olsen et al. 2002; Whiteside et al. 2007). This rapid turnover has been primarly defined on the LADs of Ovalipollis ovalis, Patinasporites densus, Vallasporites ignacii, on the increase in the Classopollis type pollens such as Classopollis meyerianus, Classopollis murphyae, Classopollis torosus, and on a blooming of trilete fern spores (fern spike) (e.g. Cornet & Traverse 1975; Cornet & Olsen 1985; Fowell & Olsen 1993; Fowell et al. 1994; Fowell & Traverse 1995; Whiteside et al. 2007). Additionally, this putative Tr –J boundary is in close stratigraphic proximity to a magnetic reversal (chron E23r: Olsen et al. 2002), a negative carbon isotope excursion (Olsen et al. 2002), and is also characterized (in the Newark and Fundy basins) by a moderate Ir anomaly (Olsen et al. 2002). However, the above palynostratigraphic criteria for placing the Tr–J boundary have been strongly critiqued (Gradstein et al. 1994; van Veen 1995; Lucas & Tanner 2007, 2008; Cirilli et al. 2009), primarly because the LAD of Patinasporites densus in the northern and southern hemispheres is documented as late Norian to early Rhaetian (Sevatian– Rhaetian boundary). The large distribution of this taxon in different palaeoclimate belts (Buratti & Cirilli 2007) would exclude the possibility of floral provincialism (Fowell & Olsen 1993) controlling a diachronous distribution of the parent plant in North American and Tethyan domains. Additionally, the increase in the Classopollis type pollens

is, as seen above, a widespread palynological event spanning the Tr–J transition in the northern and southern hemispheres and not only Jurassic in age. Recent data from the Fundy Basin of Nova Scotia, Canada, document the presence of Triassic sporomorphs in the Scots Bay Member, which overlies the North Mountain Basalt, and therefore above the last occurrence of Patinasporites densus (Cirilli et al. 2009). In the Blomidon Formation, at Partridge Island section, no significant palynological decline or turnover has been observed below the North Mountain Basalt that can be attributed to a mass extinction event. Furthermore, although a true fern spike is absent at this site, a level enriched in fern spores (e.g. Converrucosisporites cameronii, Dictyophyllidites harrisii and Dictyophyllidites sp.) is recorded 15 cm below the putative Tr– J boundary, but not above the boundary as in the Newark Basin (Fowell & Olsen 1993; Fowell et al. 1994; Whiteside et al. 2007), suggesting the relative abundance of fern spores may record a short-term episode of climate change and atmospheric acidification (i.e. related to CAMP eruptions) rather than a global event of recolonization after mass exctinction (Cirilli et al. 2009; Van de Schootbrugge et al. 2009).

Upper Triassic miospore assemblages of the southern hemisphere As a consequence of Late Triassic provincialism, different palynological zonations, mostly based on integrated dinoflagellate and spore-pollen palynostratigraphy, have been provided for eastern and western Australia (Helby et al. 1987; Brenner 1992; Backhouse & Balme 2002; Grice et al. 2005), Queensland (de Jersey 1975), Victoria Land, Antarctica (Kyle & Schopf 1982), and New Zealand (de Jersey & Raine 1990) (Fig. 8). The assemblages from New Zealand and eastern Australia, apart from a limited occurrence in the Galilee Basin of Queensland, can be assigned to the Ipswich Microflora, whereas the assemblages of northwestern Australia and Timor belong to the Onslow Microflora (Foster et al. 1994; Martini et al. 2000, 2004; Buratti & Cirilli 2007). Correlation between these palynological zonal schemes and the standard geological timescale has been based on scattered data of marine biota, primarly foraminiferans, and sparse occurrences of ammonites, conodonts and/or dinoflagellate cysts (Helby et al. 1987; Nicoll & Foster 1994, 1998).

Western Australia In western Australia three important Late Triassic palynological Zones have been proposed to correlate with conodont biozones (Dolby & Balme

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299

Fig. 8. Schematic comparison of the main palynozonations across the Late Triassic and Early Jurassic in the southern hemisphere. (1) de Jersey & Raine 1990; (2) de Jersey 1975; de Jersey & Raine 1990; (3) Dolby & Balme 1976; Helby et al. 1987; de Jersey & Raine 1990; Nicoll & Foster 1998; Backhouse & Balme 2002; Grice et al. 2005.

1976; Helby et al. 1987; Nicoll & Foster 1998; Backouse & Balme 2002; Grice et al. 2005) (Fig. 8): (1) the mostly Carnian to earliest Norian Samaropollenites speciosus Oppel Zone, marked by a significant decline of Enzonalasporites vigens and by an increase of Falcisporites australis; (2) the Norian Minutosaccus crenulatus Oppel Zone, defined by a considerable decline of the Tethyan taxa such as Enzonalasporites vigens, Patinasporites densus and Samaropollenites speciosus; and (3) the Rhaetian, probably extending into Hettangian, Ashmoripollis reducta Oppel Zone, characterized by the FAD of Zebrasporites, by the decline of bisaccates such as Falcisporites spp., and by the dominance of Classopollis spp. at the top. The transition from the Ashmoripollis reducta to Classopollis torosus Zones characterizes the Tr– J boundary, defined by the last occurrence of Triassic spores (Backhouse & Balme 2002; Grice et al. 2005).

Eastern Australia and Indian Peninsula A different Late Triassic palynozonation has been established for eastern Australia based on two long range Zones (Fig. 8): the Craterisporites rotundus Oppel Zone, assigned to the Carnian (de Jersey 1975) and recognized by the FAD of Craterisporites

rotundus at its base, by the dominance of Duplexisporites problematicus (¼ Striatella seebergensis) and Falcisporites australis and by the FAD of Polycingulatisporites crenulatus at the top; and the Polycingulatisporites crenulatus Oppel Zone, defined by the FAD of Polycingulatisporites crenulatus at the base, by a rapid decline in Falcisporites australis, by the LAD of Playfordiaspora velata and by a rapid increase in Classopollis torosus at the top. This Zone has been variously assigned to the Rhaetian– Hettangian by de Jersey (1975) and to the Norian extending to Hettangian by Helby et al. (1987). The Polycingulatisporites crenulatus Zone of Eastern Australia has been correlated with the Arcuatipollenites tethyensis Zone of the Indian Peninsula (Tripathi 2000), where the Upper Triassic –Lower Jurassic assemblages of the Dubrajpur Formation commonly contain Araucariacites, Callialasporites, Classopollis, Dictyophyllidites, Infernopollenites, Minutosaccus, Nevesisporites and Staurosaccites.

New Zealand and Tanzania The palynological zonations of eastern Australia may be reasonably compared with New Zealand

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miospore zonations (cf. Murihiku Supergroup: de Jersey & Raine 1990), considering the distribution of the Late Triassic floral and microfloral provincialism (Fig. 1). Four palynological Zones and one Subzone have been established from the Late Triassic (Oretian, Otamitan, Warepan, Otapirian) to the Hettangian (Aratauran) interval (de Jersey & Raine 1990), correlated by invertebrate faunas with the international stratigraphic stages (Fig. 8). In ascending order they are: (1) the long-ranging late Ladinian to early Norian Annulispora folliculosa Zone, defined by the FAD of Annulispora folliculosa and, in its upper part by the FAD of Annulispora microannulata, which identifies the late Carnian –early Norian Annulispora microannulata Subzone. This Subzone contains other FADs such as Alisporites warepanus, Aratrisporites flexibilis, Rogalkaisporites spp., Stereisporites antiquasporites and Striatella seebergensis in the uppermost part of the Annulispora microannulata Subzone (early Norian); (2) The early to late Norian Polycingulatisporites crenulatus Zone, marked by the FAD of Polycingulatisporites crenulatus and by the LAD of Lundbladispora denmeadi and Aratrisporites flexibilis. Polycingulatisporites mooniensis enters in the lower part of the Zone in the assemblage with Camarozonosporites rudis and Retitriletes austroclavatidites (de Jersey & Raine 1990); (3) The Rhaetian Foveosporites moretonensis Zone is defined by the acme of Densoisporites psilatus, by the LAD of Aratrisporites spp., by the scarce presence of Perinopollenites elatoides and by the presence of Rugaletes awakinoensis in association with Classopollis sp. cf. C. chateaunovi. Limbosporites antiquus is present in the upper part of the Zone; and (4) The Hettangian Retitriletes austroclavatidites Zone distinguished by the FAD of Retitriletesaustroclavatidites, by the LAD of Aratrisporites spp., by the acme of Densoisporites psilatus and by abundant Classopollis spp. n the upper part of the Zone. Taxa such as Annulispora folliculosa and Annulispora microannulata are commonly present as accessory forms also in the middle Carnian to early Norian of eastern Australia (Craterisporites rotundus Zone: de Jersey 1975), and in the Carnian to early Norian of western Australia (Dolby & Balme 1976). Polycingulatisporites mooniensis is another cosmopolitan form of the southern hemisphere that is present as an accessory taxon in the middle Norian of eastern Australia (Queensland), New Zealand (Stevens 1981) and New Caledonia (de Jersey & Grant-Mackie 1989). An Ipswich microfloral affinity has also been recognized in Tanzania where the Rhaetian microfloral assemblages of the Mkuju Formation share several taxa with the microfloras from New Zealand, southern and Eastern Australia and Malagasy such as Duplexisporites problematicus

(¼ Striatella seebergensis), Falcisporites australis, Nevesisporites limatulus, Polycingulatisporites crenulatus and Punctatisporites leighensis (Hankel 1987, 1993).

Argentina The palynomorph assemblages of Argentina are comparable with the Ipswich microflora (Zavattieri & Batten 1996), and have close similarities with those of eastern Australia, South Africa, Antarctica and other Gondwanan domains. The oldest palynological records are from the Cacheuta Formation in Mendoza Province (Orlando 1954; Jain 1968), and from various Triassic basins of western Argentina stratigraphically correlated through palynostratigraphy (Yrigoyen & Stover 1969). A Middle Triassic to Carnian microflora reported from the Cuyo Basin is composed of a large number of spores (e.g. Aratrisporites compositus, Auritulinasporites scanicus, Cadargasporites cuyanensis, Calamospora spp., Camerosporites verrucosus, Dictyophyllidites mortonii, Guthoerlisporites cancellosus, Punctatosporites spp., Rugulatisporites spp.), abundant bisaccate (e.g. Alisporites spp., Falcisporites nuthallensis, Klausipollenites schaubergeri, Minutosaccus acutus, Platysaccus spp., Protodiploxypinus spp., Triadispora spp.) and other gymnosperm pollen (e.g. Accinctisporites spp., Cycadopites spp.). The Carnian– earliest Norian has been documented from the Los Rastros and Ischigulasto Formations in the Ischigualasto–Villa Unixn Basin, San Juan –La Rioja provinces, on the basis of palynomorph assemblages containing Alisporites spp., Cacheutasporites wielandii, Cadargasporites cuyanensis, Convolutispora microrugulata, Dictyophyllidites mortonii, Klausipollenites spp., Monosulcites spp., Osmundacidites spp., Platysaccus spp., Punctatosporites walkomii, Rugulatisporites neuquenensis, Spheripollenites spp. and Vesicaspora spp., among others (Zavattieri & Batten 1996; Ottone et al. 2005). In the North Patagonian Basin, Rı´o Negro Province, the continental succession of the Cerro Puntudo Formation yielded a rich palynofloral assemblage showing a Norian to Rhaetian transitional composition. Spores such as Annulispora folliculosa, Apiculatisporis lentus, Aratrisporites spp., Craterisporites rotundus, Neoraistrickia densata, Osmundacidites spp., Striatella seebergensis, pollen such as Alisporites spp., Indusiisporites parvisaccatus, Klausipollenites decipiens, Laricoidites intragranulosus, Triadispora spp., Voltziaceaesporites heteromorphus and a few Classopollis in the upper part of the succession, co-occur with the oldest Middle Triassic –earliest Late Triassic taxa. The youngest Triassic Argentinian assemblages (Norian –early Rhaetian), have been recorded from the Chihuido Formation at the

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Fig. 9. Selected Carnian sporomorphs from the northern hemisphere. (1– 4) Patinasporites densus Leschik 1956 emend. Scheuring 1970; (5 –8) Enzonalasporites vigens Leschik 1956; (9– 12) Vallasporites ignacii Leschik 1956; (13, 14) Praecirculina granifer (Leschik 1956) Klaus 1960; (15) Pseudoenzonalasporites summus Scheuring 1970; (16) Duplicisporites granulatus Leschik 1956 emend. Scheuring, 1970; (17) Paracirculina quadruplicis Scheuring 1970; (18, 19) Lagenella martini (Leschik 1956) Klaus 1960; (20) Aratrisporites tenuispinosus Playford 1965.

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Fig. 10. Selected Carnian sporomorphs from the northern hemisphere. (1– 4) Camerosporites secatus Leschik 1956 emend. Scheuring, 1978; (5– 9) Samaropollenites speciosus Goubin 1965; (10, 11) Ovalipollis pseudoalatus (Thiergart 1949) Schuurman 1976; (12, 13) Staurosaccites quadrifidus Dolby in Dolby & Balme 1976; (14, 15) Lueckisporites’ cf. L. singhii Balme 1970; (16) Lunatisporites acutus Leschik 1956; (17) Kyrtomisporis ervii Van der Eem 1983; (18) Reticulatisporites dolomiticus Blendinger 1988; (19) Vesicaspora fuscus (Pautsch 1958) Morbey 1975; (20) Gordonispora fossulata (Balme 1970) Van der Eem 1983.

Llantenes locality, in the Malargxe Basin, southern Mendoza Province, which yielded a microflora mainly composed of Baculatisporites comaumensis, Dictyophyllidites mortonii, Guthoerlisporites cancellosus, Leptolepidites spp., Neoraistrickia taylorii, Polypodiisporites ipsviciensis, and pollen such as Alisporites spp. and Classopollis spp.

Conclusions The compositional changes of miospore assemblages from the Late Triassic to Early Jurassic appear primarily to be latitudinally, and therefore largely climatically controlled. Whereas palynostratigraphy provides reliable results on regional stratigraphic

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Fig. 11. Selected Norian and Rhaetian sporomorphs from the northern hemisphere. (1– 4) Rhaetipollis germanicus Schulz 1967; (5, 6) Tsugaepollenites pseudomassulae (Ma¨dler 1964) Morbey 1975; (7, 8, 12) Ricciisporites tuberculatus Lundblad 1954; (9, 10) Cerebropollenites macroverrucosus (Thiergart 1949) Schulz 1967; (11, 15) Heliosporites reissingeri (Harris 1957) Muir & Van Konijnenburg-Van Cittert 1970; (13, 14) Trachysporites fuscus Nilsson 1958; (16) Camarozonosporites laevigatus Schulz 1967; (17) Acanthotriletes varius Nilsson 1958; (18) Uvaesporites argenteaeformis (Bolkhovitina 1953) Schulz 1967; (19) Lycopodiacidites rugulatus (Couper 1958) Schulz 1967; (20) Ovalipollis ovalis Krutzch 1955.

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Fig. 12. Selected Norian and Rhaetian sporomorphs from the northern hemisphere. (1 –4, 7– 11, 13, 17) Classopollis torosus Reissinger 1950; (5) Classopollis murphyae (Cornet & Traverse 1975) Traverse 2008; (6, 12, 14–16, 19) Classopollis meyerianus (Klaus 1960) de Jersey 1973; (18) Granuloperculatipollis rudis Venkatachala & Go´cza´n 1964 emend. Morbey 1975.

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correlations, global correlations are complicated by many factors, such as differences in the composition of coeval assemblages, and by discrepancies in the taxonomy of sporomorphs. Nevertheless, although long distance correlations are commonly limited by climatic and environmental factors, palynology may be very useful in solving regional correlations. This is particularly evident for the European domain, where the palynostratigraphy is well documented and the available data provide a detailed regional picture of the Late Triassic and Early

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Jurassic European microfloral composition. Furthermore, microfloral associations provide useful supplementary information concerning the distribution of parent flora and the pattern of floral provinces. In the future, multidisciplinary studies based on the integration of different biostratigraphical and geochronological tools (e.g. magnetostratigraphy, chemostratigraphy) should better constrain the ages of palynomorph assemblages and improve the potentiality of palynology for world-wide correlations among marine and non-marine sections.

Fig. 13. Late Triassic sporomorphs from the southern hemisphere and Iran. (1, 2) Camerosporites secatus Leschik 1956 emend. Scheuring, 1978; (3) Patinasporites densus Leschik 1956 emend. Scheuring, 1970; (4) Vallasporites ignacii Leschik 1956; (5, 6) Enzonalasporites vigens Leschik 1956; (7, 8) Partitisporites novimundanus Leschik 1956; (9, 10) Samaropollenites speciosus Goubin 1965; (11) ‘Lueckisporites’ cf. L. singhii Balme 1970; (12) Brodispora striata Clarke 1965; (13) Perinopollenites elatoides Couper 1958; (14) Ephedripites primus Klaus 1963; (15) Trachysporites fuscus Nilsson 1958; (16) Polycingulatisporites mooniensis de Jersey & Paten 1964.

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Taxonomic note Systematic nomenclature of the Corollina – Circulina–Classopollis complex: Traverse (2004, 2008) proposed, by means of the procedures outlined in the International Code of Botanical Nomenclature (Greuter et al. 2000), to conserve the morphogeneric name Classopollis against Corollina and Circulina, which are commonly used for the same fossil pollen forms. This proposal has been accepted by the International Botanical Congress in 2005. Therefore, in the present contribution the generic name Classopollis will be used to replace Corollina and Circulina for the species: Corollina meyeriana (¼ Gliscopollis meyeriana), Corollina torosa, Corollina murphyi, Corollina sp. cf. C. chateaunovi (cf. also Cornet & Waanders 2006). Figures 9– 13 illustrate representative Late Triassic sporomorphs. First of all many thanks to Spencer Lucas as editor of this book. I gratefully acknowledge Nicoletta Buratti (Perugia University, Italy) for having made available her palynological collection and for her comments and suggestions, Roberto Rettori (Perugia University) for his constructive discussions on Triassic biostratigraphy and Lawrence Tanner (Le Moyne College, Syracuse, N.Y.) for his friendly encouragement and for reviewing the manuscript. I am also deeply grateful to the other reviewers, W. M. Ku¨rschner (Utrecht University, NL), J. Ogg (Purdue University, Indiana, USA) and J. B. Riding (British Geological Survey, UK) for their helpful comments and suggestions that surely improved the quality of this article. Part of this study was financially supported by the Italian Minister of Scientific Research (PRIN05– 07, S. Cirilli).

Appendix I Species list of cited taxa Acanthotriletes varius Nilsson 1958, Pl.III, 17 Alisporites warepanus Raine in de Jersey & Raine 1990 Annulispora folliculosa (Rogalska 1954) de Jersey 1959 Annulispora microannulata de Jersey 1962 Apiculatisporis lentus Playford 1982 Aratrisporites compositus Volkheimer & Zavattieri 1991 Aratrisporites flexibilis Playford & Dettmann 1965 Aratrisporites tenuispinosus Playford 1965, Pl.I, 20 Arcuatipollenites tethyensis (Vijaya & Tiwari 1988) Tiwari & Vijaya 1995 Ashmoripollis reducta Helby 1987 Aulisporites cf. A. astigmosus (Leschik 1956) Klaus 1960 Auritulinasporites scanicus Nilsson 1958 Baculatisporites comaumensis (Cookson 1953) Potonie 1956 Brodispora striata Clarke 1965 Pl.V, 12 Cacheutasporites wielandii Jain 1968 Cadargasporites cuyanensis Azcuy & Longobucco 1983 Calamospora tener (Leschik, 1956) Madler 1964 Callialasporites dampieri (Balme 1957) Dev 1961

Camarozonosporites laevigatus Schulz 1967, Pl.III, 16 Camarozonosporites rudis (Leschik 1955) Klaus 1960 Camerosporites pseudoverrucatus Scheuring 1970 Camerosporites secatus Leschik 1956 emend. Scheuring 1978, Pl.II, 1 –4; Pl.V, 1, 2 Camerosporites verrucosus Ma¨dler 1964 Ceratosporites helidonensis de Jersey 1971 Cerebropollenites macroverrucosus (Thiergart 1949) Schulz 1967, Pl.III, 9, 10 Cerebropollenites thiergartii Schulz 1967 Chordasporites sp. cf. C. australiensis de Jersey 1962 Cingulizonates rhaeticus (Reinhardt 1961) Schulz 1967 Classopollis meyerianus (Klaus 1960) de Jersey 1973, Pl.IV, 6, 12, 14–16, 19 Classopollis murphyae (Cornet & Traverse 1975) Traverse 2008, Pl.IV, 5 Classopollis sp. cf. C. chateaunovi Reyre 1970 Classopollis torosus Reissinger 1950, Pl.IV, 1 –4, 7– 13, 17 Concentricisporites cf. C. bianulatus (Neves 1961) Antonescu 1970 Convolutispora microrugulata Schulz 1967 Converrucosisporites cameronii (de Jersey 1962) Playford & Dettmann 1965 Craterisporites rotundus de Jersey 1970 Cycadopites stonei Helby 1987 Densosporites fissus (Reinhardt 1964) Schulz 1967 Densoisporites psilatus (de Jersey 1964) Raine & de Jersey 1988 Dictyophyllidites harrisii Couper 1958 Dictyophyllidites mortonii (de Jersey 1959) Playford & Dettmann 1965 Duplexisporites problematicus (Couper 1958) Playford & Dettmann 1965 Duplicisporites continuus Praehauser-Enzenberg 1970 Duplicisporites granulatus Leschik 1956 emend. Scheuring, 1970, Pl.I, 16 Duplicisporites verrucosus Leschik 1955 emend. Scheuring, 1978 Echinitosporites iliacoides Schulz & Krutzsch 1961 Ephedripites primus Klaus 1963, Pl.V, 14 Ellipsovelatisporites plicatus Klaus 1960 Ellipsovelatisporites rugosus Scheuring 1970 Enzonalasporites vigens Leschik 1956, Pl.I, 5– 8; Pl.V, 5, 6 Equisetosporites chinleanus Daugherty 1941 Falcisporites australis (de Jersey 1962) Stevens 1981 Falcisporites nuthallensis (Clarke 1965) Balme 1970 Foveosporites moretonensis de Jersey 1964 Gordonispora fossulata (Balme 1970) Van der Eem 1983, Pl.II, 20 Granuloperculatipollis rudis Venkatachala & Go´cza´n 1964 emend. Morbey 1975, Pl.IV, 18 Guthoerlisporites cancellosus Playford & Dettmann 1965 Heliosaccus dimorphus Ma¨dler 1964 Heliosporites reissingeri (Harris 1957) Muir & Van Konijnenburg-Van Cittert 1970, Pl.III, 11, 15 Indusiisporites parvisaccatus (de Jersey 1959) de Jersey 1963 Infernopollenites parvus Scheuring 1970

PALYNOLOGY AND PALYNOSTRATIGRAPHY Klausipollenites decipiens Jansonius 1962 Klausipollenites schaubergeri (Potonie & Klaus 1954) Jansonius 1962 Kraeuselisporites reissingeri (Harris 1957) Morbey 1975 Kyrtomisporis ervii Van der Eem 1983 Pl.II, 17 Lagenella martini (Leschik 1956) Klaus 1960 Pl.I, 18, 19 Laricoidites intragranulosus Bharadwaj & Singh 1964 Lycopodiacidites rugulatus (Couper 1958) Schulz 1967, Pl.III, 19 Limbosporites antiquus (de Jersey 1964) de Jersey & Raine 1990 Limbosporites lundbladii Nilsson 1958 Lueckisporites” cf. L. singhii Balme 1970 Pl.II, 14, 15; Pl.V, 11 Lundbladispora denmeadi (de Jersey 1962) Playford & Dettmann 1965 Lunatisporites aff. L. noviaulensis (Leschik 1956) de Jersey 1979 Lunatisporites acutus Leschik 1956 Pl.II, 16 Lunatisporites rhaeticus (Schulz 1967) Warrington 1974 Minutosaccus acutus Ma¨dler 1964 Minutosaccus crenulatus Dolby in Dolby & Balme 1976 Neoraistrickia densata Filatoff 1975 Neoraistrickia taylorii Playford & Dettmann 1965 Nevesisporites limatulus Playford 1977 Nevesisporites vallatus de Jersey & Paten 1964 Ovalipollis pseudoalatus (Thiergart 1949) Schuurman 1976 Pl.II, 10, 11 Ovalipollis ovalis Krutzch 1955, Pl.III, 20 Paracirculina quadruplicis Scheuring 1970, Pl.I, 17 Paracirculina scurrilis Scheuring 1970 Partitisporites novimundanus Leschik 1956, Pl.V, 7, 8 Patinasporites densus Leschik 1956 emend. Scheuring, 1970 (Pl.I, 1– 4; Pl. V, 3) Patinasporites toralis Leschik 1956 Perinopollenites elatoides Couper 1958, Pl.V, 13 Perinosporites thuringiacus Schulz 1962 Pinuspollenites minimus (Couper 1958) Kemp 1970 Playfordiaspora velata (Leschik 1955) Stevens 1981 Polycingulatisporites crenulatus Playford & Dettmann 1965 emend. McKellar 1974 Polycingulatisporites mooniensis de Jersey & Paten 1964, Pl.V, 16 Polypodiisporites ipsviciensis (de Jersey 1965) Playford & Dettmann 1965 Polypodiisporites polymicroforatus (Orlowska-Zwolin˜ska 1966) Lund 1977 Porcellispora longdonensis (Clarke 1965) Scheuring, 1970, emend. Morbey, 1975 Praecirculina granifer (Leschik, 1956) Klaus 1960, Pl.I, 13, 14; Protohaploxypinus sp. cf. P. microcorpus (Schaarschmidt 1963) Clarke 1965 Pseudoenzonalasporites summus Scheuring 1970, Pl.I, 15 Punctatisporites leighensis Playford & Dettmann 1965 Punctatosporites walkomii de Jersey 1962 Quadraeculina anellaeformis Maljavkina 1949 Reticulatisporites dolomiticus Blendinger 1988 Pl.II, 18

307

Retisulcites perforatus (Madler 1964) Scheuring 1970 Retitriletes austroclavatidites (Cookson 1953) Do¨ring, Krutzsch, Mai & Schulz 1963 Retitriletes gracilis Schulz 1967 Retitriletes semimuris (Danze´-Corsin & Laveine 1963) McKellar 1974 Rhaetipollis germanicus Schulz 1967, Pl.III, 1– 4 Ricciisporites tuberculatus Lundblad 1954, Pl.III, 7– 9 Rugaletes awakinoensis Raine in de Jersey & Raine 1990 Rugulatisporites neuquenensis Volkheimer 1972 Samaropollenites speciosus Goubin 1965 Pl.II, 5 –9; Pl.V, 9, 10 Semiretisporis gothae Reinhardt 1962 Staurosaccites quadrifidus Dolby in Dolby & Balme 1976 Pl.II, 12, 13 Stereisporites antiquasporites (Wilson & Webster 1946) Dettmann 1963 Striatella seebergensis Madler 1964 Trachysporites fuscus Nilsson 1958, Pl.III, 13–14, Pl. V, 15) Triancoraesporites ancorae (Reinhardt 1961) Schulz 1967 Triancoraesporites reticulatus Schulz 1962 Tsugaepollenites pseudomassulae (Ma¨dler 1964) Morbey 1975, Pl.III, 5–6 Uvaesporites argenteaeformis (Bolkhovitina 1953) Schulz 1967 Pl.III, 18 Uvaesporites gadensis Praehauser-Enzenberg 1970 Uvaesporites reissingerii (Reinhardt 1962) Lund 1977 Vallasporites ignacii Leschik 1956, Pl.I, 9 –12, Pl. V, 4 Vesicaspora fuscus (Pautsch 1958) Morbey 1975, Pl.II, 19 Voltziaceaesporites heteromorphus Klaus 1964 Weylandites magmus (Bose & Kar 1975) Van der Eem 1983

Dinoflagellate cysts Cymatiosphaera polypartita Morbey 1975 Dapcodinium priscum Evitt 1961 emend. Below 1987 Rhaetogonyaulax rhaetica (Sarjeant 1963) Loeblich & Loeblich 1968

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V ISSCHER , H., S CHUURMAN , W. M. L. & V AN E RVE , A. W. 1980. Aspects of a palynological characterization of Late Triassic and Early Jurassic ‘Standard’ units of chronostratigraphical classification in Europe. Proceedings IV International Palynological Conference, Lucknow, 2, 281–287. V ISSCHER , H. & B RUGMAN , W. 1981. Ranges of selected palynomorphs in the Alpine Triassic of Europe. Review of Palaeobotany and Palynology, 34, 115–128. V ISSCHER , H. & V AN DER Z WAN , C. J. 1981. Palynology of the Circum-Mediterranean Triassic: Phytogeographical and Palaeoclimatological Implications Geologische Rundschau, 70, 625–634. W ARD , P. D., G ARRISON , G. H., H AGGART , J. W., K RING , D. A. & B EATTIE , M. J. 2004. Isotopic evidence bearing on Late Triassic extinction events, Queen Charlotte Islands, British Columbia, and implications for the duration and cause of the Triassic– Jurassic mass extinction. Earth Planetary Science Letters, 224, 589–600. W ARRINGTON , G. 1974. Studies in the palynological biostratigraphy of the British Trias. Reference sections in West Lancashire and North Somerset. Review of Palaeobotany and Palynology, 17, 133 –147. W ARRINGTON , G. 2002. Triassic spores and pollen. In: J ANSONIUS , J. & M C G REGOR , D. C. (eds) Palynology: Principles and Applications. American Association of Stratigraphic Palynologists Foundation, 2nd edition, 2, 755– 766. W ARRINGTON , G. & I VIMEY -C OOK , H. I. 1990. Biostratigraphy of the Late Triassic and Early Jurassic: a review of type sections in southern Britain. Cahiers Universitaires Catholiques Lyon, Se´rie Science, 3, 207– 213. W ARRINGTON , G., I VIMEY -C OOK , H. C. & C OPE , J. C. W. 1994. St Audrie’s Bay, Somerset, England: a candidate Global Stratotype Section and Point for the base of the Jurassic System. Geological Magazine, 131(2), 191– 200. W ARRINGTON , G., C OPE , J. C. W. & I VIMEY -C OOK , H. C. 2008. The St Audrie’s Bay–Doniford Bay section, Somerset, England: updated proposal for a candidate Global Stratotype Section and Point for the

base of the Hettangian Stage, and of the Jurassic System. International Subcommission on Jurassic Stratigraphy Newsletter, 35(1), 2 –66. W HITESIDE , J. H., O LSEN , P. E., K ENT , D. V., F OWELL , S. J. & E T -T OUHAMI , M. 2007. Synchrony between the Central Atlantic magmatic province and the Triassic– Jurassic mass-extinction event? Palaeogeography, Palaeoclimatology, Palaeoecology, 244, 345–367. W HITESIDE , J. H., O LSEN , P. E., K ENT , D. V., F OWELL , S. J. & E T -T OUHAMI , M. 2008. Synchrony between the Central Atlantic magmatic province and the Triassic–Jurassic mass-extinction event? Reply to comment of Marzoli et al. 2008. Palaeogeography, Palaeoclimatology, Palaeoecology, 262, 194– 198. W ILLIFORD , K. H., W ARD , P. D., G ARRISON , G. H. & B UICK , R. 2007. An extended organic carbon-isotope record across the Triassic– Jurassic boundary in the Queen Charlotte Islands, British Columbia, Canada. Palaeogeography, Palaeoclimatology, Palaeoecology, 244, 290 –296. Y AROSHENKO , O. P. 2007. Late Triassic palynological flora from Western Ciscaucasia. Paleontological Journal, 41(11), 1190–1197. Y RIGOYEN , M. R. & S TOVER , Y. L. E. 1969. La palinologı´a como elemento de correlacio´n del Tria´sico en la Cuenca Cuyana. Actas IV Jornadas Geolo´gicas Argentinas, 3, 427–447. Z AVATTIERI , A. M. & B ATTEN , D. J. 1996. Miospores from Argentinian Triassic deposits and their potential for intercontinental correlation. In: J ANSONIUS , J. & M C G REGOR , D. C. (eds) Palynology: Principles and Applications. American Association of Stratigraphic Palynologists Foundation, 2nd edition, 2, 767–778. Z ERFASS , H., C HEMALE , F. J R , S CHULTZ , C. L. & L AVINA , E. L. 2004. Tectonics and sedimentation in southern South America during Triassic. Sedimentary Geology, 166, 265–292. Z IEGLER , A. M., S COTESE , C. R. & B ARRETT , S. F. 1983. Mesozoic and Cenozoic paleogeographic maps. In: B ROESCHE , P. & S U¨ NDERMANN , J. (eds) Tidal Friction and The Earth’s Rotation II., Springer-Verlag, Berlin, 240–252.

The biostratigraphic importance of conchostracans in the continental Triassic of the northern hemisphere HEINZ W. KOZUR1* & ROBERT E. WEEMS2 1

Re´zsu¨ u. 83, H-1029 Budapest, Hungary

2

926A National Center, U.S. Geological Survey, Reston VA 20192, USA *Corresponding author (e-mail: [email protected])

Abstract: Conchostracans or clam shrimp (order Conchostraca Sars) are arthropods with a carapace consisting of two chitinous lateral valves. Triassic conchostracans range in size from 2 to 12.5 mm long and are common in deposits that formed in fresh water lakes, isolated ponds and brackish areas. Their dessication- and freeze-resistant eggs can be dispersed by wind over long distances. Therefore many conchostracan species are distributed throughout the entire northern hemisphere. In the Late Permian to Middle Triassic interval, several of these forms are also found in Gondwana. Many wide-ranging conchostracan species have short stratigraphic ranges, making them excellent guide forms for subdivision of Triassic time and for long-range correlations. The stratigraphic resolution that can be achieved with conchostracan zones is often as high as for ammonoid and conodont zones found in pelagic marine deposits. This makes conchostracans the most useful group available for biostratigraphic subdivision and correlation in continental lake deposits. Upper Triassic Gondwanan conchostracan faunas are different from conchostracan faunas of the northern hemisphere. In the Norian, some slight provincialism can be observed even within the northern hemisphere. For example, the Sevatian Redondestheria seems to be restricted to North America and Acadiestheriella n. gen. so far has been found only in the Sevatian deposits from the Fundy Basin of southeastern Canada. Here we establish a conchostracan zonation for the Changhsingian (Late Permian) to Hettangian (Early Jurassic) of the northern hemisphere that, for the most part, is very well correlated with the marine scale. This zonation is especially robust for the Changhsingian to early Anisian, late Ladinian to Cordevolian and Rhaetian to Hettangian intervals. For most of the Middle and Upper Triassic, this zonation is still preliminary. Five new genera, six new species and a new subspecies of conchostracans are described that are stratigraphically important.

Half of the eight stage boundaries of the Triassic have been defined by a bio-event within a marine Global Stratotype and Point (GSSP) locality, and these definitions have been accepted by both the International Subcommission on Triassic Stratigraphy and the International Commission on Stratigraphy. The remaining four stage boundaries are nearing final definition. In the Lower Triassic, both the base of the Induan (priority: Brahmanian) Stage (¼ base of Triassic) and the base of the next younger Olenekian Stage have been firmly defined. In the Middle Triassic, there is wide agreement that the defining species for the base of the Anisian Stage should be Chiosella timorensis in the GSSP candidate site at Desli Caira (Romania), but there has not yet been a formal vote on this. The base of the overlying Ladinian Stage, however, has been firmly defined. In the Upper Triassic, the base of the Carnian has been firmly likewise defined, but there is not yet a final definition for the boundaries of the overlying Norian and Rhaetian stages. A consensus has not been reached on a defining species for the base of the

Norian or its GSSP locality, but all of the different proposals under consideration do at least fall within a rather narrow stratigraphic interval. For the base of the Rhaetian, Misikella posthernsteini Kozur & Mock has been chosen as the defining species by the International Working Group on the Rhaetian stage, and the GSSP candidate locality at Steinbergkogel (Austria) has been studied in detail by a group under the leadership of L. Krystyn (Vienna) and presented to the participants of the International Conference on ‘Upper Triassic Subdivisions, Zonations and Events’ in Bad Goisern in the autumn of 2008. The base of the overlying Hettangian stage (¼ base of the Jurassic) has been defined (so far only by a working group) as the FAD (First Appearance Datum) of Psiloceras spelae Guex, Taylor, Rakus & Bucher. The final definition of the Triassic stages within marine GSSP sections will be completed in the near future, but more than 50% of known Triassic rocks are of continental origin. Therefore, the main task of Triassic stratigraphers in the future will be subdividing and correlating terrestrial strata, both between

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 315– 417. DOI: 10.1144/SP334.13 0305-8719/10/$15.00 # The Geological Society of London 2010.

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continents and to the international scale as defined in marine, mainly pelagic successions. Currently, lithostratigraphic subdivisions and correlations are used primarily in Triassic continental basins, and this generally works well within a basin that has numerous well-exposed outcrops in the basin center. These lithostratigraphic units, however, are difficult to correlate on a global scale with other continental successions, and it is still more difficult to correlate them with marine successions. Correlations have been attempted between these deposits by use of different fossil groups, by magnetostratigraphy, and by stratigraphic evaluation of carbon isotope trends. Among fossil groups, the most suitable are sporomorphs, conchostracans, terrestrial vertebrates and, to a minor degree, also megaplants, charophytes, ostracods, fresh water bivalves, fish remains, tetrapod footprints and arthropod trails. Conchostracans have the highest biostratigraphic resolution power of all Triassic continental fossils, and in some intervals their resolution is as high as that of ammonoids and conodonts in pelagic marine beds. They have the widest distribution of identical species in both low and high latitudes, have the lowest level of endemism, and the lowest correspondence to palaeolatitudes (palaeoprovinciality) of all investigated Triassic continental fossils. Additionally, they can be found in a variety of facies including grey, black, green and red beds. Their potential for correlation with marine facies is high because they are not only common in fresh water deposits, but they also can be found in brackish deposits, deltaic marginal marine beds and on some bedding planes or in brackish intervals of very shallow marine deposits such as the Werfen facies in the Balaton Highland of west –central Hungary. In the present paper, a conchostracan zonation for the Triassic northern hemisphere is presented. In the Lower Triassic (and partly also in the Middle Triassic) many of the northern hemisphere guide forms also occur in parts of Gondwana (India, central, eastern, western and northern Africa, and South America). In the Upper Triassic, however, the Gondwanan and northern hemisphere conchostracan faunas are different.

General character, preservation and palaeoecology of fossil conchostracans Conchostracans or clam shrimp (order Conchostraca Sars) are Arthropoda of the Class Branchiopoda Latreille. They have a short, laterallycompressed body enclosed in a carapace consisting of two lateral valves. As the animal grows, it enlarges its carapace by adding bands of new shell material (called growth bands) that can be distinguished from older growth bands by a narrow line

(growth line) that marks the boundary between adjacent growth bands. The carapace typically has a chitinous composition and can range from 2–42 mm long, though among Triassic forms the range is 2– 12.5 mm and usually 3 –10 mm. The frequent preservation of the carapace as phosphatized shells indicates that there is, at least in some fossil groups, a low calcium phosphate content as well. Well-preserved fossil shells show a structure of several layers (Kozur 1982). In the suborder Spinicaudata Linder, 1945, to which all Triassic conchostracans belong (see the next section), the shells have a very small to large umbonal area without growth lines and a generally (much) larger part with growth lines. Vertical radial ribs may be present that may bear nodes or short spines at the intersection points with the growth lines. The umbonal area is smooth or bears one, or rarely two or three, often elongate nodes. Exceptionally, the umbonal node is transformed into a long, hollow spine. Additional spines may be present along the dorsal and posterodorsal margins. The space between the growth lines is smooth, punctate, reticulate or has radial lirae or anastomozing lirae. Usually only the carapace is preserved, but exceptionally even the body can be discerned. Often only prints of the valves are preserved, but these can be identified just as well as preserved valves. The preserved chitinous shells and their prints are most often strongly deformed in finegrained sediments (claystones, sometimes in limestones). Such forms with shell deformation were often regarded as distinct species or even genera, but they are only variants of undeformed species as demonstrated by Kozur (1983), Kozur & Seidel (1983a) and Goretzki (2003) for Lower Triassic conchostracans. For this reason, the exact stratigraphic range and regional distribution of fossil conchostracans requires modern studies of continuous conchostracan-bearing sections with different modes of preservation, and older publications must be evaluated with caution. Phosphatized carapaces are often undeformed. In limestones, conchostracan shells usually are also undeformed. In somewhat coarser-grained sediments (siltstones, fine-grained sandstones), the conchostracans are mostly undeformed, but finer details of the sculpture and above all the microsculpture in the space between the growth lines are not preserved. The best approach, where possible, is to study conchostracans both from claystones, shales or micritic limestones (for the microsculpture) and from siltstones or fine-grained sandstones for the undeformed outline of the carapace. Fresh-water limestones with conchostracans are rather rare, but they often contain the best-preserved fossil conchostracans (e.g. at Dalongkou in the conchostracan-rich ‘Sesame cake beds’).

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The main habitats of living and fossil conchostracans are temporary, alkaline inland ponds and small temporary fresh-water lakes. They also can occur, however, in flood-plain pools, coastal floodplains, coastal salt flats, and (in the case of some species) in brackish water estuarine facies or in deltaic plains with variable salt content (Webb 1979). Conchostracan eggs can withstand long desiccation and, in many species, also tolerate long periods of freezing. These dessication- and freeze-resistant eggs can be dispersed by wind and water over long distances. The life cycle from hatching to sexual maturity is very short, lasting only 5–23 days in modern species (Webb 1979). Therefore, conchostracans can occupy, even in semiarid and arid areas, small temporary lakes and ponds that exist for only a short time and are not necessarily even present every year.

Classification of conchostracans A robust classification of fossil and even living conchostracans (clam shrimp) has not yet been established. The order Conchostraca Sars, 1867 is often regarded as paraphyletic or polyphyletic and has been replaced by some workers with the orders or suborders Spinicaudata Linder, 1945 and Laevicaudata Linder, 1945, so the phyletic status of both the Conchostraca and the Spinicaudata (to which most of the living Conchostraca and nearly all fossil Conchostraca belong) remains uncertain. Sometimes the Cyclestherida Sars, 1899 are separated from the Spinicaudata as an independent order or suborder, and the Spinicaudata, Laevicaudata and Cyclestheriida are partly regarded as orders or suborders of the superorder Diplostraca Gerstaecker, 1866; this classification completely eliminates the order Conchostraca. In the present paper, we regard the superorder Diplostraca as containing the orders Conchostraca Sars, 1867 and Cladocera Latreille, 1829 (the latter commonly called ‘water fleas’). According to Olesen (1998), the monophyletic character of the Diplostraca is generally accepted. Therefore the monophyletic character of the Conchostraca only needs to be discussed briefly. The fossil record does not support a polyphyletic origin for the Conchostraca. Palaeozoic Conchostraca all belong to the Spinicaudata, and they do not show any indication of convergence toward different other branchiopod Crustacea, but rather are a fairly uniform group that is morphologically very similar to extant Spinicaudata. Even preserved soft parts, known only from a few fossil taxa, show no clear differences from living Spinicaudata. In the Permian and Triassic, all morphological transitions of the carapace outline can be observed between

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forerunners of Cyzicidae among the Spinicaudata to Cyclestheriidae Sars, 1899. They are mostly thick-shelled genera that contain both species with a circular and an oval carapace outline, such as Magniestheria truempyi (Kozur & Seidel) with a circular shell outline in females and Magniestheria mangaliensis (Jones) with an oval outline. It is not until the Norian that two species appear in the Germanic Basin and in the Culpeper Basin of Virginia (USA) which are thin-shelled forms with a circular carapace outline. These species, representing an as yet undescribed genus, are the oldest examples of the Cyclestheriidae. In their shell morphology, the Cyclestheriidae have never diverged far from other Spinicaudata and should be regarded as a family of this suborder. The Mesozoic Conchostraca are mainly Spinicaudata, though a few members of the Cyclestheriidae occur among them. The oldest species referable to the Laevicaudata have been reported from the Cretaceous (Tasch 1969), but it is only in the Eocene that typical fossil forms of the Laevicaudata appear (Shen et al. 2006). Thus, the fossil record supports the idea that the Conchostraca are a monophyletic group, and nearly all known Palaeozoic and Mesozoic Conchostraca are referable to the suborder Spinicaudata. The Cyclestheridae seemingly evolved during the Late Triassic from the Spinicaudata. The position of the Laevicaudata cannot be evaluated from the fossil record, because they have not been found among Palaeozoic and preCretaceous Mesozoic Conchostraca, and even in Cretaceous and younger beds they are rare. For this reason, we are concerned here only with the order Conchostraca and the suborder Spinicaudata. Because soft parts of conchostracans are found only rarely, the taxonomy of fossil conchostracans is entirely based on features of the carapace. Size, outline, convexity of carapace, number and width of growth bands, position and size of the free umbonal area (without growth lines), sculpture, and microsculpture are all important for establishing the taxonomy of fossil conchostracans. In combination with other features, the maximum size of adult conchostracans sometimes can be used to separate taxa that belong to small and large Triassic species. This is possible because the smallest Triassic species are only 2–3 mm long, whereas the largest species are 8–12.5 mm long. In several Triassic lineages, a clear increase in size can be observed over time, often even within a single species lineage. Thus, the latest Brahmanian (latest Induan) to basal Olenekian species Magniestheria subcircularis is only 3 –5.2 mm long (mostly around 3.5 mm), but the succeeding lower Smithian M. truempyi Kozur & Seidel can be up to 10 mm (5.2– 10 mm) long. The late Spathian to Aegean Euestheria albertii mahlerselli Kozur & Lepper

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is only 2.4– 3.3 mm long, but the succeeding subspecies Euestheria albertii albertii (Voltz) from the basal Bithynian is larger (around 4 mm long) and becomes still larger in the lower Bithynian Gre`s a` Voltzia where it can be up to 5 mm long. Finally, in the upper Bithynian Holbrook Member of the Moenkopi Formation, it can be up to 6–6.3 mm long (Kozur, Lucas & Morales, in prep.). Similar phylogenetic increases of size can be observed in the Upper Triassic genus Shipingia, where the largest S. olseni Kozur & Weems (up to 12.5 mm) occur immediately before its extinction at or near the end of the late Sevatian. The outline of the carapace (including its length/ height [l/h] ratio) is always an important taxonomic characteristic at lower taxonomic levels, but species with the same outline can belong to different lineages. Particularly important is the presence or absence of a straight dorsal margin, with or without distinct antero-dorsal and postero-dorsal corners. The height, degree and shape of rounding along the anterior and posterior margins are also important taxonomic characteristics. The l/h ratio may be different in females and males of the same species. Some special features of the outline, such as an incision in the upper part of the anterior margin, surrounded by an elevation, may be restricted to only one genus, for example, the Sevatian genus Redondestheria Kozur, Weems & Lucas (in Kozur & Weems 2005), but other similar features may be taxonomically important yet present in different lineages (e.g. a slight concave incision immediately below the straight dorsal margin in the upper part of the posterior margin, which in the Triassic is known in Falsisca Novozhilov, Dictyonatella Kozur, some Vertexia Ljutkevitch, and in Eosolimnadiopsis gallegoi Kozur). The outline of conchostracan shells can be strongly altered by plastic deformation, and this is common in claystones. Such deformation leads all too often to the creation of numerous synonymous taxa that simply represent different preservation of a single species. The convexity of the carapace is sometimes important for supraspecific taxonomy, but generally it is not important for separation of species within the same genus. This characteristic can be accurately evaluated only in undeformed specimens that occur in somewhat coarser sediments such as siltstones, fine sandstones, or occasionally limestones. With some experience, carapace convexity can be reconstructed from flattened specimens. In species with strong carapace convexity, the umbonal area generally overreaches the dorsal margin, especially if the specimens are flattened. The number and width of growth bands in some species are rather constant and characteristic for that species. In other species, they are highly variable

and represent only intraspecific variation. For example, the late Ladinian to early Carnian Euestheria minuta (von Zieten) consistently has 10 –20 growth lines and thereby can be readily distinguished from the somewhat younger early Carnian E. winterpockensis (Bock), which has 19 –46 growth lines. E. winterpockensis evolved from E. minuta by greatly increasing the number of growth lines (Kozur & Weems 2007), and this can be documented by the presence of transitional forms where there is a slight stratigraphic overlap in the occurrence of these two forms. The reverse trend is seen within the genus Laxitextella, where the stratigraphically oldest species L. multireticulata (Reible) has numerous and narrow growth lines, but the stratigraphically youngest species L. freybergi Kelber & Kozur has only a few very broad growth bands. In contrast, specimens of Magniestheria mangaliensis (Jones) have anywhere from 14 –36 growth lines, yet there is no trend over time toward either an increase or decrease in the number of growth lines which can be used to distinguish different taxa. Radial ribs or radially-arrayed lines of nodes also are important taxonomic features. Their presence or absence is a family level characteristic, but the number of radial ribs or the number of radial lines of nodes may constitute a generic character, a specific character or (in certain cases) only an intraspecific feature. Nodes generally develop where radial ribs intersect with growth lines. During evolution within a lineage, however, the radial ribs may disappear between the nodes leaving only the radial lines of nodes visible on the growth lines. In some advanced forms, the nodes may shift their position to the middle of the growth bands. The length of radial ribs or radially arranged lines of nodes are important for separation of species, but small variations in this trait are only an intraspecific feature. Whether radial ribs are obvious or rather subtle is an intraspecific feature. Generally, the stronger the node elevations are at the growth lines, the weaker the radial ribs are and vice versa. Presence or absence of spines at the dorsal margin and/or at the postero-dorsal corner of the carapace is an important characteristic for separating genera within the Vertexiidae, but these fragile features often are not preserved. Only the posterodorsal spine is robust enough to be usually preserved. The length of this posterodorsal spine can have taxonomic importance. Position, size and sculpture of the free umbonal area (which lacks growth lines) is a very important feature for separating species, genera and even families. However, the entire ontogenetic growth series within a species must be considered. The small juvenile stages have a relatively larger free umbonal area than adults of the same species in

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relation to the length and height of the carapace, because during growth more and more growth bands are added after the first few have formed around the free umbonal area. Simple nodes, elongated nodes, ribs or spines occur on the free umbonal area in many species, and these are also taxonomically important. These features mostly represent generic or familial characters, but in some genera their form and size distinguish different species. Hollow spines are only rarely preserved, because they stand vertically or obliquely on the free umbonal area. Therefore they only can be preserved as spines if they are squeezed onto the plane of the carapace, either in the umbonal area or beyond its dorsal margin. This means that they are much more frequently preserved on strongly deformed specimens than they are on undeformed specimens. Normally only the sediment fill within the basal part of a spine is preserved, which looks like a cone or conical node (Kozur 1983). This has resulted in different genera being established for different preservations of umbonal spines (Kozur & Seidel 1983a). Another important taxonomic feature is microsculpture, if present. Most Lower Triassic conchostracans lack or have only very indistinct microsculpture, though as a group they include the widest array of species with distinct macrosculpture (e.g. umbonal spines, spines on the dorsal margin, radial ribs). In the Middle Triassic, several forms appear that have a distinct microsculpture, but species without microsculpture still are common. By the Upper Triassic, the number of taxa with distinct microsculpture exceeds the number of species without microsculpture or with only very weak microsculpture, and the differences in microsculpture can be used to separate genera and species. Most common is a reticulation pattern on the external surface of the carapace. The polygons may be small to large and are a useful feature for distinguishing species. The arrangement of polygons also can be different, being arrayed in vertical or oblique lines (e.g. Shipingia hebaozhaiensis Shen) or lying parallel to the growth lines (e.g. S. olseni Kozur & Weems). Again, this is a useful feature to separate species. Often the microsculpture is arranged in radial lirae, a feature which also can be used to distinguish genera and species. These densely spaced radial lirae may occur in straight and parallel, irregularly curved, or anastomozing patterns, and such differences also can be used to separate species. Different types of microsculpture can be combined in a single species, for example, radial lirae may be developed on marginal growth bands, but reticulation is present on the inner growth bands and on the free umbonal area. Other types of microsculpture are a pitted surface or tiny, densely spaced nodes on the growth lines. These often represent the base of spines or hairs

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on the growth lines (growth line setae) that usually are not preserved (e.g. Shen 2003; Shen & Zhu 1990). All of these different types of microsculpture, including combinations of them, can be used to separate species or sometimes genera.

Importance of conchostracans for the biostratigraphic subdivision and correlation of the continental Triassic In the Triassic, conchostracans are very common in sediments that formed in terrestrial and brackish environments. In brackish intercalations of marine beds (e.g. in the Werfen Beds of the Tethys and brackish intercalations of ammonoid-bearing beds in northeastern Siberia) these occurrences can be correlated accurately with the marine timescale by identification of associated conodonts, bivalves and sometimes ammonoids. Radially ribbed conchostracans (e.g. Estheriella, Lioleaiina) and spined conchostracans (e.g. Cornia, Molinestheria and Vertexia) are common during the Early Triassic. These are distinctive, short-ranging excellent guide forms that have a very widespread northern hemisphere-Gondwana distribution, especially during the Gandarian (junior synonym ¼ Dienerian). The Gandarian was established by Mojsisovics et al. (1895) as the upper substage of the Brahmanian. It was well defined by the ammonoid faunas of the Lower Ceratite Limestone and Ceratite Marls of the Salt Range, Pakistan. Seventy years later Tozer (1965) established the Dienerian Stage for the same stratigraphic interval, and nobody has ever doubted that both units comprise the same time interval. We prefer to use the Gandarian Substage, partly for reasons of priority but also because the Dienerian was established in a lower diversity high latitude environment while the Gandarian was established in a low latitude environment that was much more favourable for defining an international stage or substage boundary. The conchostracan zonation that has been established for Early Triassic continental deposits (Kozur & Seidel 1983a, b; Kozur 1993b, 1999a, b; Kozur & Weems 2007) is just as detailed as the ammonoid and conodont zonations established in pelagic marine deposits. During the Upper Triassic, forms with distinct microsculpture became dominant, and a number of these are also short-ranging species with very wide regional distribution. Conchostracans are quite possibly the best available fossil group for establishing detailed Triassic continental biostratigraphy that is cross-correlative with the standard marine timescale. This assertion is based on: † The general abundance of conchostracan specimens in many continental to marginal marine settings.

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† The high percentage of specimens that represent short-ranging species. † The fact that conchostracans produce desiccation- and freeze-resistant eggs that are spread by wind, which allows them to occupy even isolated and temporary bodies of fresh to brackish water over broad geographical areas. † The resultant widespretad occurrence of shortranging species over nearly all of the northern hemisphere and also over parts of Gondwana. † The occurrence of some species in brackishestuarine and deltaic environments, or in brackish intercalations between marine environments that contain conodonts and other marine guide forms, allowing cross-correlation with the marine timescale. † The presence of many species of conchostracans in markedly different climatic zones, ranging from high latitudes to low latitudes. The wind-transport of desiccation-resistant conchostracan eggs, in combination with their short life-cycle between hatching from eggs until sexual maturity, explains the very wide, transcontinental distribution of many fossil conchostracan species. Thus, the easily recognizable, short-ranging spined Cornia germari (Beyrich) can be found in the upper Gandarian (¼ upper Dienerian) of Greenland, the Germanic Basin, the Russian Platform, northern Urals, Pricaspian depression, Timan (northern Russia), Jakutia (northeastern Siberia), Sinkiang (Xinjiang, China), India, Gabon (Africa) and Australia. Additionally, it occurs in Hungary in brackish intercalations between conodont-bearing marine strata of the Werfen Beds (Kozur & Mock 1993), in Greenland and Jakutia (Russia) in brackish beds and brackish intercalations between marine beds, and in Jakutia with ammonoids. Other shortranging species, for example, the radially ribbed Estheriella costata Weiss and E. nodosocostata (Giebel), have a regional distribution ranging from the Germanic Basin, across the Russian Platform and Gondwanan India to Angola. Magniestheria mangaliensis (Jones) occurs throughout the Lower Triassic of Eurasia (including brackish intercalations between marine strata of the Werfen Beds of the Tethys; Kozur & Mock 1993) and also in Gondwanan India and Angola. These few examples serve to show that, among Triassic Conchostraca, there are a number of short-ranging guide-forms that occur both in the northern hemisphere and in Gondwana. Endemism is not a serious problem for establishing a comprehensive latest Permian through early Hettangian conchostracan zonation scheme. As conchostracans have drought and freeze-resistant eggs that are distributed readily by wind, widely scattered ponds in an arid area may have the same conchostracan fauna. Generally, conchostracans have a very

large regional distribution. The Late Permian low and high northern latitudes have a very similar fauna, and individual species of the family Leaiidae have been reported both from very high northern palaeolatitudes of northern Siberia and from high southern palaeolatitudes of southeastern Australia. The same species of Latest Permian and Early Triassic conchostracans often occur in Greenland, the Germanic Basin, the Russian Platform, Siberia, China, Gondwanan India and central and eastern Africa. Up to the late Ladinian, guide forms of the northern latitudes also have been found in parts of Gondwana. For example, Euestheria minuta (von Zieten) is present throughout the entire northern hemisphere and in Argentina as well. This very wide distribution of many species of conchostracans in both in the northern hemisphere and across large parts of Gondwana ended during the Carnian. Upper Triassic conchostracans of Argentina, Brazil and Chile are very different from the conchostracan faunas of the northern hemisphere, and only in northwestern Africa do the same conchostracans occur as in the northern hemisphere. As this part of Gondwana in the Late Triassic was directly adjacent to the eastern United States and Canada, and also close to southwestern Europe, it is not surprising that its conchostracan fauna was much the same as that of the northern hemisphere. Within the northern hemisphere, many genera and species of Upper Triassic conchostracans are widely distributed. For example, the middle Norian species S. hebaozhaiensis can be found from China through the Germanic Basin, northwestern Africa (Morocco), and the eastern United States all the way to the southwestern United States. By the Tuvalian and Norian, some endemism within the northern hemisphere can be observed. The conchostracan faunas of the southern Newark Supergroup basins and of the southwestern United States, which were at similar equatorial palaeolatitudes (0–58N), are much more diverse, with dominant species that either do not occur farther north or occur there only very rarely. The biozonal guide forms of the northern Newark Supergroup basins and the Germanic Basin are present in these equatorial faunas, but often they also are rather rare. Thus, with adequate collecting, there are no serious problems with correlating the conchostracan faunas of these two areas, even though the overall complexion of their faunas is distinctly different. By the late Norian, endemism becomes even more pronounced. For example, the genus Redondestheria Kozur, Weems & Lucas, which is common both in the eastern and southwestern United States, is absent in the Germanic Basin and China. Similarly, Acadiestheriella cameroni n. gen. n. sp. is restricted to the upper Norian of the Fundy Basin, and there to a single level.

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However, there is a forerunner of this form in the Middle Triassic of the Germanic Basin, which is known from one specimen from the lower Anisian (Gall 1971) and one specimen from the upper Ladinian (Warth 1969). Thus, it is quite possible that A. cameroni has a much wider distribution than currently known, because it may be an extraordinarily rare form that had a very exceptional bloom within a single upper Norian horizon of the Fundy Basin. With the extinction of all known large species of conchostracans at the end of the Norian, the surviving small species again show a nearly ubiquitous distribution throughout the northern hemisphere.

Previous work North America and Europe Many groups of Triassic microfossils, today considered to be of exceptional stratigraphic importance, were in the 19th century either unknown (e.g. conodonts, holothurian sclerites, scolecodonts, charophytes, sporomorphs) or poorly known (e.g. foraminifers, ostracods, radiolarians). At most, somewhere between 1– 5% of the presently known species of these groups were described in the 19th century. Triassic conchostracans, with a shell size range of 2–12.5 mm, lay at the boundary between microfossils and macrofossils, and in the second half of the 19th century they were relatively well known compared with the true microfossils of the Triassic. Depending on the stage, about 10 –20% of the presently known and valid Triassic conchostracan species of Europe and North America were described in the 19th century. Investigation of Triassic conchostracans began in the 19th century in the Germanic Basin of Germany (von Zieten 1833; Beyrich 1857; Giebel 1857; Jones 1862, 1863; Sandberger 1871; Weiss 1875), in France (Voltz 1835), in Great Britain (Jones 1862, 1863, 1890, 1891; Jones & Woodward 1894), and in the Newark Supergroup of the eastern United States (Emmons 1856, 1857, 1863; Lea 1856; Jones 1890, 1891). After this initial flurry of research, a long interval followed in all of these regions during which only a few new Triassic conchostracan species were described. These rare contributions include investigations by Picard (1911); Schmidt (1928); Raymond (1946); Defretin (1950), and Defretin-Lefranc (1963) for the Germanic Basin. In the Germanic Basin, modern work on Triassic conchostracans began with Reible (1959, 1962) and Warth (1969). Since that time, a number of modern papers concerning Triassic conchostracans and Triassic conchostracan stratigraphy in the Germanic Basin have been published (e.g. Alexandrowicz & Slupczynski 1971; Battarel &

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Gue´rin-Franiatte 1971; Gall 1971, 1976; Kozur 1974, 1979, 1980, 1982, 1983, 1989, 1993a, b; Kozur & Seidel 1981, 1983a, b, c; Martens 1983, 2005; Kozur et al. 1993; Olempska 2004; Kozur & Weems 2006, 2007; Seegis 1997; Ptaszyn´ski & Niedz´wiedzki 2004, 2005, 2006a, b; Korte et al. 2007; Kozur & Hauschke 2008). The taxonomy of Lower Triassic and lower Anisian conchostracans has been revised and an uppermost Permian to lower Anisian conchostracan biostratigraphy established (Dockter et al. 1980; Kozur & Seidel 1983a, b, c; Mader 1984; Kozur 1993b, 1999a, b; Kozur et al. 1993; Ptaszyn´ski & Niedz´wiedzki 2004, 2005, 2006a, b; Kozur & Weems 2006, 2007; Kozur & Hauschke 2008) that has been correlated with palaeomagnetic succession, carbon isotope excursions, and Milankovitch cyclicity (Bachmann & Kozur 2004; Korte & Kozur 2005; Kozur & Weems 2006, 2007; Korte et al. 2007; Kozur & Hauschke 2008). In contrast, little work has been done on the taxonomy of Upper Triassic conchostracans of the Germanic Basin since Reible (1962) and Warth (1969). From the late 1960s until 2006, only two new genera, two new species, and one new subspecies were established from the Carnian of the Germanic Basin (Kozur 1982; Geyer 1987; Olempska 2004). Otherwise, Upper Triassic conchostracans only have been mentioned or illustrated (e.g. Hopf & Martens 1992; Reimann & Schmidt-Kaler 2002), or else described in open nomenclature (Seegis 1997). Recently, Kozur & Weems (2007; Appendix 1) described three new index species from the Tuvalian of the Germanic Basin. In the Arctic region, conchostracans have been reported from the Gandarian (Dienerian) of eastern Greenland (Defretin-Lefranc et al. 1969; Kozur & Seidel 1983a, b), where Cornia germari occurs. The other classic region for early conchostracan research in the 19th century was the Upper Triassic to Lower Jurassic Newark Supergroup rift basins in eastern North America. Following the paper on this region by Jones (1891), however, only a few conchostracans were studied or described (Wanner 1926; Bock 1946; Raymond 1946). The first modern description of Upper Triassic conchostracans from the Newark Supergroup was by Bock (1953a), who shortly thereafter also published a brief taxonomic revision that replaced the generic name Howellites Bock, 1953 with Howellisaura Bock, 1953 because of homonymy (Bock 1953b). Other authors since 1953 occasionally have noted occurrences of conchostracans in the Newark Supergroup, but they either did not attempt any taxonomic designations or else used ‘wastebasket’ names (such as Cyzicus sp.) to describe them. In a few cases small photographs were shown, with insufficient magnification for accurate taxonomic determination, or else

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drawings were presented (e.g. Gore 1986; Olsen 1988). After Bock’s work, the detailed study of Newark Supergroup conchostracans and their taxonomy languished until the work of Kozur and Weems (2005, 2006, 2007). Kozur & Weems (2005) described the upper Alaunian to lower Hettangian conchostracans of the Culpeper and Newark basins and established the first conchostracan zonation for that time interval. Kozur & Weems (2007) soon after presented an early Carnian to early Hettangian conchostracan zonation for the entirety of the Newark Supergroup and the Germanic Basin, the latter well correlated with the marine scale. The Germanic Basin and the Newark Supergroup of eastern North America have many Upper Triassic conchostracan species in common that occur in the same order of succession, but a great many other species in the Newark are undescribed or in need of major revision. Most of the new species occur in the southernmost rift basins (e.g. in the Durham and Sanford sub-basins of the Deep River Basin and in the Dan River– Danville Basin) which had palaeolatitudes around 0–58N. This equatorial fauna is more diverse than the conchostracan fauna found in the northern rift basins (e.g. in the Fundy, Hartford, and Newark basins) which had palaeolatitudes around 10–158N. Even in these more northerly basins, several new species are yet to be described. This will be done mainly in forthcoming papers, and only a few new taxa of exceptional stratigraphic importance will be described in the present paper. Rich conchostracan faunas have been found in the southwestern United States (Texas, New Mexico, Arizona, and Utah). Spathian and lower Anisian conchostracans from Arizona will be described by Kozur, Lucas & Morales (in prep.). They are identical with the conchostracans of this age from the Germanic Basin. In contrast, very few of the Late Triassic conchostracans from this region have been described. These include Anyuanestheria wingatella (Tasch) from the Adamanian (late early Tuvalian to middle Tuvalian) ‘Lake Ciniza’ facies of the Bluewater Creek Formation, New Mexico (Tasch 1978), and Redondestheria novomexicoensis Kozur, Weems & Lucas and Shipingia olseni Kozur & Weems, both from the uppermost Sevatian part of the Redonda Formation, New Mexico (Kozur & Weems 2005, 2007).

Asia Unlike in Middle and Western Europe and the eastern United States, where few studies were conducted on conchostracans during the first two-thirds of the twentieth century, comprehensive work was being done then on Triassic conchostracans in the Russian part of the former Soviet Union on the

Russian Platform, in the Pricaspian Basin, and in Mangyshlak and Siberia. There, mainly Lower Triassic and lower Anisian conchostracans were described, which generally were the same as those from the Germanic Buntsandstein (e.g. Chernyshev 1934; Lyutkevich 1938; Novozhilov 1946, 1958, 1959, 1960, 1966, 1970, 1976; Novojilov [Novozhilov] 1958a, b, c, d; Novozhilov & Kapelka 1960, 1968; Zaspelova 1961, 1965, 1973; Molin 1965a, b, 1966, 1968, 1975; Molin & Novozhilov 1965; Defretin-Lefranc 1965; Budanov & Molin 1966; Blom 1969, 1974; Lopato 1972; Menner & Lipatova 1972; Kozur et al. 1983; Lipatova & Lopato 1985; Tuzhikova 1985; Sadovnikov & Orlova 1990, 1993, 1994; Sadovnikov 1997, 2008; Orlova 1999). In the Lower Triassic and lower Anisian, nearly the same fauna is present in the Germanic Basin and in Russia. Despite this, three to five times as many species were described from Russia. This is strictly a taxonomic issue, however, because for most species several synonymous species were erected that were based on different modes of preservation or on deformed specimens (Kozur 1982; Goretzki 2003). In Russia, 10– 15, or even more, conchostracan species were often described from a single bedding plane, sometimes even from a single bedding plane within the small area of a borehole, and these on occasion were placed into more than 10 genera. It is highly unlikely that this many species of conchostracans would ever occur at a single horizon, for this is far more than have been demonstrated from a single locality anywhere else in the world. For example, in modern populations of one pond or temporary lake, or even on single bedding planes of Triassic lake deposits, usually one, two or at most three species can be observed (Kozur 1983). Even species with very strong and characteristic sculpture were assigned to numerous different species and genera. For example Cornia germari (Beyrich), with its large, hollow umbonal spine, was assigned to 11 different genera and 21 different species. An additional nine species, assigned to Cornia and one other genus, also questionably belong to Cornia germari as shown by Kozur (1982) and Kozur & Seidel (1983a). Most or all of the latter forms probably belong to this species, but the type material has not been available for study, and the original descriptions and illustrations are inadequate to confirm or reject this synonymy. Investigations of Lower Triassic conchostracans from the Russian Platform by Kozur have shown that the same species with the same ranges occur there as in the Germanic Basin and that the number of species represented is similar. The presence of numerous synonyms among the conchostracans described from Russia also was confirmed by Goretzki (2003), who presented examples where

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synonymous species, assigned to different genera, occurred on a single bedding plane of one slab. Additionally, he demonstrated that many drawings of holotypes by Novozhilov do not accurately portray the characteristics of the species being described. Even so, the above mentioned papers by Novozhilov and Molin mark an important step in the research of Triassic conchostracans, because they demonstrated for the first time that conchostracans actually are widespread and even common in continental beds that until then had been regarded as sparsely fossiliferous or barren of fossils and because they highlighted the stratigraphic importance of Middle Permian to lower Anisian conchostracans. Since these earlier studies, conchostracans also have been reported from continental and brackish water sediments between, above or below marine sediments of the Arctic, for example, in the lowermost Triassic of northeastern Siberia or in late Olenekian (Spathian) sediments from the coastal region of the Laptev Sea in northern Siberia, close to the Olenekian type area (Molin & Novozhilov 1965). Comprehensive taxonomic work on Upper Triassic and Liassic (Lower Jurassic) conchostracans has been undertaken in Asia, and in China there also has been work done on Lower and Middle Triassic conchostracans (e.g. Mansuy 1912a, b; Chi 1931; Kobayashi 1951, 1952, 1954, 1973, 1975, 1984; Chen 1974, 1982; Chang et al. 1976; Chen & Shen 1980, 1985; Liu 1982, 1987a, b, c, 1988a, b, 1989, 1990, 1994; Liu et al. 1990; Shen 1985, 2003; Li & Shen 1995). Many of these species have very widespread distributions, because they also are found in the Germanic Basin and in the Newark Supergroup.

Gondwana The study of Triassic conchostracans from Gondwana also began early with the work of Jones (1862). Today, conchostracans are known from all parts of Gondwana: southern, eastern, central and north-western Africa, Antarctica, Australia, northern India, and South America (e.g. Jones 1862, 1897; Newton 1910; Leriche 1913, 1914, 1932; Janensch 1927; Mitchell 1927; Marlie`re 1948, 1950; Teixeira 1950a, b, 1951; Defretin & Fauvelet 1951; Bock 1953a; Defretin 1953, 1958; DefretinLefrance 1967; Katoo 1971; Cockbain 1974; Ghosh & Shah 1978; Tasch & Jones 1978; Ghosh 1983; Herbst & Ferrando 1985; Tasch 1987; Ghosh et al. 1988; Gallego 1992, 1996, 1998, 1999a, b, 2001a, b, 2005; Gallego & Covacevich 1998; Shen et al. 2001, 2002; Gallego et al. 2005, 2009; Shen 2006). Tasch (1987) published a monograph about the fossil conchostracans of Gondwana, including the Triassic forms. Even in central

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Gondwana (Madagascar), where marine beds are very rare, Shen et al. (2002) recently described Magniestheria truempyi (Kozur & Seidel) from beds immediately below the second ammonoid zone of the Olenekian, thereby establishing an early Smithian age for this species. Among Gondwanan conchostracans, the best known and described are the Middle and especially the Late Triassic conchostracans from South America. In general, the correlation of Gondwanan conchostracan faunas with the international marine timescale is not very precise. In the Lower Triassic and known parts of the Middle Triassic in large parts of Gondwana (northern India, central, eastern and northern Africa, South America), conchostracan faunas are present that are very similar to those found in the northern hemisphere. In the Upper Triassic, however, the conchostracan faunas of Gondwana are rather different from those found in the northern hemisphere. Only in north-western Africa (Morocco) are the same Upper Triassic conchostracans present as in the northern hemisphere. The conchostracan faunas of Australia and Antarctica are, so far as is known, different throughout the entire Triassic from the contemporaneous faunas in the northern hemisphere, though a few taxa are probably identical at the generic level (e.g. Anyuanestheria probably ¼ Menucoestheria). Except where the same zones are present in both parts of Gondwana and the northern hemisphere, the Gondwanan conchostracan faunas will not be discussed in the present paper.

Tethys Some Triassic conchostracans have been described from the Tethys in Europe, Indochina, and from the margin of Panthalassa in Japan and Korea (Mansuy 1912a, b; Kobayashi 1951, 1952, 1954, 1973, 1975, 1984; Vada´sz 1952; Nagy 1959, 1960, 1968; Tintori et al. 1985; Kozur 1987, 1999a, b; Tintori 1990; Tintori & Brambilla 1991; Kozur & Mock 1993). These have been extremely important for correlating continental deposits with the wellestablished marine geologic timescale (Kozur 1999; Kozur & Mock 1993). They can occur in markedly different settings. They can be found in predominantly continental beds within the Tethys, such as in the Mecsek Mountains of Hungary or in Thailand. Often they occur in continental to brackish intercalations between marine beds, as in the westernmost Southern Alps, Japan and parts of Indochina. Occasionally they occur on single bedding planes or within a longer interval in very shallow water deposits that had variable salt content, such as in the Werfen Group of Hungary and in the Southern Alps. These occurrences may represent a temporary lowering of the salt content

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in the local depositional environment. Specific bedding planes with numerous conchostracans do not contain marine fossils, though marine fossils may be common in the immediately underlying and overlying beds. For example, bedding planes with Cornia germari (Beyrich) occur between immediately overlying and underlying beds that contain the marine bivalve Claraia aurita (Hauer) in the Werfen Beds of Hungary (Kozur & Mock 1993). As the Werfen Beds contain most of the conchostracan zones from the basal Triassic up to the middle Spathian, this has been a great help in correlating marine biozones with terrestrial conchostracan biozones throughout this stratigraphic interval.

Investigated regions We have studied private conchostracan collections from: the Upper Permian to basal Triassic of Dalongkou and Xialongkou (Sinkiang ¼ Xinjiang, NW China); the uppermost Permian to lowermost Hettangian of the Germanic Basin from northern, central and southern Germany, England, and southeast and southern Poland (Holy Cross Mountains, Upper Silesia); the Tethyan Triassic of the Balaton Highland (including boreholes), Mecsek Mountains (both Hungary), Werfen Beds of the Southern Alps and lower Carnian of westernmost Southern Alps (both Italy); Lower Triassic of Libya, Upper Triassic of Morocco; the Carnian to Hettangian of the Newark Supergroup basins in the eastern United States; the late Ladinian to Norian of the Fundy Basin in southeastern Canada; and the late Olenekian, early Anisian, and Carnian to Hettangian of the southwestern United States (Utah, Arizona, New Mexico, Texas); Middle and Upper Triassic of Argentina. From many of these areas we also have investigated material from museum collections, specifically the Naturkunde Museum (Stuttgart), the Institut fu¨r Geologische Wissenschaften at Martin-LutherUniversita¨t (Halle), the Yale-Peabody Museum (New Haven), the Academy of Natural Science (Philadelphia), the United States National Museum (Washington), the North Carolina Museum of Natural Sciences (Raleigh), the Texas Tech Museum (Lubbock), the New Mexico Museum of Natural History and Science (Albuquerque), and the Museum of Northern Arizona (Flagstaff).

Dalongkou (Xinjiang, NW China) During the Sino-American National Geographic Society project in Dalongkou (1996), in which H. W. Kozur took part, the stratigraphic section on the southern limb of the Dalongkou anticline was measured in detail by Dr. Spencer Lucas (New

Mexico Museum of Natural History and Science, Albuquerque), and sampled for conchostracans by Kozur. The measured section by Lucas is used here because all fossil data (not only conchostracan data) were correlated in the field with his cumulative section (recorded both in metres above the base of the Guodikeng Formation and metres below the top of the Guodikeng Formation. Additionally, both the distance between the FAD (First Appearance Datum) and LOD (Last Occurrence Datum) of important fossils and the thickness of conchostracan zones were noted. These measurements are distinctly different from an earlier Chinese measured section, but only slightly different from the measured section of Metcalfe et al. (2009), which unfortunately shows only the metres above the base or below the top of the Guodikeng Formation and not the distances between bioevents. According to the measured section of Lucas, the Guodikeng Formation is 234.7 m thick, while the section measured by Metcalfe et al. (2009) is 229.8 m, a difference of only about 2%. Using the measurements of Lucas, important data concerning the observed faunal changes and their distance from the Permian – Triassic boundary (PTB) were documented by Kozur (1998a, b), and the position of important bioevents was given in metres above the base of the Guodikeng Formation (Fig. 1). Unfortunately, all material that was obtained during the project (all palaeomagnetic samples and all conchostracan samples) was confiscated at the end of field work and, after 13 years, remains unavailable for study. All large or distinctly sculptured conchostracans (Bipemphigus, Falsisca, Megasitum, Trimpemphigus) were identified by H. W. Kozur in the field, mostly to species level by a pocket-lens, and their range was correlated with the detailed measured section made by S. G. Lucas. Smaller conchostracan genera (even those with distinct sculpture such as Tripemphigus) and genera not described at that time (e.g. species of Megasitum) could be identified in the field only to genus level or to species groups. From the F. eotriassica Zone (described below) of the uppermost Zechstein upward, the conchostracan succession of Dalongkou is identical with the succession found in the Germanic Basin. Below this level, no conchostracans are present in the upper Zechstein of the Germanic Basin, but the conchostracan succession of Dalongkou coincides at the generic level and mostly even at the species level with the conchostracan succession in the Tunguska Basin in Siberia. From this study of conchostracan ranges done during 1996 field work in Dalongkou, not only could the base of the Triassic be very precisely defined by conchostracans (Kozur 1998a, b), but it also could be shown that the base of the Triassic is not close to the FAD of Lystrosaurus, as generally

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assumed then (except by King & Jenkins 1997, who reported Lystrosaurus from the Permian of Zambia), but rather close to the LOD of Dicynodon. Later, this was independently demonstrated in the Karroo Basin in South Africa (MacLeod et al. 2000; Smith & Ward 2001; Ward et al. 2005) by locating the carbon isotope minimum at the Permian–Triassic boundary, which was found to be at the LOD of Dicynodon, and distinctly above the FAD of Lystrosaurus, which lies in the upper part of a short reversed palaeomagnetic horizon (palaeomagnetic chron ‘0r’ in Bachmann & Kozur 2004). This short reversed horizon is represented in the Germanic Basin by the lower Fulda Formation and the lowermost upper Fulda Formation of the upper Zechstein. This further confirms that the interval of co-occurrence of Lystrosaurus and Dicynodon belongs to the latest Permian as shown by Kozur (1998a, b) through conchostracan correlation in Dalongkou. In the present paper, as in Kozur (1998a, b) we place the base of the Triassic at the base of the F. verchojanica Zone (described below), located 210 m above the base of the Guodikeng Formation and 24.7 m below its top, because the boundary between the F. postera and F. verchojanica zones can be correlated directly in the Werfen beds with the conodont-defined Permian– Triassic boundary and, in the Germanic Basin, with the minimum value in d13Ccarb (Korte & Kozur 2005). The LOD of Dicynodon (219 m above the base of the Guodikeng Formation) lies at Dalongkou only nine metres above the FAD of F. verchojanica. If we consider the extremely high sedimentation rates in the Guodikeng Formation, Fig. 1. Correlation of the conchostracan zonations of the Changhsingian and in the lowermost Triassic of the Germanic Basin, northern Siberia (Tunguska Basin and Taimyr) and northwestern China (Dalongkou, Xinjiang) with the pelagic open sea conodont and ammonoid zonations of the Tethys (Iran, Transcaucasia) (the two Otoceras zones are Perigondwanan), the Tethyan shallow water conodont zonation (Southern Alps, Hungary), and the conodont zonation of intraplatform basins in South China. Not to scale. Metres in the Dalongkou column are metres above the base of the Guodikeng Formation in the southern limb of the anticline, measured by Dr. Spencer Lucas (New Mexico Museum of Natural History and Science, Albuquerque) during the American-Chinese project in Dalongkou, sponsored by the NGS, USA. The asterisk mark [*] in the figure indicates that Suchonella is not a conchostracan, but an ostracod which is typical of the Middle and Upper Permian. Abbreviations: Wuchia., Wuchiapingian; Dzhulf., Dzhulfian; Tut., Tutontchana. Additional data for Dalongkou: 171.2 m above the base of Guodikeng Formation is the FAD of Falsisca postera; at 111 and 132 m above the base of Guodikeng Formation are the lower and upper boundaries of the Falsisca turaica–F. zavjalovi Zone.

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the nine-metre difference represents only a very short time interval. Thus when conchostracans are not present, the LOD of Dicynodon can be used as a proxy for recognizing the Permian–Triassic boundary, as in South Africa. The first occurrence of Lystrosaurus at Dalongkou is not the FAD of this genus, but rather its local FOD (First Occurrence Datum) at Dalongkou. It occurs there within the F. eotriassica Zone, in the long normal interval that straddles the Permian– Triassic boundary (palaeomagnetic chron ‘1n.1n’ in Bachmann & Kozur 2004). In South Africa, Lystrosaurus begins in the short reversed interval (palaeomagnetic chron ‘0r’ in Bachmann & Kozur 2004) immediately below this long interval (Ward et al. 2005). Sediments representing this short reversed interval have not been found at Dalongkou, either because they were removed by local erosion that created a short gap (occasionally present in subaerial sediments and actually documented in some levels of the upper Guodikeng Formation) or else because they were not sampled. A combination of erosional thinning and lack of sampling of the thinned interval also is possible. In this critical interval, the section consists almost exclusively of mudstones, which were not sampled by Metcalfe et al. (2009). Metcalfe et al. (2009) have suggested that a short reversed horizon in the lowermost Guodikeng Formation and upper Wutonggou Formation corresponds to the short reversed interval that lies immediately below the long normal interval straddling the Permian–Triassic boundary. This view, however, contradicts all available biostratigraphic data. The basal Guodikeng and upper Wutonggou formations have conchostracan faunas with species that occur well below the top of the characteristic Bipemphigus –Megasitum– Trimpemphigus conchostracan fauna, but the short reversed horizon just below the Permian–Triassic boundary lies well above this fauna in an interval where these three characteristic genera are no longer present and the Falsisca species also are different. In the Germanic Basin, it can be shown that the top of this reversed interval lies in the lowermost F. eotriassica Zone, which begins in the Dalongkou section more than 100 m above the top of the short reversed interval documented by Metcalfe et al. (2009). The uppermost Permian short reversed interval often contains tuff fallout (as in the Nedubrovo Formation northeast of Moscow) or common volcanic microsphaerules (as in the lower Fulda Formation of Germany) and yields sporomorphs of the Triquitrites proratus Zone together with megaspores of Otynisporites eotriassicus Fuglewicz. O. eotriassicus is common in the upper Guodikeng Formation, but completely absent throughout the lower half of this formation (Metcalfe et al. 2009). As the top of the short reversed interval lies about 100 m below the first

appearance of O. eotriassicus, the palynological data and the conchostracan data are in agreement that the reversed horizon in the lowermost Guodikeng Formation and in the upper Wutonggou Formation are well below (and thus not correlative with) the short reversed horizon immediately below the long normal palaeomagnetic zone that straddles the Permian–Triassic boundary. The reversed horizon that ends within the lowermost Guodikeng Formation corresponds to a reversed interval found in the uppermost Zechstein 3 and Zechstein 4, as shown by Szurlies (2007). This conclusion is supported by the Milankovitch cyclicity pattern. The Guodikeng Formation includes four readily recognizable short eccentricity cycles with five readily recognisable precession cycles. Thus, the Guodikeng Formation represents about 400,000 years of time, which means that the upper limit of the reversed horizon in the upper Wutonggou and basal Guodikeng formations must have been about 400,000 years before the end of Guodikeng Formation deposition and more than 350,000 years before the beginning of the F. verchojanica Zone (¼ base of the Triassic). In the Germanic Basin, the top of the latest Permian short reversed palaeomagnetic interval lies about 200,000 years below the beginning of the Triassic (Bachmann & Kozur 2004) and, in the conodontdated beds of central and northwestern Iran and in the Southern Alps, the top of the short reversed interval also occurred about 200,000 years before the beginning of the Triassic (Kozur 2007). Previous descriptions of conchostracans from Dalongkou (Liu 1987, 1989) are rather difficult to evaluate because the taxonomy used is partly outdated. Important genera are assigned to other unrelated genera, such as the Permian genus Megasitum Novozhilov, which is assigned to the Gandarian (Dienerian) genus Cornia Lyutkevich. For other genera, such as Tripemphigus Novozhilov 1965, a junior synonym Trinodus Liu 1987 was established. Falsisca cf. F. kanandaensis Novozhilov (sensu Liu 1989) belongs to Falsisca turaica (Novozhilov, in Molin & Novozhilov 1965), and Falsisca beijiangensis Liu, 1987 is a junior synonym of F. zavjalovi (Novozhilov 1970). In addition to the southern limb of the Dalongkou anticline, the northern limb of the anticline also was studied but not sampled in great detail. The conchostracan succession was found to be the same, and the so-called ‘Sesame cake beds’ (limestones with very well-preserved conchostracans of the Beijianglimnadia –Bipemphigus–Falsisca – Polygrapta– Tripemphigus fauna) are exceptionally well developed there. In the Xiaolongkou section, we mainly studied the c. 300 m-thick Lower Triassic Jiucaiyan Formation to determine the relationship between the LOD of Lystrosaurus and Falsisca. About 200 m above the base of the

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Jiucaiyan Formation, there is a whitish sandstone that has in its upper part a thin intercalation of greenish-grey shale. The highest occurrence of Falsisca in the Xiaolongkou section is in these beds. By comparison to other sections, this interval is the highest possible LOD for Lystrosaurus. Thus, Falsisca seems to have the same uppermost range as the tetrapod Lystrosaurus, which does not occur above the Gangetian (upper Griesbachian). The base of the Gangetian was excellently defined in Mojsisovics et al. (1895) as the base of the Otoceras woodwardi Zone s.l. (including the Otoceras fissisellatum fauna) and was included as the lower substage of the Brahmanian Stage in the Himalayas. This level corresponds to the FAD of Hindeodus parvus, which marks the base of the Triassic. The Griesbachian was established as a stage 70 years later on Axel Heiberg Island, Arctic Canada by Tozer (1965) and its base was defined by the base of the Otoceras concavum Zone. The lower Griesbachian is older than the FAD of H. parvus, and the second ammonoid zone of the Griesbachian has a Permian conodont fauna (e.g. Kozur 2007). Therefore, we use here and in the following sections the Brahmanian stage and the Gangetian substage in its original definition, because use of the Griesbachian substage would improperly put the lowermost substage of the Triassic partly in the Upper Permian and partly in the Lower Triassic.

Germanic Basin During the Triassic, the Germanic Basin (Central European Basin) was a large depocentre that extended (west to east) from England through the North Sea, the Netherlands, and northern and central Germany to eastern Poland and (north to south) from southwestern Sweden and Denmark through northern, central and western Germany, eastern France, and southern Germany to nonAlpine Switzerland and southeastern France. Depending on the position of sea level at any given time and the regional climate, very different sedimentation regimes were present in the basin, ranging from fluvial in the marginal areas (and sometimes also in the central parts of the basin, e.g. parts of the Schilfsandstein) to lacustrine, brackish, marine, or hypersaline (with large volumes of gypsum and halite) in the central part of the basin. Large flood plains and hypersaline flats also were present during some parts of the Triassic. The strata within the Germanic Basin are divided into three groups; from the base, they are the Buntsandstein Group, the Muschelkalk Group, and the Keuper Group (Bachmann & Kozur 2004). It is this three-fold division that gave the Triassic System its name (von Alberti 1834). The Buntsandstein ranges in age from latest Permian to early Anisian, and mostly consists of continental deposits

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together with early Anisian marine to hypersaline facies. The Muschelkalk ranges in age from early Anisian through late Ladinian and consists of marine to hypersaline facies. The Keuper ranges in age from late Ladinian through Rhaetian and consists of continental and hypersaline deposits interbedded with subordinate marine facies. During deposition of the Lower and Middle Buntsandstein, there was a weak marine influence that originated from the northwest through a connection to the Boreal Sea. This connection allowed slightly brackish depositional environments to develop, especially during Olenekian time. During the Middle and Upper Triassic, connections to the Tethys in the south were present through gateways in the east to the Dobrogea–Ku¨re Ocean, to the northern shelf of the Meliata Ocean through the Eastern Carpathian and Upper Silesian gates, and through gateways in the S –SW to the northwestern shelf of the Meliata Ocean through the Alemannic Gate (in the present Alpenrhein depression) and to the Westmediterranean Sea in southeastern France through the Burgundian Gate. The occurrence of these thin marine intervals in the Germanic Basin allows correlation of its conchostracan-bearing continental beds to nearby marine sections in the Tethys region (e.g. Kozur 1972, 1975, 1999a, b, 2005; Brack et al. 1999; Bachmann & Kozur 2004; Kozur et al. 1993; Kozur & Mock 1993; Urlichs & Tichy 2000; Kozur & Bachmann 2005, 2006, 2008b). Each marine interval is well mapped and contains a stratigraphically useful marine fauna. Sporomorphs also have yielded important data for correlating in detail the Germanic Triassic with the international marine timescale (e.g. Schulz 1967, 1996; Scheuring 1970; Lund 1977; OrlowskaZwolinska 1983; Heunisch 1996, 2005; FialkowskaMader 1999; W. Wille, Mo¨ssingen, pers. comm.). In central and northern Germany, the conchostracanrich Lower and Middle Buntsandstein mid-basin facies (uppermost Permian to lowermost Anisian) are well correlated across the central Germanic Basin and with the international scale as summarized in Bachmann & Kozur (2004) and in Kozur & Bachmann (2005, 2006, 2008b) (see Figs 2 & 3). Correlation with fluvial marginal facies in the southern Germanic Basin has been more difficult, because no conchostracans have been found there. The correlation of the Germanic Lower Triassic with the international marine timescale has been well established. Differences in results between the various correlation methods do not exceed two or at most three short eccentricity cycles, a very low level of discrepancy for correlation of continental sequences with the international marine scale. Only the Olenekian–Anisian boundary within the Germanic Basin is controversial. Nawrocki & Szulc (2000) placed the base of the Anisian within the lower Jena Formation of the lowermost

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Fig. 2. Formations of the Lower and Middle Buntsandstein (Lower Triassic) and their correlation with the international marine timescale, their numeric ages, Milankovitch cyclicity and palaeomagnetic normal and reversed intervals (slightly modified from Kozur & Bachmann 2008b). Palaeomagnetic intervals and Milankovitch cyclicity after Szurlies (2007), but with 11 cycles shown in the Bernburg Formation. Left column: Compiled new radiometric ages from the marine Lower Triassic after Galfetti et al. (2007); Lehrmann et al. (2006); Mundil et al. (2004) and Ovtcharova et al. (2006). Right column: Extrapolated numerical ages for the Germanic Triassic in italic script; age of the base of the Triassic and the Anisian in normal script. A, Biostratigraphically correlated base of the Olenekian after Kozur & Seidel (1983), Kozur (1993, 1999a) and Kozur & Weems (2007); B, Olenekian base from palaeomagnetic correlation (Bachmann & Kozur 2004; Szurlies 2007). L., Lopingian Series; C., Changhsingian Stage.

Muschelkalk close to the base of Middle Triassic as originally chosen by von Alberti (1834). Brugman (1986) placed the Olenekian –Anisian boundary within the upper part of the Solling Formation in the Buntsandstein based on the FAD of the sporomorph Hexasaccites thiergartii (Ma¨dler) Kozur. Bachmann & Kozur (2004) and Kozur & Bachman (2005, 2006, 2008b) placed the base of

the Anisian at the base of the Stammen Beds within the upper Solling Formation (Fig. 2). Hounslow et al. (2007) and Szurlies (2007) also placed the base of the Anisian within the Solling Formation using correlation of magnetic reversals to the marine record, but their boundary is slightly lower than that of Bachmann & Kozur (2004) and Kozur & Bachmann (2005, 2006, 2008b) because they

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Fig. 3. Subdivisions of the Solling Formation in the Solling Mountains region (northwestern Germany). Wavy line represents a widespread unconformity of short duration within the Solling Formation.

placed the marine section base of the Anisian at a magnetic reversal slightly lower than the FAD of the conodont Chiosella timorensis rather than at its FAD. As the location of the base of the Anisian in the Germanic Basin falls within the upper Solling Formation, as determined by palynomorphs, conchostracans and palaeomagnetics, this correlation now seems to be well founded and is not likely to change to any great degree in future investigations. Therefore, we use this boundary. Only a few new results can be added to the synthesis of Middle Triassic stratigraphic subdivisions, correlation with the international marine timescale, numeric ages and cyclicity sequence published by Bachmann & Kozur (2004) and Kozur & Bachmann (2005, 2006). The 244.60 + 0.36 Ma age determination for the Balatonites shoshoensis Zone (Ovtcharova et al. 2006) has been cited by Kozur & Bachmann (2008b), and it fits well with the data shown in Bachmann & Kozur (2004) and Kozur & Bachmann (2005, 2006). Similarly, the 246.83 + 0.31 Ma age determination for the Platycuccoceras beds Ovtcharova et al. (2006) fits well into the scheme proposed in Bachmann & Kozur (2004) and later papers by these authors. These new data are shown together with earlier data in Figure 4. The Platycuccoceras Beds correspond to the middle Bithynian. From a similar level Lehrmann et al. (2006) reported a 246.77 + 0.13 Ma age determination, described in more detail later by Galfetti et al. (2007). The age of the stratigraphic horizon from which this date was determined is not clear, because the conodont species Chiosella gondolelloides (Bender) and Nicoraella kockeli (Tatge) both are reported from this locality but they are not known to co-occur anywhere else. Most probably N. germanica (Kozur) is the species that is present instead of N. kockeli, together with a form transitional between C. gondolelloides to N. germanicus. Such a fauna is known to occur in Pietra dei Saracini in the Sosio Valley (western Sicily, Italy) in the lower to middle Bithynian.

It is within the Norian and Rhaetian parts of the Keuper that the largest remaining problems exist with the lithostratigraphic subdivisions of the sequence and their correlation within the Germanic Basin and with the international marine timescale. This is due to unclear stratigraphic definitions, miscorrelations, and because it is unclear exactly how many disconformities are present in the sequence and whether they are local or regional in nature. As is common in continental sequences, the continental sequences within the Germanic Basin include stratigraphic gaps at various horizons. In the Buntsandstein, these gaps have been recognized for a long time and often have been used as the boundaries between formations, though some are present within a formation (e.g. the intra-Solling unconformity below the Stammen Beds). In the Keuper, unconformities for a long time either remained undiscovered or were disregarded. Beutler has in several papers drawn attention to the numerous unconformities within the Keuper and has summarized his results in Deutsche Stratigraphische Kommission (2005). All unconformities are named as disconformities, but a few actually occur only locally and are erosional surfaces at the base of fluvial sandstone bodies (e.g. within the Schilfsandstein), a few others are paraconformities, and some reflect a sharp change in facies (e.g. a depositional shift from marine to limnic or to very slightly brackish depositional environments) that was incorrectly interpreted as a disconformity (such as the transition between the middle Rhaetian Contorta Beds and the Triletes Beds). Except for the unconformity at the base of the Schilfsandstein (Stuttgart Formation), which is a basin-wide unconformity caused by a very pronounced sea-level lowstand, other verifiable unconformities within the Keuper only occur at the basin margin or on swells within the basin. In the centre of the basin, these disconformities are either missing or only represented by subtle paraconformities.

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Fig. 4. Lithostratigraphic subdivisions of the Germanic Middle Triassic and their correlation with the international marine timescale, Milankovitch cycles (short eccentricity cycles) and numeric ages. Slightly modified from Kozur & Bachmann (2008b). OB, Oolithba¨nke; TB, Terebratelba¨nke; CB, Cycloidesbank; MM., Middle Muschelkalk. Numeric ages in bold script are compiled measured radiometric data from the Tethys; numeric ages in italic script are calculated numeric ages for the base of the Anisian, Ladinian and Carnian stages, as well as for the Longobardian substage. DGB, Dolomitische Grenzbank which has the LO of Costatoria costata (Zenker) and the FAD of Myophoria vulgaris (von Schlotheim).

The largest single problem has been determining the age of the Trossingen Formation (better known as the Knollenmergel or Feuerletten), which contains the famous Plateosaurus site at Trossingen. Originally, the Knollenmergel strata were placed below

the Malschenberg Sandstone, formerly called the Bonebed- or Rhaetsandstein. Etzold & Schweizer (in Deutsche Stratigraphische Kommission 2005) recently have correlated the Malschenberg Sandstone with the ‘4. Stubensandstein’ (s4) of the

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Kraichgau region. Seegis (report to the Keuper Group of the Deutsche Stratigraphische Kommission, 2004 and pers. comm. to H. W. Kozur 2008) has found the lower Rhaetian ostracod species Rhombocythere wicheri (Anderson) in the ‘s4’ of the Kraichgau region, which would seem to confirm the correlation of Etzold & Franz (2005) because they also reported R. wicheri from the Malschenberg Sandstone in the borehole Malschenberg 1. Therefore, the ‘s4’ unit, as mapped in the Kraichgau region, can be re-assigned to the Malschenberg Sandstone, and its age can be established as early Rhaetian as in the borehole Malschenberg 1, where not only R. wicheri indicates an early Rhaetian age for the Malschenberg Sandstone but also early Rhaetian sporomorphs are present (Heunisch 1996). Correlation of the Malschenberg Sandstone with the ‘4. Stubensandstein’ of the Lo¨wenstein area, however, is not supported by any palynological or palaeontological data. In the Lo¨wenstein area, the Knollenmergel lies above strata assigned to ‘4. Stubensandstein,’ which was correlated with minimal supporting data to the lower Rhaetian Malschenberg Sandstone. This correlation required that the Knollenmergel there had to be younger than the Malschenberg Sandstone, which would place it in the upper part of the lower Rhaetian. For this reason, the Knollenmergel either was placed in an exceptionally high stratigraphic level (and the strata below the Malschenberg Sandstone assigned to the ‘Hangendletten 3’ above the ‘3. Stubensandstein’ instead of to the Knollenmergel) or else these strata were assigned to a lower Knollenmergel (to distinguish them from the ‘Knollenmergel’ above the ‘4. Stubensandstein,’ which was therefore called ‘upper Knollenmergel’). Etzold & Schweizer (in Deutsche Stratigraphische Kommission 2005) demonstrated that the Trossingen Formation (type Knollenmergel) in its type section begins above the ‘Stubensandstein 2.3.’ Thus, the oldest part of the type Trossingen Formation (Knollenmergel) corresponds to the ‘Hangendletten 2’ of the Stubensandstein. In the borehole Malschenberg 1, located in the immediately overlying basal part of the ‘3. Stubensandstein,’ there is a conchostracan fauna with Shipingia mcdonaldi n. sp. and Norestheria n. sp., both of which indicate a middle Norian age for this fauna. Therefore, the lowermost part of the Trossingen Formation is middle Norian. A distinct climate change occurred between the Norian and Rhaetian. The lower Rhaetian reflects a much wetter climate than was present in the Norian, so lower Rhaetian sediments are often grey and contain plant detritus. This makes it very improbable that the Knollenmergel strata could range from the higher part of the middle Norian all the way up to the upper part of the lower Rhaetian

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without reflecting this climatic trend. This problem has caused Kozur and Andreas Etzold (Emmendingen) to investigate the age of the Knollenmergel and related problems in detail. In a personal communication to Etzold, Wolfgang Wille (Mo¨ssingen) has stated that the sporomorph association of the ‘4. Stubensandstein’ does not correspond to the lower Rhaetian association of the Malschenberg Sandstone, but rather to the (Norian) sporomorph association of the ‘3. Stubensandstein.’ This clearly shows that the ‘4. Stubensandstein’ of the Lo¨wenstein area has been incorrectly assigned to that unit and instead is older than the Malschenberg Sandstone and older than the lower Rhaetian. This important result of Wille’s research indicates that Knollenmergel strata do not occur above the level of the lower Rhaetian Malschenberg Sandstone, but rather lie below this level. Thus, the lowermost part of the Trossingen Formation (Knollenmergel) is middle Norian in age, the Trossingen Formation lies below the lower Rhaetian Malschenberg Sandstone as previously believed, and most of the Trossingen Formation is Sevatian in age. Plateosaurus has not been found as low as the middle Norian part of the Knollenmergel at Trossingen, so therefore this famous dinosaur occurs in the Sevatian part of the Trossingen Formation at Trossingen. This is also true for the Plateosaurus occurrence at Halberstadt (Jaekel 1914; Sander 1992, 1999). In Hallau (Switzerland), below marine Hettangian beds that are at most 2 m thick, there are other beds that begin with a thick compact Steinmergel and are overlain by greenish grey marls with thin dark layers that contain plant detritus. These wetclimate beds are surely younger than the typical reddish Knollenmergel beds, which lack dark layers with plant detritus. However, according to palynologic investigations by Achilles & Schlatter (1986), they are also older than Rhaetian (older than the Malschenberg Sandstone). Therefore, this interval seems to represent a sequence that in most areas was erosionally removed between the Knollenmergel and overlying middle Rhaetian or marine Hettangian beds. The Rhaetian Exter Formation sensu Beutler (in Deutsche Stratigraphische Kommission, Menning et al. 2005) is not a very convincing stratigraphic unit. Duchrow (1984) originally introduced this name as the Exter Group and divided it into three formations: the lower Rhaetian brackish Rinteln Formation, overlain unconformably by the middle Rhaetian marine Oeynhausen Formation, in turn overlain by the upper Rhaetian to earliest Hettangian fresh water to slightly brackish Valbruch Formation. This subdivision still seems logical. As this problem remains to be discussed by Bachmann & Kozur (in prep.), we continue use of the ‘Exter Formation’ for now.

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Only a few new biostratigraphic data points have been added to the Upper Triassic framework of the Germanic Basin since Bachmann & Kozur (2004). One new result is that the ostracod Simeonella nostorica Monostori has been reported from the Lehrberg Beds of the Weser Formation. This species is common in the marine lower Tuvalian of Hungary and in the marine lower Tuvalian of Austria. In Hungary, it occurs also in slightly hypersaline marine beds. Seegis (1997) illustrated wellpreserved material of this species, but incorrectly assigned it to the late Julian species Simeonella brotzenorum alpina Bunza & Kozur, which is morphologically distinct. Other new biostratigraphic data in the Norian, Rhaetian and basal Hettangian are discussed below, under the relevant conchostracan zones. The Triassic– Jurassic boundary in the Germanic Basin also is problematic. Generally, it is drawn where the fresh water to slightly brackish Triletes beds are directly overlain by marine ammonoidbearing Hettangian beds. However, the lowest ammonoid occurrences in these beds generally are specimens of Psiloceras psilonotum (Quenstadt), which is probably a junior synonym of P. sampsoni (Portlock). This species is not known to occur as low as the base of the P. planorbis Zone, but rather first appears in the higher part of this zone. Although not yet officially defined, the proposed base of the Jurassic lies still lower than the P. planorbis Zone at the base of the P. spelae Zone. Marine beds in northern Germany below the Psilonotum Beds, generally without ammonoids, have been assumed to correlate with the ‘pre-planorbis beds,’ but in the central part of the Germanic Basin these beds probably are no older than the lower P. planorbis Zone. In western England (e.g. St. Audries Bay), the RhaetianHettangian boundary is within marine beds, but even so diagnostic ammonoids of the P. spelae Zone are missing. In the central Germanic Basin, the fresh water to slightly brackish Triletes beds probably straddle the Triassic-Jurassic boundary. The exact position of this boundary is now under study in the temporary outcrop at the A4 highway at Moseberg in the western Thuringian Basin. The sporomorphs from this locality are being investigated by the Utrecht group (Kuerschner, Bonis) and values for dCorg are being evaluated by Kraus, Korte, Bachmann & Kozur. The initial negative carbon isotope excursion just below the base of the Jurassic probably lies somewhat below the Red Levallois Clays, which for the first time have been recognised in western Thuringia. This same level in France contains the monospecific Euestheria brodieana fauna, which is characteristic of the upper Rhaetian. At Moseberg, this level has not yielded conchostracans, but immediately below this level the monospecific E. brodieana fauna does occur.

The sporomorph and carbon isotope investigations will show whether the Red Levallois Clays correspond to the similarly red Schattwald Beds in the proposed GSSP for the base of the Hettangian at Kuhjoch, in the Northern Alps. If this proves to be the case, then the top of the Red Levallois Clays would represent the most practical mapping horizon for the continental Triassic– Jurassic boundary in the Germanic Basin because the FAD of P. spelae Guex, Taylor, Rakus & Bucher is immediately above the Schattwald beds. The beginning of the marine P. sampsoni beds in the Moseberg section is much higher than the Red Levallois Clays in this section. This indicates that the Liassic marine flooding of the Germanic Basin did not begin until well after the beginning of the Jurassic. No unconformity can be recognised above the Red Levallois Clays in the Moseberg section, so the unconformity D8 probably is not present either in the western Thuringian Basin or in northern Germany. Conchostracans occur regionally and stratigraphically in various parts of the Germanic Basin. In the uppermost Permian to middle Spathian interval, they are common throughout the basin centre in northern and central Germany and in Poland, and can be found there nearly in every outcrop of shales and siltstones from the Calvo¨rde Formation up through the Hardegsen Formation. Only in the upper Calvo¨rde Formation is there an interval in which the dry Zechstein climate returned, producing hypersaline lake and sabkha deposits but no fresh water lake or pond deposits. During this time interval, conchostracans seem to have totally disappeared from the Germanic Basin, because in the uppermost Calvo¨rde Formation a new immigrant fauna appears that is unrelated to the fauna below the arid interval. Throughout other intervals in the Lower Triassic, even in the marginal, mainly sandy facies of lake deposits in central Germany, conchostracans are common in shaly and silty intercalations, such as in the Großwangen region south of Halle. Sometimes conchostracans can even occur in sandstones, such as Estheriella species, that are found around Bad Salzungen in southern Thuringia. Conchostracans also are rather common in marginal facies in southeastern Poland (e.g. Holy Cross Mountains, Ptaszyn´ski & Niedz´wiedzki 2004, 2005, 2006a, b). In marginal facies in England, Lower Triassic conchostracans are very rare but still present. Surprisingly, no conchostracans have been reported from the predominantly fluvial Lower Triassic strata of Southern Germany, not even from shaly intercalations, and personal small-scale sampling for them so far has not been successful. In the upper Spathian to lowermost Anisian Solling Formation, rich conchostracan faunas are found in surface outcrops only in northwestern

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Germany, where rich conchostracan faunas can be found throughout most of the Solling Formation in the Solling Mountains. In the uppermost part of the Solling Formation (lowermost Anisian Stammen Member in the Solling type area) and the correlative Thuringian Chirotherium Sandstone, conchostracans are rather rare but do occur in Thuringia (Kozur & Seidel 1983a) and in the northern part of southern Germany (Kozur et al. 1993). In the lower Anisian Upper Buntsandstein (Ro¨t), the eastern Germanic Basin and the basin centre in the northern Germanic Basin had marine and hypersaline environments that were not suitable for conchostracans, but conchostracans of this age can be found in the northern part of southern Germany, western Germany, and eastern France. In the lowermost part of the Muschelkalk, conchostracans occur in the western marginal facies (Lorraine, France). The largest part of the Muschelkalk, however, formed in marine or (in the Middle Muschelkalk) in hypersaline environments and therefore does not contain conchostracans. However, in the Longobardian part of the Upper Muschelkalk, above the Cycloidesbank, conchostracans are locally common in shales found between the ammonoid-rich limestones and marls. These shales also have a brackish ostracod fauna with abundant Pulviella teres (von Seebach). The associated conchostracan fauna is very monotonous, consisting mainly of Euestheria minuta (von Zieten) and sometimes also Euestheria franconica (Reible). A rich conchostracan fauna also can be found in the predominantly brackish Longobardian Erfurt Formation (Lower Keuper). Here also, E. minuta is almost always the only species present, though one specimen of Lioleaiina n. sp. was found by Warth (1969). In contrast, conchostracans are very rare in the mainly hypersaline Longobardian part of the Grabfeld Formation (Lower Gypsum Keuper) below the Cordevolian ‘Estheria’ Beds. Where conchostracans do occur, the typical Longobardian monospecific Euestheria minuta fauna is found. Rich conchostracan faunas occur in the ‘Estheria’ Beds of the upper Grabfeld Formation. The sporomorphs Patinasporites densus Leschick and Vallasporites ignacii Leschik have their lowest occurrences in the Germanic Basin in the basal ‘Estheria’ Beds, and their FAD lies close to the base of the Carnian at the GSSP Stuores Wiesen locality in the Southern Alps (Italy). Therefore, the base of the ‘Estheria’ Beds coincides with the base of the Carnian. As the FAD of Laxitextella multireticulata (Reible) also lies at the base of the ‘Estheria’ Beds, it is an excellent proxy for locating the base of the Carnian. Transitional forms between this species and L. laxitexta (Sandberger) occur in the westernmost part of the Southern Alps within the Cordevolian (Kozur & Mock 1993). In the

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middle ‘Estheria’ Beds, the late Cordevolian conchostracan fauna characteristic of the Laxitextella laxitexta Zone occurs, which includes the first appearance of the genus Gregoriusella n. gen. From this level, Gregoriusella occurs upward into the lowermost Rhaetian. Rich conchostracan faunas also are found in the Basisschichten (Osterhagen Horizon) of the Schilfsandstein. The preservation of the recovered conchostracan is poor, however, though Laxitexella of the L. laxitexta group and Gregoriusella of the G. fimbriata group definitely are present. Rich conchostracan faunas occur also in lacustrine strata of the lower Schilfsandstein between the top of the Basisschichten and the Gaildorf Horizon. They are well-preserved but have not been studied yet in any detail. Several species belonging to the L. laxitexta group, however, definitely are present. A rich and diverse conchostracan fauna occurs also in the Gaildorf Horizon at the base of the upper Schilfsandstein. This fauna consists partly of new species of the L. laxitexta group, and also includes transitional forms between Laxitextella and Anyuanestheria, as well as Palaeolimnadia n. sp. The rich conchostracan fauna of the Schilfsandstein remains to be described. During the Tuvalian, the distribution of conchostracans in the Germanic Basin became very restricted. Before the Tuvalian they occurred throughout the entire Germanic Basin, but in the Tuvalian they are absent in the hypersaline central parts of the basin and also in the fluvial deposits that formed within a narrow marginal belt around the basin centre. During the Tuvalian, they only occur in those marginal areas where fluvial fresh water input changed a spatially restricted marginal region of hypersaline marine environments into brackish water environments. Such environments have been found in some marginal areas around the inverse estuary of the basal Tuvalian Dolomie de Beaumont (e.g. at Stuttgart, Haussmannstrasse 44, which has a rich conchostracan fauna, Kozur & Weems 2007); and in eastern France in the Hombourg-Buange D 978 road cut, region Thionville (Bourquin & Durand 2006), or in larger fresh water lake to brackish areas around the hypersaline marine ingression of the Lehrberg beds in the eastern Germanic Basin of Poland (e.g. in Krasiejo´w, Upper Silesia) or in southern Germany (Seegis 1997). Conchostracans also occur in fresh water pond deposits close to rivers that terminated in hypersaline flats, as in the Coburg Sandstein (Hassberge Formation, upper Tuvalian; Fig. 5) of the Hassberge and Steigerwald regions of Franconia (northern Bavaria). The only conchostracan occurrences from the Coburg Sandstein in the Germanic Basin are (Kelber, pers. comm.): (1) Coburg Sandstein quarry just east of Passmu¨hle in the Ebelsbach

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Fig. 5. Lithostratigraphic subdivisions of the Germanic Upper Triassic and their correlation with the international marine timescale and numeric ages. Modified after Kozur & Bachmann (2008b). Date with an asterisk mark [*] is 40 Ar/39Ar data from the Adamanian of Ischigualasto, Argentina (Rogers et al. 1993), corresponding to a middle to late Tuvalian level between the Lehrberg Beds and the top of the Weser Formation. The 230.91 + 0.33 Ma date of Furin et al. (2006) is from the basal Carnepigondolella zoae Zone, a level somewhat older than the Lehrberg Beds of the Weser Formation. The 225 + 3 Ma date (Gehrels et al. 1986, 1987) is from volcanics in the lower Norian E. quadrata Zone in SE Alaska. The 201.5 Ma date for the Triassic–Jurassic boundary is based on a biostratigraphic re-dating (Kozur & Weems 2007) as latest Rhaetian of the lower lava flow of the CAMP volcanics in the Newark Supergroup, and on radiometric data from a well-dated Rhaetian– Hettangian boundary section in Peru by Schaltegger et al. (2008). Calculated numeric ages for the base of the Carnian, Norian and Rhaetian stages are in italic script. Wavy line in the upper Arnstadt Formation represents an unconformity of short duration underlain by pedogenic sediments. Wavy line below the Contorta Beds represents an unconformity of short duration, especially where the Lower Contorta Clay is missing. 2.3, 3, 4, Stratigraphic position of the Stubensandstein subdivisions designated as Stubensandstein 2.3, Stubensandstein 3 and Stubensandstein 4; OBM, Obere Bunte Mergel; U., Upper; L., Lower.

valley, the enlarged quarry Fra¨nkische Schleifsteinwerke GmbH (the former quarry Ankenbrand): 508000 56.1900 N; 108400 55.5200 E.; (2) Coburg Sandstein quarry Vetter, southeast of Eltmann: 498570 37.7400 N; 108400 22.8000 E (Reimann & Schmidt-Kaler 2002); and (3) Coburg Sandstein quarry at the Steinberg northwest of Obersteinbach: 498540 31.0700 N; 108310 15.5200 E (Reimann & Schmidt-Kaler 2002). Conchostracans occur in a

similar geological setting in the Heldburggipsmergel Member of the uppermost Weser Formation (borehole Groß-Scho¨nebeck-1 at 778 m, northern Germany) and in the Weser Formation at the Schwanberg at Iphofen (Franconia, northern Bavaria). The occurrence in the Groß-Scho¨nebeck-1 borehole is the only occurrence of conchostracans in the Weser Formation of northern Germany. According to Gerhard Beutler (pers. comm.), in this region

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a significant fresh water influx prevented the accumulation of gypsum in the Heldburgipsmergel Member of the Weser Formation, which accords well with the occurrence of conchostracans in this region. In the Norian and Rhaetian conchostracans again are found throughout the entire Germanic Basin, though at most localities they are rare. They are common only in a few levels, especially in the lower and middle Arnstadt Formation. In the upper Arnstadt Formation, they have been found only in northern Germany. Lower Rhaetian conchostracans have been found in boreholes in northern Germany, and also occur rarely in the western Thuringian Basin (e.g. the outcrop at the A 4 highway, 800 m west-northwest of Moseberg west of Eisenach) and in southern Poland (at Lipie, Upper Silesia). The upper Rhaetian monospecific Euestheria brodieana fauna is much more widely distributed and is found in England (Lilstock Formation), in northern Germany, in the Moseberg section in the western Thuringian Basin, and occasionally mentioned but not yet figured in southern Germany. Especially interesting is the occurrence of the monospecific E. brodieana fauna in the Red Levallois Clays at the very top of the Rhaetian section in eastern France (Battarel & Gue´rinFraniatte 1971). This occurrence indicates that the genus Bulbilimnadia, common in the early Hettangian, was not present yet in the latest Rhaetian. Conchostracans from continental basal Hettangian (correlative to the marine P. spelae Chron) deposits in the Germanic Basin so far have been found only in the lower Hettangian Sołtyko´w exposure, Holy Cross Mountains, Poland (Pien´kowski 2004; Pien´kowski & Niedz´wiedzki 2009). SEM photographs and data detailing the age of this sequence, sent to us by Drs Grzegorz Pien´kowski and Grzegorz Niedz´wiedzki (Warsaw), show that this fauna contains Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp. This oldest Hettangian species also is found in the Culpeper Basin of Virginia (eastern United States) and in the upper Whitmore Point Member of the Moenave Formation in the St. George-Zion area of Utah and northern Arizona (western United States).

Newark Supergroup rift basins The Newark Supergroup includes all strata and volcanic flows and tuffs deposited in 30 tectonic basins that filled mostly during the Late Triassic and earliest Jurassic in what is now the eastern United States and eastern Canada (Fig. 6). A comprehensive bibliography of much of the earlier literature concerning the Newark Supergroup is in Margolis et al. (1986), and Newark Supergroup stratigraphic nomenclature has been summarized thoroughly by

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Luttrell (1989), including history of units, age of units, and location of type sections or areas. Since the summary of Luttrell, the only regional studies that have been conducted are by Weems & Olsen (1997) and Faill (2003). The stratigraphic nomenclature for formations and members followed here is that used by Luttrell (1989) and slightly modified by Weems & Olsen (1997). Prior to the work of Weems & Olsen (1997), group names were established for individual basins, creating an unnecessarily complex stratigraphic nomenclature. Weems & Olsen (1997) chose instead to define only three regionally recognizable groups within the Newark Supergroup based on the common tectonic framework of the supergroup. These are the Chatham Group, which includes the pre-CAMP (Central Atlantic Magmatic Province) sedimentary sequence, the Meriden Group, which includes the CAMP volcanics and interbedded sedimentary sequences, and the Agawam Group, which includes the post-CAMP sedimentary sequence (Fig. 7). Within the Chatham Group, Olsen (1997) recognized a regionally extensive unconformity that separated his tectonostratigraphic sequence TS II from sequence TS III. Kozur & Weems (2007) have confirmed the presence of this unconformity, because the youngest dated strata beneath this unconformity are early Cordevolian in age and the oldest dated strata overlying this unconformity are late Julian. This unconformity therefore encompasses some or perhaps all of late Cordevolian (late early Carnian) through early Julian (early middle Carnian) time. All strata in the small and shallow Scottsburg, Randolph, Roanoke Creek, Briery Creek, Flat Branch, and Deep Creek basins lie below this unconformity, and the basal strata of the Farmville (basal sandstone and overlying unnamed coalbearing shale-rich unit), Richmond (Tuckahoe Formation and Vinita Shale), Taylorsville (Stagg Creek and Falling Creek formations), and Gettysburg basins (basal beds here informally named ‘Irishtown beds’) also are below this unconformity. The basal units of the Culpeper Basin (Seneca Creek Conglomerate, Reston Conglomerate, and Rapidan Member of the Manassas Sandstone) also probably lie below this unconformity, but definitive fossil evidence bracketing the age of the unconformity above these units is lacking. In the Newark Basin, Olsen & Rainforth (2003, fig. 4) have placed an unconformity between the Cutalossa and Prallsville members of the Stockton Formation. This also may represent the regional unconformity between TS II and TS III, and we also have noted that the Solebury Member of the Stockton Formation (which lies directly beneath the Prallsville Member) possibly belongs below this unconformity because of the occurrence there of an unusual amphibian not found

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Fig. 6. Map showing the names and distribution of basins comprising the Newark Supergroup in the eastern United States and eastern Canada. Basins are numbered from south to north. Basin numbers match basin column numbers in Figures 8 through 10.

anywhere above the TS II –TS III unconformity (Kozur & Weems 2007). Definitive conchostracan or other evidence for this conclusion, however, is lacking. Throughout the Newark Supergroup, Kozur & Weems (2005, 2007) have recognized a second

major stratigraphic unconformity, possibly as much as four or five million years in duration, which encompasses latest Norian and all but latest Rhaetian time (Fig. 7). This unconformity is located at the boundary between tectonostratigraphic sequences TS III and TS IV of Olsen (1997).

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Fig. 7. Graphic representation of the age ranges of the groups within the Newark Supergroup. Stratigraphy after Weems & Olsen (1997). CAMP is an acronym for Central Atlantic Magmatic Province.

Although Olsen did not recognize an unconformity at this boundary, he did note ‘sediments interbedded and overlying the basalts typically have much higher sedimentation rates than underlying sequences. . .’ As the TS I –TS II and TS II– TS III boundaries are both marked by significant unconformities (Olsen 1997), and as there is a sudden major increase in the rate of deposition with the beginning of sequence TS IV as described above, it is not surprising to find that an unconformity also marks the TS III –TS IV boundary. In most basins, modern erosional rubble from the basal lava flows has

obscured this contact. However, in the exposure near Old Wife Rock in the Fundy Basin of Nova Scotia, this low angle unconformity is prominently exposed at the top of a thick whitish paleosol that formed during the interval of nondeposition before the accumulation of the North Mountain Basalt (shown on the cover of Olsen et al. 1989). An angular unconformity also has been recognized in the Deerfield Basin between the Upper Triassic Sugarloaf Arkose and the Lower Jurassic Fall River beds (Hubert & Dutcher 1999), and it also is inferentially present in at least the northern Hartford

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Basin because the basal Talcott Basalt is missing there as it is in the Deerfield Basin to the north (lost within this unconformity). In the Newark and Culpeper basins, the extremely condensed Rhaetian section indicates that an unconformity is present, but sufficient outcrops to demonstrate this conclusively have not been found. The stratigraphy of those basins that have produced conchostracans is briefly summarized below. Numbers refer to the basin number designations that are used in Figures 6, 8 –10. We have made some fairly significant changes in interpretation since Kozur & Weems (2007), and these are discussed below.

Sanford and Durham Sub-basins of the Deep River Basin (4, 5) The first major work on the Deep River Basin was done by Emmons (1856), who referred to the entire stratigraphic sequence as the ‘Chatham series’ but did not erect a more detailed stratigraphy for units within any of its constituent sub-basins. Campbell & Kimball (1923) later established names for units within the middle Sanford Subbasin, and these were mapped in considerable detail by Reinemund (1955). The units recognized were (from oldest to youngest) the Pekin, Cumnock and Sanford formations. The Pekin and Sanford formations are lithologically similar redbed formations, but they are readily distinguished in the Sanford Sub-basin by the presence of the intervening Cumnock Formation, which contains abundant brown, tan, grey, green, and black shales and coals that formed in paludal to lacustrine depositional environments. This is the interval from which Bock (1953a, b) described Howellisaura berryi (Bock). Unfortunately, the Cumnock lithology only can be traced into the southernmost Durham Sub-basin. Therefore, the central and northern depositional sequences in that sub-basin are not so readily divided because they mostly are redbeds from bottom to top. Hoffman & Gallagher (1989) divided the stratigraphic sequence in the Durham Sub-basin into three successive lithofacies associations, of which their ‘lithofacies association I’ corresponds rather closely to all of the Pekin Formation and most of the Cumnock Formation. Their ‘lithofacies association II,’ well exposed in the Triangle Brick Quarry pit, in the past has been correlated with the lower part of the Sanford Formation (e.g. Huber et al. 1993a, fig. 8; Lucas et al. 1998). However, the conchostracan fauna preserved within this red, but lacustrine, stratigraphic interval is virtually identical with the conchostracan fauna that occurs near the top of the Cumnock Formation in the Sanford Sub-basin, indicating that ‘lithofacies association II’ instead is correlative with the

uppermost part of the Cumnock Formation. This leaves ‘lithofacies association III’ as being correlative with all of the Sanford Formation. Kozur & Weems (2007) placed the age of ‘lithofacies association II’ at the base of the Norian, citing the occurrence of the reptile Aetosaurus in the Triangle Brick Quarry pit which Lucas et al. (1998) suggested appeared no earlier than the base of the Norian. An earlier origin for Aetosaurus has been suggested based on phylogenetic evidence presented by Sues et al. (2003), but this has been disputed (Lucas 2010). The recent report of dicynodont bones from this pit (Peyer et al. 2008) in the past would have added a Carnian flavor to the fauna found there, but this is no longer the case as a Norian (or perhaps even Rhaetian) dicynodont has been documented recently in Poland (Dzik et al. 2008). Thus, based only on evidence from vertebrates, ‘lithofacies association II’ is best interpreted as representing early Norian (Lacian) time. The conchostracans from the Triangle Brick Quarry provide an even more definitive age for this unit. They are species characteristic of the Euestheria buravasi–Euestheria n. sp. Zone, which occur in the Germanic Basin in the lower Norian lower Arnstadt Formation (lower Steinmergelkeuper). This fauna is quite different from the fauna found in the uppermost Carnian, which includes Laxitextella freybergi Kelber & Kozur and other late Tuvalian guide forms that also occur in the uppermost Carnian of the Newark Supergroup (e.g. the Fulton site in the Gettysburg Basin, Fig. 8). Below the Euestheria buravasi– Euestheria n. sp. Zone, there is in the Germanic Basin a very short interval in the basal Norian that is characterized by a monospecific fauna of Palaeolimnadia schwanbergensis Reible. This fauna occurs also in the basal Norian Warford Member (basal Passaic Formation) of the Newark Basin. In southeastern Asia, Euestheria buravasi Kobayashi also occurs in the lower Norian. Thus, the conchostracan fauna from the Triangle Brick Quarry section belongs to the lower Norian close to, but a little above, the Carnian– Norian boundary in agreement with Lucas et al. (1998).

Dan River/Danville Basin (7) Early reconnaissance mapping suggested that there were two different basins in this area, but subsequent detailed mapping demonstrated that only a single elongate basin is present. The location of this basin, astride the North Carolina–Virginia border, resulted in the creation of two different stratigraphies for the constituent strata. Meyertons (1963), working in the northern Danville (Virginia) end of the basin, recognized a basal Dry Fork Formation that generally is overlain by, but partly

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Fig. 8. Correlation chart of the lower strata in the southern and central part of the Newark Supergroup, showing their stratigraphy and the biostratigraphically important horizons in each basin.

intertongues with, a Leakesville Formation, which in turn is overlain disconformably by a Cedar Forest Formation. Within the Leakesville, Meyertons recognized two intertonguing members: the Cow Branch Member and the Cascade Station Member.

Thayer (1970), mapping in the southern Dan River (North Carolina) end of the basin, erected a distinctly different stratigraphy. Thayer named his basal coarse and mostly fluvial unit the Pine Hall Formation, which is overlain gradationally by his

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Fig. 9. Correlation chart of the strata in the Dan River/Danville, Culpeper, and Gettysburg basins in the west– central part of the Newark Supergroup, showing their stratigraphy and the biostratigraphically important horizons in each basin.

finer and mostly lacustrine Cow Branch Formation. This in turn is overlain by, and partly intertongues with, his Stoneville Formation. More recently, Kent & Olsen (1997) have published a measured section from the base of the lower Cow Branch Formation into the lower part of the Cascade Station

Member that considerably increases the reported thickness of this interval. Aspects of all of these stratigraphies have been adopted here (Fig. 9). Conchostracans have been found so far only in the Cow Branch Formation at two different localities, one near Pine Hall in the ‘lower Cow Branch’ and

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Fig. 10. Correlation chart of the strata in the Newark, Hartford, and Fundy basins in the northern part of the Newark Supergroup, showing their stratigraphy and the biostratigraphically important horizons in each basin.

the other at the Solite Quarry in the ‘upper Cow Branch’ (Kozur & Weems 2007).

Scottsburg, Briery Creek, and Farmville basins (8, 11, 12) No formal lithostratigraphic names have been erected for the strata in any of the basins that lie along the trend of the Farmville Basin. Two basins (Randolph and Roanoke Creek) have yielded no

lacustrine strata so far, but the other three have paludal to lacustrine strata that contain the fish Dictyopyge (in the Scottsburg Basin) and conchostracans (in the Briery Creek and Farmville basins). The conchostracans and fish are all characteristic of the lower Cordevolian Laxitextella multireticulata Zone. These basins probably represent remnants of the ‘keel’ of a once continuous and much more extensive basin along this trend. The most completely-preserved stratigraphic section is

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in the Farmville Basin, which has a basal sandstone unit, an intermediate paludal/lacustrine unit that contains conchostracans and fish, and a disconformably overlying fluvial to fanglomeratic unit. This sequence is very reminiscent of the stratigraphy found in the Richmond and Taylorsville basins. Wilkes (1982, 1987) and Goodwin et al. (1986) have provided general mapping of these basins. Detailed mapping has been published only for the northern part of the Farmville Basin (Marr 1980).

Richmond and Deep Run basins (13, 15) The Deep Run Basin is a small, shallow eastern structural outlier of the Richmond Basin. It is separated from the main basin by only a few kilometers and has a stratigraphy identical to the basal part of the main basin (Shaler & Woodworth 1899). The stratigraphy of the Richmond Basin was established by Shaler & Woodworth (1899) and has been revised since by Cornet & Olsen (1990). The interpretation of Cornet & Olsen (1990) is complex and based in large part on interpretation of seismic lines. In the absence of any surface control, we choose to largely follow the original stratigraphy of Shaler & Woodworth (1899), in which the basal unit of the basin is the Tuckahoe Formation (‘Tuckahoe Group’ in Shaler & Woodworth 1899; reduced to Tuckahoe Formation by Cornet & Olsen 1990), the middle unit is the Vinita Shale, and the upper unit is the Otterdale Sandstone. In this stratigraphy, the productive coal beds constitute the upper member of the Tuckahoe Formation. The coal bed interval is the horizon from which Bock (1953a, b) described Isaura midlothianensis (Euestheria sp. indet. of E. minuta group) and Howellisaura winterpockensis (¼ Euestheria winterpockensis).

Taylorsville Basin (16) The stratigraphy of the exposed Taylorsville Basin was described and mapped by Weems (1980, 1981, 1986). LeTourneau (2003) revised the stratigraphy of Weems and added several new formations for subsurface stratigraphic units that are higher than any exposed at the surface. Kozur & Weems (2007) rejected most of LeTourneau’s revisions of the exposed stratigraphy. The only unit in the exposed portion of the Taylorsville Basin that has yielded conchostracans is the paludal to lacustrine Falling Creek Formation. The conformably underlying Stagg Creek Formation and the disconformably overlying Newfound Formation both formed in fluvial depositional environments that were not favorable for conchostracan colonization. Recently, Paul Olsen (Columbia University) has made available to the authors conchostracans from the

subsurface Port Royal Formation that were mentioned in LeTourneau (2003) but not described or illustrated. These specimens are identical with the fauna found in the upper (but not uppermost) Lockatong Formation of New Jersey, and they confirm the estimated stratigraphic position assumed for this unit in Kozur & Weems (2007) (see Fig. 8 for approximate sample horizon).

Culpeper Basin (19) The stratigraphy of the Culpeper Basin has gone through a number of iterations. Roberts (1928) recognized a basal Manassas Sandstone overlain by a Bull Run Shale, but he did not describe the higher portions of the stratigraphic column. Lindholm (1979) provided the first inclusive stratigraphic compilation and applied stratigraphic names to the entire column. His work subsequently was revised and partially duplicated by Lee & Froelich (1989). Weems & Olsen (1997) synthesized the best aspects of both stratigraphies, and their stratigraphy is used here (Figs 8 & 9). The lowest formation in the basin is the Manassas Sandstone, which at its base includes the Tuscarora Creek Member in the northern part of the basin, the laterally equivalent Reston Conglomerate Member in the central part of the basin, and the Rapidan Member in the southern part of the basin. Unconformably above all of these is an upper Poolesville Sandstone Member. The Manassas is overlain by the Bull Run Formation, which consists of a lower Balls Bluff Siltstone Member and an upper Groveton Member. The Bull Run is overlain successively by the Catharpin Creek Formation, the Mount Zion Church Basalt, the Midland Formation, the Hickory Grove Basalt, the Turkey Run Formation, the Sander Basalt, and the Waterfall Formation. The Manassas Formation is composed of fluvial sandstones and siltstones that so far have yielded no conchostracans. Much of the overlying Balls Bluff siltstone of the Bull Run Formation also represents fluvial deposits, though they are quite fine-grained and reflect much lower energy conditions than the Manassas. One locality within the Balls Bluff (the Rutiodon and coelacanth locality described in Weems, 1979 and in Weems & Kimmel 1993) has yielded impressions that may be conchostracans but this is not definitive and the specimens, if real, are not well enough preserved to be useful for stratigraphic correlation. Above the Balls Bluff, the Groveton Member of the Bull Run Formation and all but the uppermost Catharpin Creek Formation have yielded many abundant conchostracan faunas of Norian age. These have been described in some detail in Kozur & Weems (2005, 2007) and can be assigned to the Shipingia hebaozhaiensis Zone, the Redondestheria grovetonensis Zone, and the

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Shipingia olseni zones of Alaunian and Sevatian (middle and late Norian) age. The highest beds of the Catharpin Creek Formation in the Culpeper Basin, about 15 m below the lowest Mt. Zion Church Basalt, have recently yielded a fauna consisting entirely of Euestheria brodieana, indicating its late (and probably latest) Rhaetian age. Only about 10 m below this, beyond a covered interval, a rich fauna dominated by Shipingia olseni demonstrates that this bed and beds below it are late Norian and older. Above the Mt. Zion Church Basalt, a fish-bearing unit in the middle part of the Midland Formation (the ‘Midland fish bed’) has yielded a more varied conchostracan fauna that includes E. brodieana, Bulbilimnadia killianorum (named in this paper) and very rare Bulbilimnadia sheni. The latter two species are completely unknown from the Rhaetian (even the uppermost Rhaetian) in the Germanic Basin region, but B. killianorum recently has been found by Grzegorz Niedz´wiedzki and Grzegorz Pien´kowski, Warsaw, in the lowermost Hettangian continental strata of the Holy Cross Mountains (Poland). Therefore, the age of the Mt. Zion Church Basalt is just before (or possibly at) the Triassic –Jurassic boundary, the age of the lowest Midland Formation is unknown, and the middle and upper Midland are early Hettangian in age. In the middle Midland, E. brodieana and B. killianorum are by far the more abundant species, with B. sheni occurring only very rarely. By the upper part of the Midland, B. sheni is distinctly more common than B. killianorum and E. brodieana remains abundant. By the base of the Waterfall Formation, B. killianorum is gone and the fauna consists almost entirely of E. brodieana and B. sheni. About 800 m above the base of the Waterfall, B. froelichi makes its first appearance, and by the top of the Waterfall Formation E. brodieana and B. sheni both are gone. Thus, a rather rapid series of changes in the conchostracan fauna can be documented in the upper part of the Culpeper Basin column. As discussed in Kozur & Weems (2007), a rich conchostracan fauna in the Yale Peabody Museum collection, allegedly from the Turkey Run Formation, almost certainly is from the basal Waterfall Formation. So far, the Turkey Run Formation has yielded no conchostracans.

Gettysburg Basin (20) The stratigraphy of the southeastern Pennsylvania Gettysburg Basin was defined by Stose & Bascom (1929), and this stratigraphy later was extended into the Maryland part of the basin by Jonas & Stose (1938). These authors recognized two formations, a basal, predominantly fluvial New Oxford Formation and an overlying, predominantly

343

lacustrine Gettysburg Shale. This unit was later renamed the Gettysburg Formation (Glaeser 1963) to more accurately reflect the presence of the many facies within the unit that were much coarser than shale. Within the Gettysburg, Stose & Bascom (1929) also distinguished a ridge-forming interval that they named the Heidlersburg Member. They considered this interval to be sandstone-rich, but Smoot (1999) has pointed out that it is not much different from other parts of the Gettysburg Formation in its overall lithology. Cornet (1977) recognized two thin units at the very top of the Gettysburg Basin fill, a basalt that he informally called the Aspers basalt, and sedimentary strata above it. Weems & Olsen (1997) formalized this name as the Aspers Basalt and named the strata above the basalt the Bendersville Formation. Based on our conchostracan studies (Kozur & Weems 2007), it is now clear that the basal sequence of the New Oxford Formation is distinctly older than the strata that overlie them and is separated from the overlying strata by a disconformity, identifiable in the stratigraphic column shown in Stose & Bascom (1929) by a prominent quartzose conglomerate that lies 35 m above the basal unconformity. Stose & Bascom gave a detailed section of this sequence in an outcrop near Irishtown, and because of this these beds are here informally called the ‘Irishtown beds.’ In Kozur & Weems (2007) they were designated simply as an unnamed unit below the New Oxford Formation. A few horizons within the Gettysburg Basin have yielded abundant conchostracans, and it seems likely that many more are to be found. One long-recognized conchostracan-bearing interval, considered in the past to lie within the New Oxford Formation, is a lacustrine sequence that occurs along Little Conewago Creek in Pennsylvania just south of the Susquehanna River. The geological map of York County, Pennsylvania (Stose & Jonas 1939) shows this locality as being near the middle of the New Oxford outcrop belt, and this was used by Cornet (1977) to estimate a relatively low horizon for this locality. However, Wanner (1926) estimated that this locality is nearly 2000 metres above the base of the New Oxford Formation, estimating strata thickness from an average regional 208 dip in this area. This is very nearly the full thickness of the New Oxford in its type area as measured by Stose & Bascom (1929). Stose & Jonas (1939) show a major fault immediately northwest of the Little Conewago Creek area within the New Oxford strike belt, seemingly with its west side up, so the New Oxford Formation probably is largely fault-repeated within this area. Thus, the stratigraphic level calculated by Wanner is probably close to correct, and the lake bed strata found along Little Conewago Creek probably are the basal lake beds of the Gettysburg

344

H. W. KOZUR & R. E. WEEMS

Formation and not an anomalous lake sequence within the normally fluvial New Oxford Formation. This interval has produced remains of phytosaurs (Rutiodon carolinensis and possibly Rutiodon manhattanensis, described in Doyle & Sues 1995), metoposaurs (Buettneria perfecta), fish, and conchostracans including abundant specimens of ‘Estheria’ pennsylvanica Wanner. This is the type area for ‘E.’ pennsylvanica Wanner, 1926, and the authors have collected numerous new specimens from this area and reassigned this species to a new genus (Wannerestheria n. gen.) described in the taxonomic portion of this paper. Comparison with conchostracan collections from a number of horizons within the Newark Basin in the Delaware River Valley, made by the authors and by Paul Olsen (Columbia University), indicates that Wannerestheria pennsylvanica occurs only in the upper part of the Lockatong Formation. Therefore, the lacustrine strata along Little Conewago Creek correlate with the upper part of the Lockatong, which is upper Tuvalian (uppermost Carnian) in age. Therefore, the Gettysburg Formation correlates with both the uppermost Lockatong and all of the Passaic Formation of the Newark Basin. A second conchostracan-bearing interval, the Fulton site in Maryland (Kozur & Weems 2007), lies well to the southwest of the type area of the New Oxford and Gettysburg formations and also beyond several structural anomalies within the basin. These structural complications make it impossible to directly trace the Maryland lacustrine interval in a simple manner across the intervening region to the type area of the New Oxford and Gettysburg formations, and this has made its precise stratigraphic horizon rather problematic. Jonas & Stose (1938) mapped this horizon in Maryland as the basal interval of the Gettysburg Formation, and its lacustrine character is certainly typical of the Gettysburg. Conchostracans found at and near the Fulton site constitute a diverse upper Tuvalian fauna that can be readily correlated with the upper Tuvalian Laxitextella freybergi Zone of the Germanic Basin. Even the index species L. freybergi Kozur is present. At the same time, however, the more endemic form Wannerestheria pennsylvanica (Wanner) is present. Therefore the Fulton site seems to be at (or at least very close to) the same horizon as the type locality of Wannerestheria pennsylvanica along Little Conewago Creek. This supports the assignment of the Fulton site lake beds to the base of the Gettysburg Formation (see Figs 8 & 9).

Newark Basin (22) The strata of the New Jersey portion of the Newark Basin were divided into three formations by

Kummel (1897), and these units were later carried over into the Pennsylvania portion of the Newark Basin by Bascom et al. (1909). In ascending order, these units were a predominantly fluvial Stockton Formation, a lacustrine Lockatong Formation, and a predominantly redbed Brunswick Formation. The Brunswick Formation included the bulk of the stratigraphic column of the Newark Basin, and this interval later was divided into seven formations by Olsen (1980), who, in doing so, abandoned the name ‘Brunswick.’ The interval represented by the former Brunswick Formation now is occupied by (in ascending order) the Passaic Formation, the Orange Mountain Basalt, the Feltville Formation, the Preakness Basalt, the Towaco Formation, the Hook Mountain Basalt and the Boonton Formation (Fig. 10). The Triassic part of the column (Stockton, Lockatong and Passaic formations) since then has been subdivided into a considerable number of members (Olsen et al. 1999). Conchostracans occur at many levels in the Newark Basin, and much work remains to be done on the details of its conchostracan biostratigraphy. One fauna is now known from the upper part of the Stockton Formation. It is yet to be studied in detail, but it is obviously different from overlying faunas and belongs to a different conchostracan biozone than the biozone of the basal Lockatong (Kozur & Weems 2007). A number of horizons have been collected from the Lockatong Formation, and it now is known to include three successive faunas that can be readily distinguished from each other. The conchostracan fauna of the upper Passaic Formation belongs within the Shipingia olseni Zone (Kozur & Weems 2007), but more recently discovered faunas from much lower in the Passaic, yet to be described, clearly belong below the upper Alaunian S. hebaozhaiensis Zone. A conchostracan fauna from the Warford Member of the Passaic, toward the base of the formation, is clearly early Norian (‘Lacian’) in age. As conchostracan faunas from the upper (but not uppermost) part of the underlying Lockatong Formation (Smith Corner Member) are still of late Tuvalian (uppermost Carnian) age, the Carnian –Norian boundary lies somewhere between the Smith Corner Member of the Lockatong and the Warford Member of the Passaic. A distinct change in the regional climate has been documented in the Germanic Basin at the Carnian– Norian boundary (Kozur & Bachman 2010), and it may be that the Lockatong –Passaic boundary (which also reflects a major climatic shift) similarly corresponds to the Carnian –Norian boundary. Until conchostracan faunas are recovered from strata between the Smith Corner and Warford members, no more precise location for the Carnian –Norian boundary can be established in the Newark Basin. However,

TRIASSIC CONCHOSTRACANS

the monospecific Palaeolimnadia schwanbergensis conchostracan fauna of the Warford Member is identical with the fauna from the basal Weser Formation (basal Norian) of the Germanic Basin. As the oldest Newark Basin occurrence of the Norian genus Aetosaurus also is in the Warford Member (¼ base of the Neshanician ¼ base of Revueltian), this level may well represent the basal Norian as assumed by Kent & Olsen (2000). Conchostracans apparently are rare in the Lower Jurassic part of the Newark Basin column. A few poorly preserved specimens, possibly but not certainly belonging to Bulbilimnadia froelichi, have been identified on a slab containing fossil fish from a horizon high in the Boonton Formation (see locality Table 1). This suggests that the Boonton Formation does not range higher than the lowest part of the Portland Formation in the Hartford Basin, which is consistent with the much thinner section of the Boonton as compared with the Portland.

Hartford Basin (25) Krynine (1950) applied the first formal stratigraphy to the sediments of the Hartford Basin. In ascending order, he recognized a predominantly fluvial New Haven Arkose, a Meriden Formation that consisted of interbedded lava flows and fluvial to lacustrine strata, and a Portland Formation that is predominantly lacustrine in its lower part but fluvial toward its top. Rodgers et al. (1959) named the lava flows as members of the Meriden, and called them (in ascending order) the Talcott Lava Member, the Holyoke Lava Member, and the Hampden Lava Member. Lehman (1959) slightly modified these names and additionally named the sedimentary packages between them as formations. In ascending order these were the Talcott Basalt, the Shuttle Meadow Formation, the Holyoke Basalt, the East Berlin Formation and the Hampden Basalt (Fig. 10). The name Meriden was abandoned, but since has been revived as a group name by Weems & Olsen (1997). Fossils are rare in the New Haven Arkose, but skeletal remains of Rutiodon and Stegomus (Aetosaurus) clearly indicate a late Carnian to Norian age for that unit (Lull 1953). Cornet (1977, p. 225) reported a lower Norian palynoflora from near the base of the New Haven, which suggests that deposition in the Hartford Basin did not begin until after the end of Carnian time, much later than in any other Newark Supergroup basin. Overlying units are much more fossiliferous (McDonald 1996). We have examined conchostracans from only two horizons, one from the middle of the East Berlin Formation and the other from the middle of the Portland Formation, but other

345

horizons have been reported (McDonald 1996). The specimens from the East Berlin Formation are not well-preserved but seem to be Bulbilimnadia froelichi, which was first described from the Waterfall Formation in the Culpeper Basin (Kozur & Weems 2005). The fauna from the middle of the Portland Formation seems to be composed of new species, more similar to forms found in the Kayenta Formation of the southwestern United States than to forms from lower horizons within the Newark Supergroup. Notably, B. froelichi seems to have disappeared before the middle Portland Formation, which suggests that the middle Portland is upper Hettangian in age.

Fundy Basin (29) Powers (1916) named strata in the Fundy Basin (in ascending order) the Annapolis Formation (consisting of a lower Wolfville Sandstone member and an upper Blomidon Shale member), the North Mountain Basalt, and the Scots Bay Formation. Later, Klein (1962) abandoned the name Annapolis Formation and raised the Wolfville and Blomidon to formational rank (Fig. 10). Beds immediately beneath the Wolfville Formation, called the Lower Economy beds, have been ascribed a Middle Triassic age (Baird & Olsen 1983). Because they are considerably older than any other beds usually included in the Newark Supergroup, they generally have been excluded from the definition of the Newark Supergroup (e.g. Weems & Olsen 1997). Baird & Olsen (1983) discussed a vertebrate fauna from the middle of the Wolfville Formation, which they assigned to the Carnian, and a vertebrate fauna from the upper Wolfville that they assigned to the Carnian –Norian. Conchostracans found in the lower part of the Wolfville Formation near Evangeline Beach constitute a monospecific Euestheria minuta (von Zieten) fauna of probable late Ladinian age, slightly older than other any conchostracan fauna so far known from anywhere else in the Newark Supergroup. Since this interval is close in age to the lower Cordevolian units in the United States, the question arises if there might be an upper Cordevolian–lower Julian unconformity separating these strata from higher strata in the Wolfville just as there is farther south. Unfortunately, the Evangeline Beach section lies across the Southern Bight of the Minas Basin from the type Wolfville area so, in the absence of any known outcrops of the intervening stratigraphic interval, it is not possible to determine if this is true or not. Conchostracans have been reported (but not illustrated) from the upper part of the type area of the Wolfville Formation at Medford Beach near Paddy Island (Cameron & Gould 2000); their described size is much too large to be E. minuta, so even without

Locality

Age

Sanford Basin, North Carolina Pomona Pipe Co. pit near Early Tuvalian Gulf Bethany Church Late early-middle Tuvalian Late early-middle Tuvalian

Carbonton Dam site

Upper Tuvalian

River Road near Horseshoe Bend

Lacian

Stop 1.3 in Olsen et al. (1989)

Lacian

Durham Basin, North Carolina Lacian Triangle Brick Quarry, NMMNH and NC Museums Dan River Basin, North Carolina Solite Quarry Late Tuvalian U.S. 220 road cut on N side of Dan River

Early Tuvalian

Briery Creek Basin, Virginia Worthy property (YPM Early Cordevolian 34659 and YPM 220198) Flournoy’s coal pit *(YPM 34664) Farmville Basin, Virginia Little Willis River (YPM 34647)

Early Cordevolian

Early Cordevolian

Unit

Latitude

Longitude

Pekin Formation, near middle of unit Cumnock Formation, about 80 m above the base Cumnock Formation, about 100– 200 m above base Cumnock Formation, near middle of unit Cumnock Formation, near top of unit

35.56948N

79.29768W

35.55848N

79.29218W

Not reported

Other data

Conchostracan Zone

Goldston Quadrangle, Chatham Co., N.C. Goldston Quadrangle, Chatham Co., N.C.

Zone indeterminate

Not reported

Goldston Quadrangle?, Chatham Co., N.C.

Howellisaura princetonensis Zone

35.52018N

79.34808W

35.45538N

79.38458W

Goldston Quadrangle, Chatham County, N.C. Goldston Quadrangle, Moore County, N.C.

Sanford Formation, near bottom of unit

35.50818N

79.30048W

Putnam Quadrangle, Lee County, N.C.

Anyuanestheria wingatella Zone Euestheria buravasi– Euestheria n. sp. Zone Norestheria barnaschi– Shipingia mcdonaldi Zone

‘Lithofacies Association II’ of Hoffman & Gallagher 1989

35.86948N

78.89608W

Green Level Quadrangle, Durham County, N.C.

Euestheria buravasi– Euestheria n. sp. Zone

Upper Cow Branch Formation, about 100 m below top? Lower Cow Branch Formation, about 10 m above the base

36.54128N

79.66988W

Northeast Eden Quadrangle, Rockingham County, N.C.

Zone indeterminate

36.38758N

79.94118W

Mayodan Quadrangle, Rockingham County, N.C.

New zone just below A. wingatella Zone

Lower lacustrine unit, about 100 m above base Lower lacustrine unit, about 100 m above base

37.21908N

78.44458W

Hampden Sydney Quadrangle, Prince Edward County, Va.

Laxitextella multireticulata Zone

37.19498N

78.47548W

Hampden Sydney Quadrangle, Prince Edward County Va.

Laxitextella multireticulata Zone

Within unnamed lacustrine unit about 450 m above its base

37.41138N

78.39408W

Willis Mountain Quadrangle, Cumberland Co., Va.

Laxitextella multireticulata Zone

Howellisaura princetonensis Zone

H. W. KOZUR & R. E. WEEMS

Deep River coal basin (ANSP 31194)

346

Table 1. Conchostracan localities in the Newark Supergroup

Richmond Basin, Virginia Midlothian coal mine spoil (ANSP 31497)

Early Cordevolian

Winterpock (ANSP 31197)

Early Cordevolian

Deep Run Basin, Virginia Three Chopt Road fish locality

37.50298N

77.64028W

Midlothian Quadrangle, Chesterfield Co., Va.

Laxitextella multireticulata Zone

37.34598N

77.72048W

Winterpock Quadrangle, Chesterfield Co., Va.

Laxitextella multireticulata Zone

Early Cordevolian

Lower Vinita Shale, about 50 m above base

37.63518N

77.58168W

Glen Allen Quadrangle, Henrico Co., Va.

Laxitextella multireticulata Zone

Taylorsville Basin, Virginia Stagg Creek, older level

Early Cordevolian

37.78318N

77.54508W

Stagg Creek, younger level

Early Cordevolian

37.78338N

77.54508W

Little River site, W of C & O RR bridge

Early Cordevolian

37.82258N

77.42858W

Hanover Academy Quadrangle, Hanover Co., Va. Hanover Academy Quadrangle, Hanover Co., Va. Ashland Quadrangle, Hanover Co., Va.

Laxitextella multireticulata Zone Laxitextella multireticulata Zone Laxitextella multireticulata Zone

Falling Creek, north of Ashland

Late Cordevolian

37.78828N

77.48558W

Ashland Quadrangle, Hanover Co., Va.

Laxitextella multireticulata Zone

Wilkins Well (in LeTourneau 2003)

Tuvalian

Lower Falling Creek Fm., 40 m above base Lower Falling Creek Fm., 44 m above base Middle Falling Creek Fm., about 155 m above base Upper Falling Creek Fm., near top (exact horizon uncertain) Port Royal Fm., horizon not specified

38.24148N

77.01258W

Dahlgren Quadrangle, Westmoreland Co., Va.

Howellisaura ? ovata Zone

Groveton Mbr., Bull Run Fm., 350 m above base Groveton Mbr., Bull Run Fm., 1250 m above base Groveton Mbr., Bull Run Fm., 1850 m above base Groveton Mbr., Bull Run Fm., 2310 m above base Groveton Mbr., Bull Run Fm., 2550 m above base Groveton Mbr., Bull Run Fm., 2750 m above base

38.64208N

77.59288W

38.80148N

77.52648W

Nokesville Quadrangle, Fauquier Co., Va. Gainesville Quadrangle, Prince William Co., Va.

Shipingia hebaozhaiensis Zone Shipingia hebaozhaiensis Zone

38.85128N

77.52158W

Gainesville Quadrangle, Fairfax Co., Va.

Shipingia hebaozhaiensis Zone

38.81368N

77.54608W

Gainesville Quadrangle, Prince William Co., Va.

Redondestheria grovetonensis Zone

38.94998N

77.52948W

Arcola Quadrangle, Loudoun Co., Va.

Redondestheria grovetonensis Zone

38.85998N

77.56738W

Gainesville Quadrangle, Loudoun Co., Va.

Shipingia olseni Zone

Culpeper Basin, Virginia Carriage Ford

Alaunian

I-66 at VA Route 234, NW corner

Alaunian

Sudley Road

Alaunian

Groveton Cemetery, Bull Run Battlefield

Sevatian

Arcola

Sevatian

Catharpin

Sevatian

347

(Continued)

TRIASSIC CONCHOSTRACANS

Lower Vinita Shale, probably within basal 50 m Lower Vinita Shale, probably within basal 50 m

348

Table 1. Continued Locality

Age Sevatian

Lenah

Sevatian

Cedar Run

Sevatian

Haymarket

Sevatian

Casanova

Sevatian

Locality HAYM of Cornet (1977)

Late Rhaetian

Midland fish bed on Licking Run Killian Property

Basal Hettangian

Avalon Farm

Early Hettangian

McDonald site on I 66 (YPM 202438 and YPM 202439)

Early Hettangian

Opal #1 core

Early Hettangian

Catletts Branch

Early Hettangian

Early Hettangian

Gettysburg Basin, Maryland and Pennsylvania Hanover Street, SE of town Early Cordevolian of New Oxford Rheems locality of De Wet et al. (1998)

Early Tuvalian or late Julian

Groveton Mbr., Bull Run Fm., 2900 m above base Catharpin Creek Fm., 200 m above base Catharpin Creek Fm., 1050 m above base Catharpin Creek Fm., 100 m below top and 2350 m above base Catharpin Creek Fm., 50 m below top and 2400 m above base Catharpin Creek Fm., 9.4 m below top and 2440 m above base Midland Fm., 100 m above base Midland Fm., 150 m above base Lower Waterfall Fm., 250 m above base Lower Waterfall Fm. About 300 m above base Upper Waterfall Fm., 800 m above base Upper Waterfall Fm., 150 m below top and 1100 m above base

Irishtown beds, below New Oxford Fm., about 10 m above base Basal Gettysburg Fm., about 30 m above base

Latitude

Longitude

Other data

Conchostracan Zone

38.96358N

77.55168W

Arcola Quadrangle, Loudoun Co., Va.

Shipingia olseni Zone

38.96128N

77.57008W

Shipingia olseni Zone

38.65308N

77.67688W

38.84818N

77.63528W

Arcola Quadrangle, Loudoun Co., Va. Catlett Quadrangle, Fauquier Co., Va. Thoroughfare Gap Quadrangle, Prince William Co., Va.

38.65228N

77.70448W

Catlett Quadrangle, Fauquier Co., Va.

Shipingia olseni Zone

38.81248N

77.64588W

Thoroughfare Gap Quadrangle, Prince William Co., Va.

Euestheria brodieana Zone

38.61728N

77.72538W

38.77348N

77.69638W

38.75518N

77.65978W

38.82198N

77.68558W

Midland Quadrangle, Fauquier Co., Va. Thoroughfare Gap Quadrangle, Fauquier Co., Va. Thoroughfare Gap Quadrangle, Prince William Co., Va. Thoroughfare Gap Quadrangle, Prince William Co., Va.

Upper Bulbilimnadia killianorum Zone Bulbilimnadia sheni Zone Bulbilimnadia sheni Zone Bulbilimnadia sheni Zone

38.63338N

77.80008W

38.82348N

77.69988W

Warrenton Quadrangle, Fauquier Co., Va. Thoroughfare Gap Quadrangle, Prince William Co., Va.

Bulbilimnadia froelichi Zone Bulbilimnadia froelichi Zone

39.84508N

77.03408W

McSherrystown Quadrangle, Adams Co., Pa.

Laxitextella multireticulata Zone

40.13158N

76.58318W

Elizabethtown Quadrangle, Lancaster Co., Pa.

Conchostracans not identifiable

Shipingia olseni Zone Shipingia olseni Zone

H. W. KOZUR & R. E. WEEMS

Broad Run west of power line

Unit

Locality 11 of Wanner (1926)

Late Tuvalian

Locality 1 of Wanner (1926)

Late Tuvalian

Fulton site on Beaver Branch

Late Tuvalian

U.S. Rt. 15 SSE of Heidlersburg

Late Alaunian

Black Rock Tunnel (ANSP 16849) Eureka quarry (ANSP 16847, ANSP 16848, ANSP 65528, ANSP 65530) Princeton Library

Middle Tuvalian Late early-middle Tuvalian Late early-middle Tuvalian Middle Tuvalian

Granton Quarry (ANSP 31192) Skunk Hollow

Late Tuvalian

Smith Corner

Late Tuvalian

Warford Creek

Lacian

Milford

Lacian

150 m east of Cornet (1979) localities GHT 1, 3 Exeter, Constitution Avenue 14

Sevatian Sevatian

40.07038N

76.73538W

Yorkhaven Quadrangle, York Co., Pa.

Wannerestheria pennsylvanica Zone

40.09328N

76.72648W

York Haven Quadrangle, York Co., Pa.

Wannerestheria pennsylvanica Zone

39.63558N

77.33478W

39.93938N

77.14328W

Emmitsburg Quadrangle, Frederick Co., Md. Biglerville Quadrangle, Adams Co., Pa.

Laxitextella freybergi Zone Norestheria barnaschi-Shipingia mcdonaldi Zone

Upper Stockton Fm., about 1050 m above base Lockatong Fm., about 200 m above base Lockatong Fm., stratigraphic level unknown Lockatong Fm., about 200 m above base Lockatong Fm., about 300 m above base Lockatong Fm., about 600 m above base

40.41148N

75.03058W

Lumberville Quadrangle, Hunterdon Co., N.J.

Zone not yet established

40.14388N

75.51298W

40.26028N

75.18158W

Phoenixville Quadrangle, Chester Co., Pa. Doylestown Quadrangle, Montgomery Co., Pa.

Howellisaura ? ovata Zone Howellisaura princetonensis Zone

40.34978N

74.65788W

40.84478N

73.98398W

Not reported

Not reported

Howellisaura princetonensis Zone Howellisaura ? ovata Zone Wannerestheria pennsylvanica Zone

40.45448N

75.06758W

40.46948N

75.06118W

Princeton Quadrangle, Mercer Co., N.J. Central Park Quadrangle, Bergen Co., N.J. Material provided by Paul Olsen, no specific locality given Lumberville Quadrangle, Hunterdon Co., N.J. Lumberville Quadrangle, Hunterdon Co., N.J.

40.57258N

75.11238W

40.31858N

75.78508W

40.31348N

75.84358W

Lockatong Fm., about 750 m above base Passaic Fm., Warford Mbr., about 200 m above base Passaic Fm., about 300 m above base Passaic Fm., 1750 m from top Passaic Fm., 25 m from top

Frenchtown Quadrangle, Hunterdon Co., N .J. Birdsboro Quadrangle, Berks Co., Pa. Birdsboro Quadrangle, Berks Co., Pa.

TRIASSIC CONCHOSTRACANS

Newark Basin, Pennsylvania and New Jersey Raven Rock Quarry Early Tuvalian

Basal Gettysburg Fm., about 1650 m above base Basal Gettysburg Fm., about 1700 m above base Gettysburg Fm., basal lake bed Gettysburg Fm., lower Heidlersburg Mbr., about 200 m above base

Wannerestheria pennsylvanica Zone Palaeolimnadia schwanbergensis Zone Zone indeterminate Shipingia olseni Zone Shipingia olseni Zone (Continued) 349

350

Table 1. Continued Locality Fish bed at Boonton dam (specimens on fish slab YPM 12349)

Age Early Hettangian

North Branford

Early Hettangian

Miner Brook (Westfield fish bed) Kelsey– Ferguson Quarry

Early Hettangian Late Hettangian

Fundy Basin, Nova Scotia (Canada) Evangeline Beach Late Ladinian Medford Beach locality of Cameron & Gould (2000) Blomidon Cliff (YPM 34664) Wasson Bluff, Station 11 of Olsen and Et-Touhami (2008)

Carnian Sevatian ?Earliest Hettangian

Latitude

Longitude

Other data

Conchostracan Zone

Boonton Fm., about 400 m above base

40.89588N

74.39788W

Boonton Quadrangle, Morris Co., N.J.

Bulbilimnadia froelichi Zone

Shuttle Meadow Fm., About 50 m above base Middle East Berlin Fm., about 80 m below top Middle East Berlin Fm., 82 m below top Middle Portland Fm., about 1500 m above base

Not reported

Not reported

Durham Quadrangle, Middlesex Co., Conn.

Material not examined

41.33568N

72.76868W

41.59188N

72.70068W

41.98068N

72.60788W

Branford Quadrangle, New Haven Co., Conn. Middletown Quadrangle, Middlesex Co., Conn. Broad Brook Quadrangle, Hartford Co., Conn.

Bulbilimnadia froelichi Zone Bulbilimnadia froelichi Zone Zone not yet defined

45.13898N

64.32938W

Euestheria minuta Zone

45.18388N

64.35758W

45.23438N

64.35468W

Approximately 45.40008N

Approximately 64.25008W

Wolfville Quadrangle (21 H/ 1), Kings Co., Nova Scotia Wolfville Quadrangle (21 H/ 1), Kings Co., Nova Scotia Wolfville Quadrangle (21 H/ 1), Kings Co., Nova Scotia Parrsboro Quadrangle (21 H/8), Cumberland Co., Nova Scotia

Wolfville Fm., about 350 m above base Wolfville Fm., about 1,10 m above base Blomidon Fm., about 1000 m above base McCoy Brook Fm., near base of unit

Zone indeterminate Shipingia olseni Zone Material not examined

H. W. KOZUR & R. E. WEEMS

Hartford Basin, Connecticut Durham fish bed of ?Earliest McDonald (1992) Hettangian

Unit

TRIASSIC CONCHOSTRACANS

better identification they still must be at least Carnian in age. Conchostracans from the Blomidon Formation also have been reported but not illustrated by Cameron & Ford (1998). There are specimens, however, that we have seen and studied from high in the Blomidon Formation at Cape Blomidon. These demonstrate that the upper part of the Blomidon belongs within the Shipingia olseni Zone and thus is upper Norian (Sevatian) in age (Kozur & Weems 2007). This fauna also includes a new and so far endemic genus and species, Acadiestheriella cameroni n. gen. n. sp., described in the present paper. No conchostracan data are available yet from the highest (Scots Bay) strata above the North Mountain Basalt, but the basal Scots Bay has recently produced Rhaetian palynomorphs (Cirilli et al. 2009). Based on the early Hettangian fauna now known from the middle of the stratigraphically equivalent Midland Formation in the Culpeper Basin, the age of this unit may range upward at higher levels into the early Hettangian. Olsen et al. (2003) created the name Partridge Island Member of the Blomidon Formation for a thin sequence of beds lying immediately beneath the North Mountain Basalt on Partridge Island near Parrsboro, Nova Scotia. They stated that these beds, which are at least partly lacustrine or paludal in origin, have a Corollinadominated palynoflora. In the Newark Supergroup, this flora first appears in uppermost Rhaetian beds immediately above the major unconformity that includes uppermost Norian and most of Rhaetian time (Fig. 7). Although this palynoflora formerly was considered to mark the base of the Hettangian (e.g. Olsen et al. 2003), its Rhaetian age in the Fundy Basin recently has been established by Cirilli et al. (2009). This suggests that the Partridge Island Member is correlative with the beds just below the Mt. Zion Church Basalt in the Culpeper Basin (described above) that have produced a monospecific Euestheria brodieana conchostracan fauna. At the eastern end of its type section, the Partridge Island Member rests on weathered fluvial sandstones with apparent unconformity. The age and stratigraphic position of this unit, at its type locality, is well defined and firmly established. In other parts of the Fundy Basin, however, such as at the bluffs near Old Wife Rock and at Cape Blomidon, very light-coloured strata that have been referred to the Partridge Island Member almost certainly do not belong to that unit. Instead, these stratigraphic intervals appear to represent thick paleosols at the top of the Blomidon Formation immediately below the North Mountain Basalt. Unlike the Partridge Island Member, these paleosols are unfossiliferous, blocky in texture, and entirely lack the rhythmically interbedded gray and red pattern so obviously developed at Partridge Island. These paleosols are some

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of the strongest physical evidence yet found for the presence of a long hiatus between deposition of the bulk of the Blomidon Formation and the time of extrusion of the North Mountain Basalt. The Partridge Island Member apparently occupies broad swales cut into the top of the underlying part of the Blomidon Formation, and therefore occurs only discontinuously beneath the North Mountain Basalt (Fig. 10). The age range of Newark Supergroup strata in the Fundy Basin seems to be approximately as long as the age range of Newark Supergroup strata in the other Newark Supergroup basins, but the thicknesses of the Fundy strata are only about half the thicknesses of age-equivalent strata documented elsewhere. This suggests that the rate of deposition in the Fundy Basin during the Upper Triassic and basal Jurassic generally was much slower than it was in the other basins of the Newark Supergroup, though curiously the North Mountain Basalt is much thicker than its correlative basalts farther south.

Southwestern United States The tectonic setting of the Triassic strata in the southwestern United States is very different from the setting of the Newark Supergroup, which was deposited in a series of rift-related valleys across the heart of central Pangaea. In the southwestern United States, deposition during the Triassic (and also during the Early Jurassic) was within a back-arc basin in a broad coastal floodplain setting along the western margin of Pangaea. The main occurrences of continental deposits are in two regions, the southern High Plains of the western United States in northern Texas and eastern New Mexico, and on the Colorado Plateau in northwestern New Mexico, northeastern Arizona, and southern Utah (Fig. 11). The stratigraphic units in different parts of these regions in some areas have been given identical names and sometimes different names, depending on local mapping history and variations in facies that are found in different areas (Fig. 12). Unlike in the eastern United States, part of the Lower and Middle Triassic are represented by the Moenkopi Formation (Ward 1901), which consists of terrestrial shales, sandstones, and occasional impure limestones. In northern Arizona, the Moenkopi has been divided into three members: the Wupatki (McKee 1951, 1954), the Moqui (McKee 1951, 1954), and the Holbrook (Hager 1922; McKee 1951, 1954) from oldest to youngest. The oldest Triassic conchostracans in the southwestern United States have been found within the upper part of the Wupatki Member near Meteor Crater in Arizona (see Table 2 for details). This fauna

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Fig. 11. Map showing areas in Texas, New Mexico, Arizona, and Utah where stratigraphically important collections of Upper Triassic and Lower Jurassic conchostracans have been made. 1, St. George region, Utah (Whitmore Point Member of Moenave Formation); 2, Moab region, Utah (Kayenta Formation); 3, Potter Canyon, Arizona (Moenave Formation, Whitmore Point Member); 4, Petrified Forest region, Arizona (Moenkopi Formation, Blue Mesa Member of Petrified Forest Formation, Bluewater Creek Formation); 5, Fort Wingate region, New Mexico (‘Lake Ciniza’ in Bluewater Creek Formation); 6, Taos region, New Mexico (Whitaker Coelophysis quarry in Rock Point Formation); 7, Apache Canyon region, New Mexico (Redonda Formation); 8 – Lubbock region, Texas (Tecovas and Cooper Canyon formations). Localities 1 through 6 are on the Colorado Plateau; localities 7 and 8 are beneath the High Plains region of the west–central United States.

has yielded the following conchostracans: Palaeolimnadia alsatica detfurthensis Kozur & Seidel, very rare P. nodosa (Novozhilov) and Euestheria exsecta (Novozhilov) (Kozur, Lucas & Morales, in prep.). These forms indicate a Late Olenekian (Spathian) age for the Wupatki. This locality also is the type locality of the temnospondyl amphibian Wellesaurus peabodyi (Welles & Cosgriff 1965). The upper Moenkopi Holbrook Member at the Cottonwood Ruins locality near Winslow, Arizona has yielded large Euestheria albertii albertii (Voltz), which is the guide form for the lower Anisian Bithynian substage. Specimens as large as 6 mm in length, which are common among the Cottonwood Ruins material, are found only in the upper Bithynian. Therefore, the Holbrook can be placed confidently in the lower Anisian (Lucas & Schoch 2002). Above the Moenkopi there is a long hiatus in deposition extending well into the Carnian (Fig. 12). This hiatus ended with deposition of the widespread Shinarump Conglomerate (Gilbert 1875) (later designated as a formation of the Chinle Group) and its probable stratigraphic equivalent the Camp Springs Conglomerate (Beede & Christner 1926) [later designated as a formation of the Chinle (or Dockum) Group]. This interval has a distinctive vertebrate fauna that is representative

of the Otischalkian faunachron, but so far no conchostracans are known from this interval. Immediately above the Shinarump, the next younger unit is the Bluewater Creek Formation (Lucas & Hayden 1989) (originally named as a member of the Chinle Formation, now part of the Chinle Group). The Bluewater Creek has yielded conchostracans in a lacustrine sequence known as ‘Lake Ciniza’. Kozur has collected specimens from this interval that are dominantly Anyuanestheria wingatella (Tasch), but also include rare specimens of Laxitextella seegisi Kozur, which is the guide form found in the Lehrberg Beds in the Germanic Basin (upper part of lower Tuvalian), and also specimens of Howellisaura princetonensis (Bock), which also allows correlation of this interval with the lower part of the Lockatong Formation in the Newark Basin. Additionally, the vertebrate fauna from this unit includes the lowest occurrence of vertebrates characteristic of the Adamanian faunachron. Thus, this locality has produced a key fauna for correlation. The Lake Ciniza conchostracan fauna can be closely correlated with conchostracan faunas in Europe, with a part of the Tecovas Formation of Texas, and with the lower Lockatong Formation of the Newark Basin, but its correlation relative to

TRIASSIC CONCHOSTRACANS

353

Fig. 12. Triassic and part of Lower Jurassic stratigraphy in part of the southwestern United States, showing the correlation of units across the area and their stratigraphic positions relative to the international marine timescale. Horizons that have produced conchostracan are shown by asterisks. Although Lucas (2010; Lucas & Tanner 2007) have reported Protosuchus from the Dinosaur Canyon Member of the Moenave Formation, the locality they reported is south of the Colorado River where the upper Dinosaur Canyon lithology lies at a horizon laterally equivalent to the Whitmore Point Member in northwestern Arizona and southeastern Utah. Therefore, the base of the Wassonian LVF does not extend down section as far as the Dinosaur Canyon Member in the area indicated on our chart. The boundary between the Adamanian and the Revueltian land vertebrate faunachrons could lie at the top, within, or at the base of the Sonsela Member of the Petrified Forest Formation.

another allegedly Bluewater Creek conchostracan fauna is problematic. According to Spencer Lucas (Albuquerque, personal communication) the Lake Ciniza strata belong within the middle part of the Bluewater Creek Formation. However, conchostracans from another (allegedly lower) Bluewater Creek locality (Placerias Quarry in Arizona) constitute a monospecific fauna of Palaeolimnadia n. sp., which is very similar and perhaps even identical with Palaeolimnadia n. sp. from the upper Schilfsandstein (uppermost Julian) of central Europe. This fauna is considerably older than the Lake Ciniza fauna. The overlying lower Blue Mesa Member of the Petrified Forest Formation is dominated by Gregoriusella n. sp., which is very similar to (and probably identical with) Gregoriusella n. sp., which occurs in the upper Tuvalian Coburg Sandstein. A single fragmentary

conchostracan from this locality has the microsculpture of Laxitextella dorsorecta and, based on the size and shape of this fragment, was from a specimen the same size as this species. Thus, a late Tuvalian age is probable for the lower Blue Mesa Member. This is in reasonable accord with the late early Tuvalian to middle Tuvalian age for the underlying Lake Ciniza fauna, but the ‘lower Bluewater Creek’ fauna seems much too old relative to these other two localities to be at all likely to belong within the Bluewater Creek Formation. This oldest fauna better matches the age that has been established for the Shinarump and Santa Rosa formations. In the latter formation, Bachmann, Kozur and Franz found a macroflora that is similar to that of the Lunz Beds in the Northern Alps and the Schilfsandstein in the Germanic Basin (late Julian). A similar flora from the Santa Rosa was

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Table 2. Conchostracan localities in the Southwestern United States Locality Arizona Meteor Crater (MNA locality 1435) Cottonwood Ruin (MNA locality 1436) Placerias Quarry (NMMNH locality 858)

Age

Longitude

Moenkopi Formation, Wupatki Member Moenkopi Formation, Holbrook Member Bluewater Creek Formation, lower member

On file at MNA On file at MNA 34.43598N

On file at MNA On file at MNA 109.45788W

Petrified Forest Formation, Blue Mesa Mbr. Moenave Formation, Whitmore Point Member Moenave Formation, Whitmore Point Member

34.94208N

109.76678W

36.88028N

112.84668W

36.88028N

112.84668W

Late early (middle) Bluewater Creek Formation, Tuvalian middle part Sevatian Rock Point Formation

35.46148N

108.55638W

36.33558N

106.46388W

Late Sevatian to early Rhaetian

34.99768N

103.40708W

Late Bithynian (early Anisian) ?Early Tuvalian

P-50786 (NMMNH locality Upper Tuvalian 3631) Potter Canyon (bed 22, Latest Rhaetian locality 7734) Potter Canyon (bed 43, Earliest Hettangian locality 7735) New Mexico Lake Ciniza beds (NMMNH locality 1864) Whitaker quarry (NMMNH locality 3115) Apache Canyon (NMMNH locality 3590)

Texas (Texas Tech locality 3628)

Redonda Formation

Late early (middle) Tecovas Formation Tuvalian

(Texas Tech locality 0690)

Tuvalian

Tecovas Formation, lower part

Utah (NMMNH locality 6817)

Earliest Hettangian

Moenave Formation, Whitmore Point Mbr. Kayenta Formation

Moab (NMMNH locality 3293)

Sinemurian

Other data

Conchostracan zone

Navajo Co., Az.

Euestheria exsecta Zone

Navajo Co., Az.

Upper Euestheria albertii albertii Zone Zone indeterminate, but no younger than lower Tuvalian Zone indeterminate but upper Tuvalian Euestheria brodieana Zone Bulbilimnadia killianorum Zone

Salado Quadrangle, Apache Co., Az. Adamana Quadrangle, Apache Co., Az. Mocassin Quadrangle, Mohave Co., Az. Mocassin Quadrangle, Mohave Co., Az. FortWingate Quadrangle, McKinley Co., N.M. Ghost Ranch Quadrangle, Rio Arriba Co., N.M. Apache Canyon Quadrangle, Quay Co., N.M.

Anyuanestheria wingatella Zone Shipingia olseni Zone Redondestheria grovetonensis Zone, Shipingia olseni Zone, Gregoriusella polonica Zone

On file at On file at Collette Springs Texas Tech Texas Tech Quadrangle, Crosby Co., Texas On file at On file at Kalgary Quadrangle, Texas Tech Texas Tech Crosby Co., Texas

Anyuanestheria wingatella Zone

37.09968N

113.53498W

38.64968N

109.73498W

Bulbilimnadia killianorum Zone Zone not yet defined

St. George Quadrangle, Washington Co., Utah Jug Rock Quadrangle, Grand Co., Utah

Zone indeterminate

H. W. KOZUR & R. E. WEEMS

Latitude

Late Olenekian

Unit

TRIASSIC CONCHOSTRACANS

described by Ash (1988). Therefore, the Santa Rosa and Shinarump formations may be time equivalents and facies equivalents to the Schilfsandstein as has been suggested by Prof. G. H. Bachmann (Halle, pers. comm.). The lower Bluewater Creek Formation with Palaeolimnadia n. sp. also may be a lateral facies equivalent of the upper Shinarump and upper Santa Rosa formations, or else it could represent an unusual facies of the Shinarump. The Blue Mesa Member of the Petrified Forest Formation (Heckert & Lucas 2002; originally a member of the Chinle Formation of Gregory, 1950) has yielded conchostracans known elsewhere from the late Tuvalian. This stratigraphic assignment is in accord with the late Tuvalian age assigned to this unit by Litwin et al. (1991). There are several unconformities at the base and within the overlying Sonsela Member of the Petrified Forest Formation (Lucas et al. 2007a). Neither the Sonsela Member of the Petrified Forest Formation (Heckert & Lucas 2002) nor the laterally equivalent Trujillo Formation (Gould 1907) have yielded any well-preserved conchostracans. The vertebrates that are present in this interval are suggestive of the lower part of the Revueltian faunachron. The base of this LVF, however, is defined by the first occurrence of Aetosaurus and it has not been found in this unit yet, leaving open the possibility that the Sonsela is uppermost Adamanian. The Sonsela has yielded late Tuvalian palynomorphs (Litwin et al. 1991), which could be either in place or reworked from the beds immediately below. Therefore, whether the Sonsela is Carnian or Norian and whether it is Adamanian or Revueltian remains uncertain. The base of the Norian cannot be defined as yet by conchostracans in the southwestern United States. It may lie at the Adamanian-Revueltian tetrapod boundary, which lies at the top, within, or at the base of the Sonsela Member of the Petrified Forest Formation. Unfortunately, the Sonsela has yielded no well-preserved concostracans so far. In the underlying Blue Mesa Member, only a single wellpreserved but low diversity conchostracan fauna is known which contains Gregoriusella n. sp. This form also occurs in the late Tuvalian L. freybergi Zone of the Germanic Basin. Its presence tends to support a late Tuvalian age for the Blue Mesa Member, but it is not a robust support because neither the late Tuvalian guide form L. freybergi nor any early Norian guide form (e.g. Euestheria buravasi or Palaeolimnadia schwanbergensis) is present. Therefore, the Norian base as defined by conchostracans is undeterminable for now in the southwestern United States. At present there are two mutually exclusive ages that have been proposed for the Carnian– Norian boundary, each based on radiometric dates. These

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conflicting dates have produced a ‘short Norian model’ and a ‘long Norian model’. New radiometric ages from the Chinle Group suggest that the ‘short Norian model’ (e.g. Ogg et al. 2008) may be correct. This model assumes a Norian duration of only about 13 Ma with its base at around 216.5 Ma. Irmis & Mundil (2008) dated the base of the Blue Mesa Member of the Petrified Forest Formation as being 219.2 + 0.7 Ma. J. Ramezani presented very similar values (219.4 Ma) for the base of the Blue Mesa Member in a lecture (New Chinle Ages) at the Colorado Plateau Coring Project meeting in Albuquerque (May 2009). Heckert and Lucas also have an unpublished analysis by Jim Mortensen that yielded an c. 219 Ma age for the base of the Blue Mesa (Spencer G. Lucas, Albuquerque, pers. comm., 2009). In his talk on New Chinle Ages, Ramezani also presented radiometric ages of 217.84 + 0.05 Ma for the Sonsela Member and 209.78 + 0.12 Ma for the Black Forest Beds of the Painted Desert Member. The latter unit earlier was dated by Riggs et al. (2003) as being between 209 and 213 Ma, though they assumed that the actual age was closer to 209 Ma. Together with the radiometric ages of Ramezani for the Sonsela Member, the base of the Norian would seem to lie either around 218 Ma (base of Sonsela) or 217 Ma (top of Sonsela), either of which ages supports the ‘short Norian model’ of Ogg et al. (2008). Further studies are necessary to solve this problem, however. The only conchostracans known from the Petrified Forest Formation are from the basal Blue Mesa Member, and they represent a low diversity fauna of equivocal age. If the vertebrate biostratigraphic data for the Petrified Forest Formation can be confirmed by robust conchostracan data, then the ‘short Norian model’ would be probable. Even if the age of the Blue Mesa Member proves to be younger than previously assumed, however, the ‘long Norian model’ remains possible if the upper part of the Adamanian LVF ranges up through the lower Norian (Lacian). However, such a conclusion is not supported by conchostracan stratigraphy anywhere outside of the southwestern United States. We used the ‘long Norian model’ in the present paper, specifically the submodel of Bachmann & Kozur (2004) and Kozur & Weems (2007) that proposes a duration for the Norian of c. 17–20 Ma based on an age for the Carnian –Norian boundary of 225 + 3 Ma. Most other authors who advocate the ‘long Norian model’ prefer a much longer duration for the Norian (e.g. Gallet et al. 2003 with a Norian duration of 25 Ma). The main suport for the ‘long Norian model’ is an age of 225 + 3 Ma that was obtained from lower Norian rhyolites in southeastern Alaska (Gehrels et al. 1987). According to the conodont data of Savage & Gehrels

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(1987), the rhyolites from which the radiometric age was obtained are bracketed below by beds with Metapolygnathus primitius, 1987, which was used in 1987 as a group of species of latest Carnian to earliest Norian age, and above by Epigondolella abneptis, 1987, which was used in 1987 as a group of lower Norian species that includes E. quadrata Orchard and E. triangularis (Budurov), though the holotype of E. abneptis comes from the middle Norian. Thus, Gehrels et al. (1987) clearly have established that the reported radiometric age was taken from within lowermost Norian strata. A modern re-investigation of this radiometric age is necessary, but it is noteworthy that a 230.91 + 0.33 Ma for the late early Tuvalian (Furin et al. 2006) and a corrected age of 231.4 Ma for the Adamanian LVF of Ishigualasto (Argentina) (Irmis & Mundil 2008) also support the ‘long Norian model’ because, if the base of the Norian lies at about 217 or 218 Ma, then the Tuvalian would be longer than the entire Norian, which seems very improbable. Above the Sonsela Member, neither the Painted Desert Member of the Petrified Forest Formation (Heckert & Lucas 2002), the laterally equivalent Bull Canyon Formation of New Mexico (Lucas & Hunt 1989) nor the Cooper Canyon Formation of Texas (Lehman et al. 1992), have yielded conchostracans. However, Litwin et al. (1991) have reported Norian palynomorphs from this interval. This interval also has produced abundant vertebrates of the Revueltian faunachron. No Triassic strata are preserved above the Cooper Canyon Formation in Texas, and there is a major unconformity above the laterally equivalent Bull Canyon Formation in eastern New Mexico as well. In northwestern New Mexico and northeastern Arizona a somewhat younger unit occurs, the Owl Rock Formation (Lucas 1993; Lucas et al. 1997; originally described as the Owl Rock Member of the Chinle Formation by Stewart 1957) that also has not yielded conchostracans. This unit has yielded vertebrates, however, which are referable to the Revueltian faunachron (Lucas et al. 2007a; Spielmann et al. 2007) and indicate that it is not greatly younger than the underlying Petrified Forest Member. Above this level, there is a minor unconformity within the middle Norian (Alaunian). Deposition resumed regionally in the late Norian (Sevatian) and laid down a series of deposits previously mapped as the Church Rock Member (Stewart 1957) in southern Utah, the Church Rock Formation of the Chinle Group (Lucas 1993; Lucas et al. 1997) in north-northeastern Arizona, the Rock Point Formation (Lucas 1993; Lucas et al. 1997; originally described as the Rock Point Member of the Wingate Sandstone by Harshbarger et al. 1957, later moved to the Chinle Formation by Dubiel

1989) in east-northeastern Arizona and northwestern New Mexico, and the Redonda Formation in eastern New Mexico (Griggs & Read 1959; originally described as the Redonda Member of the Chinle Formation by Dobrovolny & Summerson 1946). The Church Rock and the Rock Point are different names for what is essentially the same stratigraphic interval. These names were applied in different geographical areas and both were used for the first time in the geological literature in the year 1957. Lucas (1993) has suggested that the name Church Rock (with no designated type locality) be replaced by the name Rock Point (which does have a type locality), and this suggestion is followed here. This sequence has yielded conchostracans from two localities. One occurrence, at the Apachean-age Whitaker Coelophysis quarry in the Rock Point Formation of northwestern New Mexico, has yielded a new species of very large Shipingia, different from S. olseni and bearing a welldeveloped microsculpture that makes it more advanced than S. hebaozhaiensis. This form is too advanced to be Alaunian and therefore can be assigned a Sevatian age. The other occurrence is in the Redonda Formation of northeastern New Mexico, where several conchostracan species have been found including the type material of Redondestheria novomexicoensis and large specimens of Shipingia olseni. Above the S. olseni Zone, Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej occurs. This species characterizes the latest Sevatian(?) and basal Rhaetian in northern Germany and Lipie (Poland), so the highest strata in this sequence appear to range upward into the lowest part of the Rhaetian. Because the vertebrates within this sequence were the basis for creating the Apachean faunachron, the discovery of Sevatian-age conchostracans within it means that its stratigraphic range must be revised downward to include the Sevatian. The immediately overlying Dinosaur Canyon Member of the Moenave Formation (Harshbarger et al. 1957) has yielded no conchostracan as yet. However, conchostracan have been reported from the Whitmore Point Member of the Moenave (Wilson 1967; Lucas & Milner 2006). Specimens found in the upper Whitmore Point Member that we received from Dr Andrew R. C. Milner (St. George, Utah) are mainly Euestheria brodieana (Jones) but also include a few B. killianorum. Additionally, Dr Spencer G. Lucas (Albuquerque) has provided material from his bed 43 (a 0.5 m-thick purple-brownish mudstone unit) of the upper Whitmore Point Member in Potter Canyon, northern Arizona (NMMNH locality 7735), about 1 km west of the type section of the Whitmore Point Member. This well-preserved material, from about 3.5 m below the Springdale Member of the

TRIASSIC CONCHOSTRACANS

Kayenta Formation, is dominated by E. brodieana but also has common B. killianorum. These two occurrences demonstrate that the upper part of the Whitmore Point Member is earliest Hettangian in age and correlative with the basal Hettangian strata of Poland and the middle part of the Midland Formation above the basal Mt. Zion Basalt in the Culpeper Basin in the eastern United States. This conclusion is in accord with the palynological results of Cornet & Waanders (2006). Dr Spencer G. Lucas and Dr Wolfram Ku¨rschner (Utrecht) have discovered a third and lower horizon in the lower Whitmore Point Member at Potter Canyon. This material comes from bed 22 of Lucas (a 2.6 m-thick black and greenish shale unit) that lies about 6 m above the Dinosaur Canyon Member of the Moenave Formation. More than 400 well-preserved conchostracan specimens were collected by Spencer G. Lucas and Wolfram Kuerschner from these shales and all belong to Euestheria brodieana. A monospecific E. brodieana fauna is characteristic of the upper Rhaetian E. brodieana Zone, and this age designation is in agreement with data presented by Wolfram Ku¨rschner at the Colorado Plateau Coring Project meeting in Albuquerque (May 2009), based on the presence of a carbon isotope minimum at this stratigraphic level that is the initial carbon isotope minimum within the uppermost Rhaetian. Thus, the Triassic – Jurassic boundary lies within the Whitmore Point Member, somewhat higher in the section than so far assumed. The palynoflora from the upper part of the underlying Dinosaur Canyon Member is close in age (Litwin et al. 1991; Cornet & Waanders 2006) and belongs to the top of the Rhaetian according to the new conchostracan and isotopic data from the lower Whitmore Point Member. Of particular importance is the fact that B. killianourm has never been found anywhere in Rhaetian strata, even in the uppermost Rhaetian strata of the Germanic Basin and England. Latest Rhaetian conodonts (including Misikella posthernsteini Kozur & Mock) are present in England at a level where E. brodieana occurs in the absence of B. killianorum, and this level clearly is below the level of the marine Psiloceras spelae Zone. Therefore, in continental beds, the B. killianorum Zone starts at a level no lower than that at which Psiloceras spelae Guex et al. occurs. On the other hand, the B. killianorum Zone also begins below the marine P. planorbis Zone. Therefore the B. killianorum Zone begins in (and may correspond to) the P. spelae Zone of the marine realm. The Kayenta Formation (Baker 1936), of probable late Hettangian and/or Sinemurian age (Padian 1989), lies disconformably above the Moenave Formation. This is the youngest unit on the Colorado Plateau that has yielded

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conchostracans. Their study is beyond the scope of the present paper, but it is worth noting that they seem to represent forms younger than the Bulbilimnadia froelichi Zone. They also may be younger than the conchostracans found in the middle of the Portland Formation, which in turn suggests that the youngest possible age for the Portland Formation of the Newark Supergroup is late Hettangian or early Sinemurian.

Regional occurrence of conchostracans during the Early, Middle and Late Triassic Early Triassic Early Triassic to early Anisian conchostracans are very common in the Germanic Basin, where they have been long known and well studied, sometimes in continuous successions (see references in the last section). The synonyms of all of these species also are well known (Kozur 1982, 1983; Kozur & Seidel 1983a; Kozur & Hauschke 2008). Additionally, Lower Triassic conchostracans are well known (but mainly without modern revision) from the Russian Platform and adjacent areas such as the Precaspian Basin and Mangyshlak (see references in the previous section). During the Early Triassic, conchostracans were common, not only in the Germanic Basin, on the Russian Platform, and in adjacent areas such as the Precaspian Basin and Mangyshlak, but also in the Tethys in the uppermost Permian to Upper Olenekian (Spathian) Werfen Group (Kozur 1987, 1999; Kozur & Mock 1993; Bachmann & Kozur 2004). Lower Triassic or uppermost Permian to Lower Triassic conchostracans also are well known from Siberia (Chernyshev 1934; Novozhilov 1946; Novojilov [Novozhilov], 1958b; Molin 1965b; Sadovnikov & Orlova 1990, 1993; Orlova 1999), Kazakhstan (Zaspelova 1961), and China (Chang et al. 1976; Liu 1982, 1987b, 1988, 1989, 1990; Chen & Shen 1985; Liu et al. 1990). A few occurrences of Lower Triassic conchostracans are known from continental beds in North America (upper Wupatki Member of the Moenkopi Formation; Kozur, Lucas & Morales, in prep.) and brackish water beds in Greenland (Defretin-Lefranc 1969; Kozur & Seidel 1983a). Lower Triassic conchostracans are common as well in Gondwana, including India (Jones 1862; Ghosh & Shah 1978; Ghosh 1983; Ghosh et al. 1988; Tasch 1987), Australia, (Mitchell 1927; Cockbain 1974; Tasch & Jones 1979; Tasch 1987) and Africa (Janensch 1927; Leriche 1932; Marlie`re 1950; Teixeira 1950a, b, 1951; Kozur & Seidel 1983a, b). In South America, Lower Triassic conchostracans have not

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been mentioned until now (see summary by Tasch 1987).

Middle Triassic Middle Triassic conchostracans are not so widely distributed as the Lower Triassic ones. Conchostracans were abundant in a few areas during the early Anisian (Aegean and lower Bithynian of the Ro¨t and Gre`s a` Voltzia) and late Ladinian (brackish equivalents of the upper Hauptmuschelkalk above the Cycloides Bank, and especially the Erfurt Formation ¼ Lower Keuper) (Fig. 4). They have been recognized for quite a long time and are rather well studied (von Zieten 1830; Voltz 1835; Jones 1862, 1863, 1890, 1891; Defretin 1950; Reible 1962; Defretin-Lefranc 1963; Warth 1969; Gall 1971, 1976; Kozur 1974, 1982, 1999b; Kozur et al. 1993) and well correlated with the marine scale. In the middle Anisian through lower Ladinian (Pelsonian through Fassanian) interval, however, known strata are either marine or hypersaline and do not contain conchostracans. Outside the Germanic Basin, Middle Triassic conchostracans have been studied from occurrences that are not well dated or else are isolated and lack context to establish the local conchostracan succession. Aegean and lower Bithynian conchostracans are known from the Mecsek Mountains in southern Hungary (Vada´sz 1952; Nagy 1959, 1960, 1968; Kozur et al. 1993), where they occur in brackish beds intercalated with brachyhaline marine deposits containing Costatoria costata (Zenker) and, in the Bithynian, Myophoria vulgaris (von Schlotheim). Well-dated conchostracan-bearing upper Bithynian (uppermost lower Anisian) strata with low faunal diversity are known from the Holbrook Member of the uppermost Moenkopi Formation of Arizona, USA (Kozur, Lucas & Morales, in prep.). A monospecific Euestheria minuta (von Zieten) fauna has been found in the lower Wolfville Formation in the Fundy Basin and will be described in a future paper. It indicates a Longobardian age for the lower Wolfville. Low-diversity, but characteristic upper Anisian (Illyrian) conchostracan faunas were described from northern Siberia close to the shore of the Laptev Sea; these have been well dated by bivalves and ammonoids (Novozhilov 1946, 1958b, c). Middle Triassic conchostracans of similar age (Anisian to lower Ladinian) are known from China, where they are taxonomically well studied (especially in Chang et al. 1976) but no more specifically dated than Middle Triassic. Little is known about the Middle Triassic conchostracans of Gondwana. Mitchell (1927), Raymond (1946), Tasch & Jones (1979), and Tasch (1987) have described a few Middle Triassic conchostracans from Australia, mainly under the

name of the long-ranging form Palaeolimnadia wianamattensis (Mitchell). These occurrences actually may comprise more than one species. Gallego (1992) described an upper Ladinian fauna with Euestheria minuta (von Zieten) from Mendoza and San Juan in Argentina and a Middle Triassic fauna from the Santa Maria Formation off Rio Grande do Sul, Brazil. Katoo (1971) described different conchostracans from the lower part of this formation, among them forms he classified as Lioestheria sp. and some Magniestheria sp. of a size and outline that are known elsewhere only in the Olenekian to lower Anisian. On the basis of these identifications, he and Tasch (1987) assigned the Santa Maria Formation to the Late Triassic. Anderson & Anderson (1993) concluded that the tetrapods of the upper Santa Maria Formation indicate a Norian age for the upper part of that formation. However, Lucas (2002) has concluded that the tetrapods of the upper Santa Maria Formation are not Norian but Middle Triassic to late Carnian in age. As no tetrapods are known from the lower Santa Maria Formation, where conchostracans occur, either age is compatible with the age suggested by conchostracans for the lower Santa Maria Formation. Besides Middle Triassic forms, Katoo (1971) has illustrated from the upper Santa Maria Formation, under the name Euestheria cf. E. minuta (von Zieten), a Laxitextella of the L. seegisi Kozur group, which does not range higher than the Upper Carnian (middle Tuvalian). This latter occurrence supports the age range suggested by Lucas for the upper Santa Maria Formation.

Upper Triassic Upper Triassic conchostracans are as widely distributed as Lower Triassic ones. However, generic assignment of many of these species is not yet clear. Among described species, many are synonyms with some species being assigned by different authors to several different species and genera. In other cases, new species have been established by authors who were unaware that these same species already had been described from other regions. In addition to these considerable problems, many species were established on poorly preserved material that was neither well described nor well documented and needs to be revised by re-sampling from the type localities. Finally, a large percentage of Upper Triassic conchostracans have yet to be described. The most complete lower and middle Carnian conchostracan successions are known from the Germanic Basin, and these successions are well correlated with the marine timescale. The taxonomy of Cordevolian conchostracans is well known (Reible 1962; Warth 1969; Kozur 1982; Kozur & Weems

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2007). Julian conchostracans of the Schilfsandstein have been sampled and photographed but are not yet described. The lower Julian has not produced conchostracans because there is a rather long stratigraphic gap in this part of the section in the Germanic Basin. In other parts of the world, well dated lower Julian conchostracans sensu stricto (not including Cordevolian forms incorrectly assigned to the lower Julian) are unknown. Four horizons with conchostracans are known from the Tuvalian Weser Formation and time equivalents. They occur in eastern France and southern Germany at the base of the Tuvalian (in equivalents of the Dolomie de Beaumont), in the higher part of the lower to middle Tuvalian Lehrberg Beds, in the upper Tuvalian Coburg Sandstein, and in the uppermost Tuvalian Heldburggipsmergel Member of the Weser Formation. The Tuvalian conchostracans of the Germanic Basin are taxonomically well studied and well dated (Reible 1962; Olempska 2004; Kozur & Weems 2007). Even so, if there is to be a comprehensive Tuvalian conchostracan zonation for the northern hemisphere, a synthesis still needs to be made with the Tuvalian conchostracan succession of the Newark Supergroup rift basins in eastern North America and with Tuvalian conchostracan faunas in the southwestern United States (Arizona, New Mexico, Texas). In the Germanic Basin, no conchostracans are present between the four stratigraphic levels mentioned above, all of which represent independent zones with only the two upper zones in direct succession. The intervening stratigraphic levels, barren of conchostracans, represent times of hypersaline deposition in which conchostracans could not live. In the Newark rift basins, the Tuvalian has in several rift basins continuous conchostracan successions in lake deposits without any prolonged hypersaline intervals. Additionally, the equatorial conchostracan fauna of the southern Newark rift basins is more diverse than the conchostracan fauna from the northern Newark rift basins and the still more northerly Germanic Basin. A conchostracan fauna nearly identical with that of the equatorial southern Newark rift basins occurs in the Tuvalian of the southwestern United States, but so far only Anyuanestheria wingatella (Tasch) has been established in the literature (Tasch 1978). Carnian conchostracans also have been reported from other parts of Europe in fresh to brackish water intervals between marine sediments, for example, in the lower Carnian of the westernmost Southern Alps and western Carpathians (Tintori 1990; Tintori & Brambilla 1991; Kozur & Mock 1993). During the middle Carnian wet intermezzo of late Julian age (Kozur & Bachmann 2008), a large influx of fresh water took place in the northwestern Tethys region, resulting in brackish horizons within the

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Lunz Formation that contain conchostracans (e.g. Hornung 2007); these have yet to be studied in detail. Kobayashi (1952, 1954, 1975) has described or mentioned Upper Triassic conchostracans that occur in the Yamaguchi Prefecture, Japan within the Atsu Group just below the Halobia shale. ‘Estherites’ nakazawai Kobayashi is a Palaeolimnadia from which only one specimen was illustrated. The holotype (and most of the illustrated specimens) of ‘Estherites’ atsuensis Kobayashi also belong to Palaeolimnadia, and it may be that P. nakazawai is only a slightly deformed P. atsuensis. There is a second taxon, however, that he included with this species that is an Anyuanestheria with radial lirae. This fauna was assigned to the early Carnian by Kobayashi (1952, 1954, 1975) because the immediately overlying beds contained the bivalve Halobia multistriata Kobayashi & Aoti (¼ Halolobia aotii Kobayashi & Ichikawa), which then was regarded as an early Carnian species (Kobayashi 1963). These conchostracans, however, are closely related to Palaeolimnadia schwanbergensis Reible, which was originally described in the Germanic Basin from the latest Carnian ‘Berggips Beds’ (Heldburggipsmergel Member of the Weser Formation). This species subsequently was found at another locality within the Heldburggipsmergel Member, where it also occurs together with upper Tuvalian conchostracans such as Laxitextella freybergi Kelber & Kozur. It occurs also as a monospecific fauna in the lowermost Arnstadt Formation (lowermost Norian) of the Germanic Basin, and in a similarly monospecific lowermost Norian fauna in the Warford Member of the Passaic Formation of the Newark Basin. Thus, the P. atsuensis/P. nakazawai fauna also is probably latest Carnian or possibly earliest Norian in age, which is in strong conflict with its present lower Carnian age assignment. According to Prof. Christopher A. McRoberts (Cortland, pers. comm., February, 2009), Halobia multistriata has been found in well dated beds from north-eastern Siberia that range in age from the lower Norian (e.g. upper Dawsoni Zone) up into the early middle Norian (Rutherfordi Zone), with an acme in the upper lower Norian (Magnus Zone). This means that a late Tuvalian to early Lacian age for the P. atsuensis conchostracan fauna from Japan now does agree with the age range of its associated molluscan fauna. Norian conchostracans are widely distributed and in part well dated, but many taxa still have to be revised and about 50% of the discovered taxa are yet to be described. Only a few of these new taxa will be described in the present paper. Norian conchostracans are accurately dated in several places in Asia (Kobayashi 1954). In eastern Asia,

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well dated and rather diverse lower Norian conchostracans are known from Nam Phrom in Thailand (Kobayashi 1973, 1975, 1984) and also from the Norian of China (Chen 1974; Chang et al. 1976; Chen & Shen 1980). Well dated, but poorly preserved Norian conchostracan faunas are known from Hikiji in Japan, from Tonjin and other localities in Korea, and from Mon Cay in Vietnam (Ozawa & Watanabe 1923; Kobayashi 1975, 1984). In western Asia, Novozhilov & Kapelka (1960) described a diverse conchostracan fauna from Madygen in western Kirgistan that they assigned to the lower Norian. Kobayashi (1975) has cast doubt on such an exact age determination for continental beds that lack marine intercalations or marine underlying or overlying beds, but the fauna contains six species that occur in better-dated Norian beds from other places in Asia. Thus a Norian age can be accepted, though it is not clear if the age is precisely lower Norian or perhaps slightly younger. Very rich conchostracan faunas occur in the Newark rift basins, where a continuous Norian conchostracan standard zonation can be developed. The richest lower Norian (Lacian) faunas are known from the Deep River, Culpeper and Newark basins, but these faunas are not yet fully revised and described. Only a few of these forms were illustrated in Kozur & Weems (2007). Middle Norian (Alaunian) faunas are known from the Culpeper, Gettysburg and Deep River basins. The richest upper Norian (Sevatian) faunas are known from the Culpeper, Newark and Fundy basins. Several uppermost Alaunian and Sevatian conchostracans of the Newark rift basins were described in Kozur & Weems (2007), and an additional form is described in the present paper. A very important upper Norian conchostracan fauna from within the Shipingia olseni Zone occurs in the upper Redonda Formation of New Mexico. A somewhat older upper Norian conchostracan fauna occurs in the Ghost Ranch Quarry of New Mexico in the Rock Point Formation. This fauna cannot be precisely located yet relative to the Sevatian conchostracan faunas from the Newark rift basins or from Germany, but the co-occurrence of this fauna in the Ghost Ranch Quarry with the parasuchian Redondasaurus (Lucas & Tanner 2007) clearly places it within the Apachean land–vertebrate faunachron (LVF). These conchostracans are more advanced than those of the uppermost Alaunian, which therefore means that they are Sevatian, but they also are older than the uppermost Sevatian conchostracans of the Apache Canyon locality from the uppermost Shipingia olseni Zone. Norian conchostracans occur in the Germanic Basin in the mid-basin Arnstadt Formation (Steinmergelkeuper) of northern and central Germany,

the contemporaneous Argiles bariole´es dolomitiques Formation of eastern France and in the more marginal basin Stubensandstein Formation (Lo¨wenstein Formation) of southern Germany. Except for a few local occurrences with rich faunas, conchostracans occur only rarely in the lower and middle Norian, and in the Sevatian they are only known from boreholes in northern Germany. Almost nothing has been published concerning the Norian conchostracans in the Germanic Basin. Warth (1969) illustrated a species from the lower part of the ‘3. Stubensandstein’ which he identified as Palaeestheria dorsorecta (Reible). The same material was reillustrated by Kozur & Weems (2007), determined to be an advanced Shipingia hebaozhaiensis Shen and assigned to the uppermost Alaunian. Hopf & Martens (1992) illlustrated (but did not identify) a single conchostracan from the lower Steinmergelkeuper (lower Norian) of western Thuringia. Other occurrences have been mentioned but neither described nor illustrated (e.g. Bourquin & Durand 2006 from the Argiles bariole´es dolomitiques Formation of eastern France). Norian conchostracans from Gondwana are often either inadequately described, based on poorly preserved material, or not well dated and called Upper Triassic without any more detailed placement. Only in Morocco have well-preserved middle Norian Shipingia hebaozhaiensis Shen been found by Paul Olsen (Columbia), which will be described in a future joint paper with him. This, however, is not part of a true Gondwanan fauna, but a fauna of northern hemisphere aspect. Well-dated Rhaetian conchostracan horizons in the northern hemisphere contain only low diversity faunas that include only small species of conchostracans. Lower Rhaetian conchostracans occur rarely in Western and Central Europe, but also are known from boreholes in northern Germany and from a quarry at Lipie in Upper Silesia (Poland). These lowest Rhaetian faunas consist nearly exclusively of Gregoriusella polonica n. gen. n. sp., both at Lipie and in northern Germany (in the basal few metres of the Exter Formation in a borehole at Tarnow). Only a few as yet undescribed additional generalized species of Euestheria also are present in the lower Rhaetian. Above this interval, the faunas start to include both Gregoriusella polonica n. gen. n. sp. and Euestheria brodieana. At first Euestheria brodieana is rare, but it progressively replaces Gregoriusella polonica up-section until, in the upper part of the lower Rhaetian, the faunas become monospecific Euestheria brodieana faunas that persist to the top of the Rhaetian. In the region of the Germanic Basin (England, France, Germany and Poland) upper Rhaetian faunas are much more widely distributed. They

TRIASSIC CONCHOSTRACANS

consist of monospecific faunas of Euestheria brodieana (Jones), described from England (Jones 1862, 1891; Reible 1962) and France (Battarel & Gue´rin-Franiatte 1971), with similar faunas also present in Germany (e.g. in the upper Rhaetian Triletes Beds in a temporary outcrop along the A4 near Moseberg close to Eisenach). The same monospecific E. brodieana fauna occurs in the upper Rhaetian beds immediately below the first lava flow in the Culpeper Basin. The so-called ‘Rhaetian –Liassic’ faunas with large conchostracans described by Bock (1953a), such as the Congestheriella olsoni (Bock) fauna of Venezuela, were incorrectly dated; they belong instead to the Late Jurassic and not to the Triassic (Gallego et al. 2009).

Conchostracan zonation from the latest Permian through the Hettangian Depending on the local geological setting (i.e. the presence or absence of continuous sequences with lake deposits), different areas at different stratigraphic intervals are most suitable for establishing a Triassic conchostracan zonation for the northern hemisphere. It is not unusual for conchostracan zonations to be combined from different regions. This is especially necessary to provide accurate correlation of the conchostracan zonation with the marine timescale. The uppermost Permian conchostracan zonation, including the conchostracan zonation across the Permian–Triassic boundary and the correlation of the conchostracan zonation with the marine scale around the Permian– Triassic boundary, can be best established by combining the conchostracan zonations at Dalongkou (Xinjiang) with the Tunguska Basin (in northern Siberia) and with the Germanic Basin of central Germany and southeastern Poland (Holy Cross Mountains). Conchostracan faunas around the Permian– Triassic boundary in northeastern Siberia and in the basal Werfen Beds of Hungary must be also considered for accurate correlation with the marine scale. The conchostracan zonation of the Lower Triassic and lower Anisian can be best established in continuous successions within the Germanic Basin in central and north-western Germany and in the lower Anisian in southern Germany and eastern France. The correlation of the Lower Triassic conchostracan zonation with the marine scale is best determined in the Werfen Group of Hungary, in the Southern Alps, and in the basal Triassic and Olenekian of northern Siberia. In the lower Anisian, intercalations of marine and conchostracanbearing beds occur in the Germanic Basin and also in the Mecsek of southern Hungary and in northern Siberia. For the Middle Triassic above the lower

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Anisian, the conchostracan successions of China are the best available for establishing a continuous conchostracan zonation, but for correlation with the marine timescale the conchostracans of northern Siberia and the upper Ladinian conchostracans of the Germanic Basin also must be considered. The zonation of Lower Carnian conchostracans was established in the Germanic Basin, and this is the area where the future zonation of late Julian conchostracans also will be developed once the conchostracans and conchostracan successions of the Schilfsandstein are described. No well dated early Julian conchostracans are yet known. The Tuvalian conchostracan zonation was established partly in the Germanic Basin and partly in the Newark Supergroup. Supplementary zonations need to be established for the diverse equatorial conchostracan faunas of the southern Newark Supergroup Basins and for the southwestern United States. The best areas for developing a detailed Norian conchostracan zonation are in the Newark Supergroup and in the Germanic Basin, but additional zones also must be described from the southwestern United States and perhaps also from China and southeast Asia. The Rhaetian conchostracan zonation has been established in the Germanic Basin. The early Hettangian conchostracan zonation has been developed in the Newark Supergroup, but important additional data are now available from the southwestern United States and from marginal parts of the Germanic Basin in southeastern Poland.

Upper Permian to basal Triassic conchostracan zonation The conchostracan zonation of the Upper Permian to lowermost Triassic (Fig. 1) is remarkably uniform from the very high palaeolatitudes of northern Siberia to the low palaeolatitudes of Dalongkou (Xinjiang) and the Germanic Basin. Close to the Permian –Triassic boundary, correlations with the marine scale are possible using the conchostracan faunas from northeastern Siberia and the basal Werfen Beds in Hungary. We present here a conchostracan zonation that starts with the beginning of the tuffaceous part of the Siberian Trap in the Tunguska Basin. The beginning of this zonation can be best established from one section in the Guodikeng Formation on the southern limb of the Dalongkou anticline, Xinjiang. From the Falsisca eotriassica Zone upward, this zonation previously was established in the Germanic Basin by Kozur & Seidel (1983b) and Kozur (1993b). The Falsisca eotriassica and F. postera zones established in the Germanic Basin are also present at Dalongkou and in northern Siberia (Tunguska Basin and Taimyr) in the same succession and with the same

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conchostracans. Below the F. eotriassica Zone, no conchostracans are present in the upper Zechstein (Late Permian) of Germany; only the lower Zechstein (Zechstein 1) has in its marginal facies yielded some conchostracans, among them a Megasitum n. sp., a species which is younger than the youngest Tatarian Megasitum but older than the upper Changhsingian Megasitum of the Guodikeng Formation (Dr Brigitte Hammerich, Freiburg, pers. comm.). Falsisca zavjalovi –Tripemphigus minutus Zone Definition. A distinctive fauna defined by the joint occurrence of the genera Bipemphigus, Falsisca, Megasitum and Tripemphigus. Lower boundary. First appearance of Falsisca zavjalovi and F. turaica. Upper boundary. Disappearance of the genera Bipemphigus, Tripemphigus and Megasitum (the latter genus in the upper part of the zone rare, may disappear somewhat earlier than the other two genera); in high latitudes additionally marked by the disappearance of Hemicycloleaia and Echinolimnadia. The upper boundary of this zone marks the strongest extinction event within Permian conchostracan faunas, well below the Permian– Triassic boundary. Important conchostracan species. Bipemphigus gennisi Novozhilov; Bipemphigus cf. B. gennisi Novozhilov (determined as Cornia beijiangensis Liu by Liu Shuwen in Cheng Zhengwu et al. 1989); Bipemphigus cf. B. liaoningensis Shen & Li; Tripemphigus minutus (Liu); T. sibiricus (Novozhilov) (occurrence at Dalongkou not quite certain, could be also a T. minutus with weaker microsulpture); T. khovorkiliensis Novozhilov (only in northern Siberia); Megasitum vanum (Novozhilov) (in Dalongkou determined as Megasitum cf. M. vanum); Falsisca zavjlovi (Novozhilov 1970) [determined by Liu Shuwen in Cheng et al. (1989) as Falsisca beijiangensis Liu 1987]; F. turaica (Novozhilov) [determined by Liu Shuwen in Cheng et al. (1989) as Falsisca cf. F. kanandaensis Novozhilov]. Additionally, there is Beijianglimnadia qitaensis Liu (¼ B. elegans Liu) (only at Dalongkou), B. dalongkouensis Liu (¼ B. minuta Liu, only at Dalongkou), Polygrapta subovata (Liu) (only at Dalongkou), Echinolimnadia mattoxi Novozhilov (only in northern Siberia), Hemicycloleaia mitchelli (Etheridge) with a distribution both in high northern palaeolatitude Siberia (Tunguska Basin) and high southern palaeolatitude eastern Australia (Newcastle CM Group).

Type locality. Dalongkou (Xinjiang), southern limb of the Dalongkou anticline, and 65–107 m above the base of the Guodikeng Formation. Occurrence. Dalongkou (Xinjiang), on the southern and northern limbs of the anticline; Guodikeng Formation of Xialongkou (Xinjiang), Changhsingian, on the southern limb of the anticline between 65 and 107 m above the base of the formation; northern Siberia (Tunguska Basin), in the Lebedeva and lower Khungtukun tuffs, Bugarikta Formation. Remarks. Both Tripemphigus sibiricus and T. minutus are very small forms and in almost all features very similar. T. minutus has a stronger microsculpture, but it is of the same type as in T. sibiricus. Possibly these are only preservational and documentation differences, because the Dalongkou material is well-preserved and the microsculpture of T. minutus is shown in SEM photographs, while the microsculpture of T. sibiricus is shown only in a light optic photograph. Thus, it is entirely possible that T. minutus (Liu 1987) is a junior synonym of T. sibiricus (Novozhilov 1965), but this can be clearly determined only after investigation of the Dalongkou material, which is still inaccessible [see section titled ‘Dalongkou (Xinjiang, NW China)’]. If T. minutus is a junior synonym of T. sibiricus, the name of the zone will need to be changed to the Falsisca zavjalovi –Tripemphigus sibiricus Zone. No change of the content of the zone would be necessary, however. We could avoid this problem by defining a Falsisca zavjalovi –Tripemphigus sibiricus Zone in the Tunguska Basin, but this is a remote area and the succession of conchostracans there is not in a successive section, so definition of this zone in the Dalongkou section is much more suitable. The lowermost known range of Hemicycloleaia mitchelli in the Tunguska Basin is the base of the Falsisca zavjalovi – Tripemphigus minutus Zone, but elsewhere this species surely ranges lower. The uppermost range of this species, however, does accurately reflect the top of the Falsisca zavjalovi – Tripemphigus minutus Zone. This species cannot be used for correlation of the base of the F. zavjalovi – T. minutus Zone to eastern Australia, for example, because there it appears earlier than the F. zavjalovi –T. minutus Zone, and this also may be the case in northern Siberia. However, the top of the F. zavjalovi –T. minutus Zone can be recognized in Australia by the disappearance of H. mitchelli in the uppermost part of the Newcastle CM Group of the Sydney Basin (for range of H. mitchelli in the Sydney Basin, see Jones & Chen 2000). The most characteristic genera in the F. zavjalovi –T. minutus Zone are typical Permian

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genera (Megasitum) or Changhsingian genera (Bipemphigus, Tripemphigus). The lower boundary is defined by the appearance of Falsisca zavjalovi, but Falsisca straddles the Permian– Triassic boundary and the most diverse Falsisca faunas occur in the Changhsingian part of its range. Liu in Cheng et al. (1989) tends to define the Permian– Triassic boundary by the first appearance of Falsisca. However, Falsisca jeskinoica (Novozhilov) occurs as low as the Tatarian on the Russian platform, and F. secreta likewise occurs in the Tatarian of eastern Kazakhstan. Thus, the base of the F. zavjalovi –T. minutus Zone in some areas is not defined by the level where the first member of the genus Falsisca appears, though in Dalongkou and northern Siberia this is the case. Beijianglimnadia also begins close to the base of the F. zavjalovi–T. minutus Zone and ranges above this zone, though it does not range into the Triassic. Polygrapta is common in the F. zavjalovi –T. minutus Zone and also in the underlying Changhsingian beds (lower Guodikeng Formation, Wutonggou Formation). Large representatives of Polygrapta disappear at the top of the F. zavjalovi –T. minutus Zone, though rare small Polygrapta persist into the lowermost Triassic. The northern and southern high latitude Hemicycloleaia mitchelli belongs to a typical Palaeozoic genus that died out at the top of the F. zavjalovi –T. minutus Zone. The northern latitude Echinolimnadia mattoxi is restricted to the F. zavjalovi– T. minutus Zone. The disappearance of several typical Palaeozoic, Permian or Changhsingian genera at the top of the F. zavjalovi –T. minutus Zone marks its upper boundary. This is the strongest and most pronounced conchostracan turnover anywhere within the Permian, and it is well below the Permian– Triassic boundary and even below the FAD of the dicynodont Lystrosaurus. At that time, which was within the C. changxingensis –C. deflecta conodont zone (Fig. 1), no distinct changes in the low latitude marine faunas are seen except for temporary replacements of the warm water conodont faunas by cool water conodont faunas outside the equatorial belt. As seen in the Tunguska Basin, the strong conchostracan faunal changes above the Falsisca zavjalovi –Tripemphigus minutus Zone occur within the very thick Khungtukunian tuffs of the Siberian Trap. This is a global bioevent in continental facies, probably caused by climatic changes related to this huge explosive volcanic event that occurred just before the effusion of the Siberian Trap flood basalts. At that time, which was within the C. changxingensis– C. deflecta conodont zone (Fig. 1), no distinct changes in the low latitude marine faunas are seen except for temporary replacements of the warm water conodont faunas by cool water conodont faunas outside the equatorial belt.

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The difference between the low latitude fauna and northern high latitude faunas of the F. zavjalovi –T. minutus Zone are rather minor. Hemicycloleaia mitchelli is restricted to high northern and southern latitudes. Echinolimnadia mattoxi is restricted to northern high latitudes. All other genera are widespread and occur both in low and northern high latitudes. Falsisca turaica –Falsisca zavjalovi Zone Definition. Co-occurrence of Falsisca turaica (Novozhilov) and F. zavjalovi (Novozhilov) without Bipemphigus, Megasitum, Tripemphigus and F. eotriassica Kozur. Lower boundary. Disappearance of the important Palaeozoic, Permian and Changhsingian genera Megasitum, Bipemphigus, Tripemphigus, and (in high latitudes) Hemicycloleaia and (in high northern latitudes) Echinolimnadia; this is the strongest conchostracan turnover within the Permian. Upper boundary. eotriassica.

Appearance

of

Falsisca

Important conchostracan species. Falsisca turaica (Novozhilov); F. zavjalovi (Novozhilov) (¼ Falsisca beijiangensis Liu and F. semicircularis Liu); F. qitaiensis Liu; F. dalongkouensis Liu; Beijianglimnadia ? rotunda Liu, 1989 (¼ B. ? multilinearis Liu 1989). Type locality. Dalongkou, on the southern limb of the anticline, 111–132 m above the base of the Guodikeng Formation. Occurrence. Middle Guodikeng Formation of Dalongkou and Xiaolongkou (both Xinjiang), lower part of upper Changhsingian. Upper Khungtukun tuffs of the Tunguska Basin. Age. The F. turaica –F. zavjalovi Zone is a short zone below the upper Changhsingian E. eotriassica Zone. As the latter zone begins distinctly above the base of the upper Changhsingian, the F. turaica– F. zavjalovi Zone belongs to the early part of the late Changhsingian. Remarks. After the extinction of most Permian species at the end of the F. turaica–F. zavjalovi Zone, a late Changhsingian to Gangetian low-diversity conchostracan fauna begins. Falsisca eotriassica Zone Definition. Range Zone of Falsisca eotriassica Kozur & Seidel. Lower Boundary. First appearance of Falsisca eotriassica.

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Upper boundary. Disappearance of F. eotriassica and F. zavjalovi. Important conchostracan species. F. eotriassica Kozur & Seidel; Palaeolimnadia aff. cishycranica (Novozhilov), only in the lower part of the zone but common; Euestheria gutta (Lyutkevich); Falsisca bolodekitensis (Novozhilov); Falsisca postera Kozur & Seidel, rare in the upper part of the zone; Falsisca zavjalovi (Novozhilov). Type locality. Quarry at Caaschwitz, La¨useberg, in Thuringia, Germany. Upper Bro¨ckelschiefer, uppermost Zechstein. Occurrence. Upper Bro¨ckelschiefer (uppermost Zechstein) of the Germanic Basin, including the uppermost part of the short reversed palaeomagnetic interval below the long normal interval that straddles the Permian–Triassic boundary and the lower part of the long normal interval. Dalongkou (Xinjiang), on the southern limb of the anticline from 132 m to c. 180 m above the base of the Guodikeng Formation, slightly above the FAD of F. postera at 171.2 m above the base of the Guodikeng Formation. Lower part of the co-occurrence interval of Dicynodon and Lystrosaurus. Upper part of the upper Khungtukunian tuffs and possibly also in sedimentary intercalations within the lower Putorana flood basalt of the Siberian Trap of the Tunguska Basin and Taimyr (only Falsisca bolodekitensis is reported from the latter beds). Also possibly in the uppermost Permian of NE Siberia (only P. cishycranica is reported). Age. Late Changhsingian below the base of the Buntsandstein and the age-equivalent base of the Boundary Clay at Meishan, central and northwestern Iran, and in Transcaucasia. Remarks. Kozur & Seidel (1983b) introduced this zone as the lower subzone of the F. eotriassica Zone s.l. that comprised the entire Falsisca fauna of the Germanic Basin from the base of the F. eotriassica Zone s.s. to the disappearance of the genus Falsisca. Kozur (1993) later regarded the F. eotriassica Zone as an independent zone. The index species is very similar to Falsisca bolodekitensis (Novozhilov) found in sedimentary intercalations within the lower Putorana basalt. However, F. bolodekitensis displays much narrower and much more numerous growth bands than F. eotriassica. F. bolodekitensis is rarely present in the eotriassica Zone of the Germanic basin and was assigned by Kozur & Seidel (1983a, pl. 1, fig. 6) to F. eotriassica. The occurrence of F. bolodekitensis in the (upper) eotriassica Zone of the Germanic Basin may indicate that this zone ranges perhaps up to the level of the lower Putorana flood

basalts of the Siberian Trap together with Lystrosaurus. However, Falsisca zavjalovi is also present in the F. eotriassica Zone of the Germanic Basin. In the Tunguska Basin, F. zavjalovi is common in the Khungtukun tuffs, but it does not occur as high as the sediments intercalated within the Putorana flood basalts. Thus, the F. eotriassica Zone (with F. zavjalovi) of the uppermost Zechstein of the Germanic Basin is no younger than the tuffs below the Putorana flood basalts. This is confirmed by the presence of F. podrabineki in the sediments intercalated within the Putorana flood basalts (Sadovnikov & Orlova 1993), because F. podrabineki is restricted to the F. postera Zone of the lowermost Buntsandstein in the Germanic Basin. Thus, the distinct climatic change at the Zechstein –Buntsandstein boundary, which coincides with the base of the Boundary Clay at the base of the C. meishanensis –H. praeparvus Zone (Kozur 1998a, b; Bachmann & Kozur 2004), also coincides with the beginning of the eruption of wide-spread flood basalts in the Siberian Trap (Fig. 1). Falsisca postera Zone Definition. Range of F. eotriassica.

F.

postera

without

Lower boundary. Disappearance of Falsisca eotriassica Kozur & Seidel, F. zavjalovi (Novozhilov) and probably F. bolodekitensis (Novozhilov); the last species may range a little higher. Upper boundary. Disappearance of F. postera, which remains common even into the uppermost F. postera Zone. The FAD of Falsisca verchojanica (Molin) usually follows a distinct Falsisca-free interval, so this FAD generally cannot be used for defining the top of the F. postera zone. In the southern limb of the anticline at Dalongkou, the youngest F. postera was found 199 m above the base of the Guodikeng Formation, the FAD of F. verchojanica is at 210 m above the base of the Guodikeng Formation, and the LOD of Dicynodon is at 219 m above the base of the Guodikeng Formation. These distances represent only very short time intervals because the sedimentation rate at Dalongkou is extremely high. An exception to this pattern of succession is found in the Zachełmie quarry east of the village Zagnan´sk, north of Kielce, Holy Cross Mountains. There is only a 1.4 m-thick interval without Falsisca that separates the highest occurrence of F. postera from the lowest occurrence of F. verchojanica. As continental beds generally do not yield conchostracans or other fossils in every layer, this can be regarded as an essentially continuous succession from the F. postera zone to the F. verchojanica Zone. The first conchostracans

TRIASSIC CONCHOSTRACANS

from this section were reported by Ptaszyn´ski & Niedz´wiedzki (2004, 2006) and F. postera was correctly assigned to the uppermost Permian in Ptaszyn´ski & Niedz´wiedzki (2006). In addition to F. postera, juvenile F. postera were incorrectly reported as F. eotriassica. F. verchojanica was not found in these first studies, but it was later reported in Kuleta et al. (2007). Important conchostracan species. Falsisca postera Kozur & Seidel, in some beds abundant; F. podrabineki (Novozhilov), rare; Euestheria gutta (Lyutkevich), dominant; E. oertlii Kozur, very rare; E. jakutica (Novozhilov), very rare. Type locality. Section Nelben (formerly called the ‘old quarry close to the Saale bridge near Ko¨nnern’), NW of Halle, Sachsen-Anhalt, Germany. Cycle 1 and basal part of Cycle 2 below the horizon of the alpha 2 oolitic limestones. Occurrence. Cycle 1 and lower part of cycle 2 (below the horizon of the alpha 2 oolitic limestones) of the Calvo¨rde Formation (lower Graubank Horizon s.s.) of the lowermost Buntsandstein in central and northern Germany. The Clayey Sandstone member (except its uppermost part) of the Calvo¨rde Formation of Thuringia. Lower threefourths of the Jaworzna Formation of the basal Buntsandstein in the abandoned Zachełmie quarry on the western part of Chełm Hill east of the village of Zagnan´sk, about 8 km north of the town of Kielce, Holy Cross Mountains. This occurrence lies at the southeastern margin of the northern Germanic Basin and is especially important because the F. postera Zone is separated from the overlying F. verchojanica Zone of the basal Triassic by only a 1.4 m-thick Falsisca-free interval. Basal Werfen Group of the Balaton Highland, Hungary. Upper, but not uppermost part of the co-occurrence interval of Lystrosaurus and Dicynodon at Dalongkou (Xinjiang). F. postera is known with certainty from 171.2 m to 199 m above the base of the Guodikeng Formation on the southern limb of the anticline. The co-occurrence interval of Lystrosaurus and Dicynodon is from 161 m to 219 m above the base of the Guodikeng Formation. Putorana flood basalts of the Siberian Trap. In the Tunguska Basin Falsisca podrabineki is restricted to sedimentary intercalations between the Putorana flood basalts (Sadovnikov & Orlova 1993). In the Germanic Basin this species is restricted to the F. postera Zone. Therefore the Putorana flood basalts can be correlated with the F. postera Zone. Age. Late Changhsingian. The correlation of the postera Zone with the latest Permian shallow water H. praeparvus Zone in the lowermost Werfen Group of Hungary and the correlation of

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the overlying F. verchojanica Zone with the basal Triassic in northern and northeastern Siberia places the continental Permian–Triassic boundary between the F. postera and F. verchojanica zones. This was later confirmed by Korte & Kozur (2005), who found the first minimum of the negative excursion in d13Ccarb around the Permian–Triassic boundary in the lower oolite bank alpha 2 horizon immediately above the top of the F. postera Zone in the Nelben section. In marine sections this minimum is situated at the base of the Hindeodus parvus Zone, which defines the base of the Triassic. Remarks. The postera Zone was introduced by Kozur & Seidel (1983b) as a subzone of the F. eotriassica Zone s.l. At that time F. postera was regarded as a subspecies of F. eotriassica. in the Nelben section, both Falsisca postera and the F. postera zone (formerly subzone) were established by Kozur & Seidel (1983a, b). In this section the uppermost Zechstein and the cycle 1 and lower cycle 2 of the Calvo¨rde Formation are exposed. Two horizons with oolitic limestone occur, one in the lower part of cycle 1 and one in the lower part of cycle 2 (Szurlies 1998). The upper horizon of oolitic limestone was assigned to the ‘oolitic limestone beta’ by Kozur & Seidel (1983). However, Paul & Klarr (1988) later subdivided the alpha oolitic limestone horizon into two horizons, alpha 1 and alpha 2. The oolitic limestone horizon in the basal part of cycle 2 of the Nelben section does not correspond to the beta horizon, but rather to the alpha 2 horizon. This is the level at which F. postera disappears in the Nelben section. The lower part of cycle 2 is the highest and most prominent level of grey-coloured rock in the Graubank Horizon sensu Schulze (1969). Kozur (1998a, b) used the Graubank Horizon in the sense of Schulze (1969) and therefore put the top of the F. postera Zone (and with it the top of the Permian) within the upper Graubank Horizon. Since then, Bru¨ning (1986) restricted the Graubank Horizon to the lower part of cycle 2, and this restricted definition of the Graubank horizon is preferred (Szurlies 1999). Bachmann & Kozur (2004) placed the Permian–Triassic boundary at the base of the minimum of the d13Ccarb, which occurs at the alpha 2 oolitic limestone horizon within the Graubank Horizon in the lower part of cycle 2 of the Nelben section (Korte & Kozur 2005). This coincides perfectly with the top of the F. postera Zone in this section and so it was chosen as the Permian –Triassic boundary by Kozur (1998a, b). Falsisca verchojanica Zone þ Falsisca cf. F. verchojanica Zone Definition. Range of Falsisca verchojanica (Molin) and of F. cf. F. verchojanica of the

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advanced F. verchojanica group. Falsisca cf. F. verchojanica (Molin) is a separate species, but for now we cannot describe this species because we have no suitable holotype. The well-preserved material from Xiaolongkou was confiscated and is unavailable for study [see section titled ‘Dalongkou (Xinjiang, northwest China)’]. F. cf. F. verchojanica from the Germanic Basin (see Kozur & Seidel 1983a) is clearly a different species from F. verchojanica, but all known specimens are either not fully preserved or deformed. Additionally, in the Chinese material there appear to be two different species that occur above the interval that contains F. verchojanica. Lower boundary. LOD of F. postera Kozur & Seidel, FAD of F. verchojanica (Molin). Upper boundary. The upper boundary of the F. verchojanica Zone s.s. is at the FAD of F. cf. F. verchojanica (¼ Falsisca n. sp.), illustrated by Kozur & Seidel (1983a, pl. 3, figs 6 & 7) under the name Falsisca eotriassica n. subsp. The upper boundary of the Falsisca cf. F. verchojanica Zone is characterised by the disappearance of the genus Falsisca Novozhilov, which is at the same horizon where the tetrapod Lystrosaurus disappears. Important conchostracan species. Falsisca verchojanica (Molin); Euestheria gutta (Lyutkevich), rare; E. oertlii Kozur. In the F. cf. F. verchojanica Zone the latter two species and Falsisca n. sp. (¼ F. cf. F. verchojanica ) are present. Type locality. Dalongkou (Xinjiang), on the southern limb of the anticline. Uppermost Guodikeng Formation (from G 210 m to G 232.7 m) and lower Jiucaiyan Formation. Gangetian (lower Brahmanian, basal Triassic). A latest Changhsingian age cannot be excluded for the lowermost F. verchojanica Zone below the level where Dicynodon disappears (in Dalongkou between 210 m and 219 m above the base of the Guodikeng Formation). Occurrence. Uppermost Guodikeng Formation and lower Jiucaiyan Formation of Dalongkou, Xinjiang. Together with the F. cf. F. verchojanica Zone in the lower 200 m of the Jiucaiyan Formation of Xiaolongkou, Xinjiang. Lowermost Triassic of northeastern Siberia. Lowermost Triassic of the Russian Platform. The Falsisca cf. F. verchojanica Zone occurs in the variegated lower part of the sandy Claystone member of the Calvo¨rde Formation in Thuringia (Gangetian Substage of the Brahmanian Stage) and in the middle and upper Jiucaiyan Formation of Xiaolongkou, Xinjiang. Uppermost Jaworzna Formation of the basal Buntsandstein above the F. postera Zone (see under this zone) in the abandoned Zachełmie quarry on the western

part of Chełm Hill east of the village of Zagnan´sk, about 8 km north of the town of Kielce, Holy Cross Mountains. F. verchojanica also occurs in sedimentary intercalations between the Marininskii flood basalts of Taimyr (Sadovnikov & Orlova 1993). Age. Gangetian Substage of the Brahmanian (Induan) Stage. The base of the F. verchojanica Zone was first correlated with the base of the Triassic at the FAD of H. parvus by Kozur (1998a, b), but the lowermost part of the F. verchojanica Zone that overlaps with the last occurrences of Dicynodon at Dalongkou could possibly belong to the latest Changhsingain. Falsisca verchojanica in Siberia always has been reported as coming from the basal Triassic, but the basis for this conclusion is rather indirect. In the Verchojanian Range in NE Siberia, F. verchojanica occurs only along its western slope (pers. comm. Dr N. G. Sadovnikov, Moscow), whereas Otoceras and lowermost Triassic ammonoids are found only along its eastern slope. In the Germanic Basin, the minimum d13Ccarb value lies in the Nelben section immediately above the top of the F. postera Zone (Korte & Kozur 2005), but in Nelben F. verchojanica has not yet been found immediately above the last F. postera. In the Holy Cross Mountains, F. postera is common up to the top of the F. postera zone, and 1.4 m above this zone F. verchojanica is present. The correlation of the postera/verchojanica boundary in the Nelben section with the marine Permian– Triassic boundary ultimately was confirmed by establishing the relationships between magnetostratigraphy, magnetosusceptibility and cyclostratigraphy (MSEC) in the Nelben section by Bachmann, Hansen, Szurlies, and Toft. These unpublished results, by comparison with the Meishan (south China) GSSP for the base of the Triassic, Jimsar (Xinjiang) and other sections, indicate that the location of the PTB is within the Graubankzone s.s., in the lower part of Cycle 2 in the alpha 2 oolitic limestone horizon of the Nelben section (pers. comm. from Prof. G. Bachmann, Halle and Dr H. J. Hansen, Copenhagen whom we thank very much for this information and permission to cite it here as unpublished data). Remarks. The equivalents of both the F. postera Zone and the F. verchojanica Zone in the lowermost Buntsandstein can be found in the Tunguska Basin and in Taimyr in sedimentary intercalations within the Putorana and Marininskii flood basalts, respectively. Therefore, the Permian–Triassic boundary lies within the flood basalts of the Siberian Trap. In undoubtedly Lower Triassic beds (e.g. in the uppermost Guodikeng Formation above the last occurrence of Dicynodon in Dalongkou, in the

TRIASSIC CONCHOSTRACANS

Jiucaiyan Formation of Dalongkou and Xiaolongkou (both Xinjiang), as well as in the lower Sandy Claystone Member of the Calvo¨rde Formation of the lower Buntsandstein in Thuringia) advanced very large (.7 mm), slender Falsisca occur which belong to the F. verchojanica group. In general they have a very fine reticulation, which unfortunately is usually not recognizable or only visible in the posteroventral part of the carapace. On the basis of the latter feature, Liu (1987a) established the genus Difalsisca, which is here regarded as a junior synonym of Falsisca. In all investigated sections, the underlying F. postera zone is separated from this zone by a short interval in which no Falsisca are present. In the rapidly deposited Dalongkou section, the interval between the last F. postera and the first Falsisca verchojanica group is very short (above G 199 m and below G 210 m). In this section, the lowermost part of the F. verchojanica Zone þ F. cf. verchojanica Zone (lower 9 m of more than 200 m) may belong to the uppermost Permian because in this interval are the highest occurrences of the Permian tetrapod genus Dicynodon. Most of the F. verchojanica Zone þ F. cf. F. verchojanica Zone coincides with the range of Lystrosaurus without Dicynodon. Most of the concurrent range zone of Lystrosaurus and Dicynodon can be assigned to the Changhsingian [see section titled ‘Dalongkou (Xinjiang, NW China)’], and the Lystrosaurus fauna without Dicynodon certainly belongs to the Lower Triassic. The upper range of the F. verchojanica Zone þ F. cf. F. verchojanica Zone is well established in the Xiaolongkou section (Xinjiang), where it ends about 200 m above the base of the Jiucaiyan Formation (there about 300 m thick). This level coincides with the highest occurrence of Lystrosaurus. Nowhere is Falsisca known to range into the Gandarian (Dienerian), Lystrosaurus also probably does not range higher than the Gangetian, which would make its total range upper Changhsingian through Gangetian. The time of disappearance of the conchostracan Falsisca and the tetrapod Lystrosaurus coincides with the time at which (warm and) cold-water adapted conodonts such as Hindeodus disappeared worldwide in the marine realm. This extinction event coincides with a strong warming that occurred in the uppermost Gangetian and at the base of the Gandarian (base of Dienerian). Lower Triassic conchostracan zonation. A Lower Triassic conchostracan zonation (Figs 13 –15) was established in the Germanic Basin by Kozur & Seidel (1983a, b) and later refined (Kozur 1993a, b; Bachmann & Kozur 2004; Kozur & Weems 2007; Kozur & Hauschke 2008; see Figs 13 & 14). This zonation can be traced throughout the entire northern hemisphere (wherever conchostracans are present)

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and across large parts of Gondwana. It is also closely correlated with the marine scale (e.g. Kozur & Seidel 1983a, b; Kozur 1993a, b, 1999; Kozur & Mock 1993; Bachmann & Kozur 2004; Kozur & Weems 2007). Among Triassic conchostracans, the taxonomy of the Lower Triassic forms is best known (e.g. Kozur & Seidel 1983a; Shen et al. 2002; Bachmann & Kozur 2004; Goretzki 2003; Kozur & Hauschke 2008). The F. verchojanica þ F. cf. F. verchojanica zones were discussed in the previous section. The Lower Triassic conchostracan zones, and their correlation with the Germanic Triassic and the marine scale, are shown in Figures 13 and 14. In Figure 15 the uppermost Permian to lower Smithian conchostracan zones are correlated with Milankovich cyclicity (short eccentricity cycles), the palaeomagnetic data of Szurlies (2004, 2007) and the d13Corg curve of Korte & Kozur (2005). Many of these conchostracan zones can be seen to have very short time spans. Molinestheria seideli Zone Definition. Range of Molinestheria seideli Kozur and Vertexia tauricornis Lyutkevich without Estheriella. Lower boundary. First occurrence of Molinestheria seideli Kozur and Vertexia tauricornis Kozur. Upper boundary. FAD of Estheriella bachmanni Kozur & Hauschke. Important conchostracan species. Molinestheria seideli seideli Kozur, very common (especially in the lower part of the zone); Vertexia tauricornis tauricornis Lyutkevich, common in the lower part of the zone; V. tauricornis transita Kozur & Seidel, common only in the middle and upper part of the zone; Cornia germari (Beyrich), absent in the basal part of the zone, then rare, and from the middle part of the zone increasingly common toward the top; Euestheria gutta (Lyutkevich), in the lower part of the zone locally common, above that very rare; E. oertlii Kozur, generally rare; Magniestheria ? malangensis (Marlie`re), rare; M. lerichi (Marlie`re), very rare. Type locality. Lindenschlucht at Su¨sser See, west of Halle/Saale (Sachsen-Anhalt, Germany). Upper Calvo¨rde Formation (upper part of eccentricity cycle 10 of the uppermost Calvo¨rde Formation) and lower Bernburg Formation. Lower Gandarian (lower Dienerian). Occurrence. Uppermost Calvo¨rde Formation (upper eccentricity cycle 10) and lower Bernburg Formation (eccentricity cycles 1 to lower

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Fig. 13. Conchostracan zonation of the Lower Buntsandstein of the Germanic Basin. This is the standard conchostracan zonation for the upper Changhsingian to lower Smithian interval in the Boreal realm, the low and middle latitudes of the northern hemisphere and in northern to central Gondwana. Vertical distances not to scale. Updated after Kozur & Weems (2007). The ranges of the index species and some selected other species are shown. For Magniestheria rybinskensis, only the range below the Volpriehausen Formation is shown. Scale ¼ 1 mm. Arrows indicate times of conchostracan immigration into the Germanic Basin following facies-controlled conchostracan-free intervals (Sabkha deposits) that are present throughout the entire Germanic Basin. Q.S., Quickborn Sandstein; 1, Falsisca eotriassica Kozur & Seidel; 2, Falsisca postera Kozur & Seidel; 3, Falsisca verchojanica (Novozhilov); 4, Molinestheria seideli Kozur; 5, Vertexia tauricornis Lyutkevich; 6, Estheriella bachmanni Kozur & Hauschke n. sp.; 7, Cornia germari (Beyrich); 8, Estheriella marginostriata Kozur; 9, Estheriella nodosocostata (Giebel); 10, Estheriella costata Weiss; 11, Magniestheria subcircularis (Chernyshev); 12, Magniestheria truempyi Kozur & Seidel; 13, Lioleaiina radzinskii Kozur & Seidel; 14, Magniestheria rybinskensis (Novozhilov), for illustration and upper range see Figure 14. See also explanation to Figure 15.

eccentricity cycle 4) of the northern and central Germanic Basin. Russian Platform (lower Krasnobakov podsvita of the upper Vochmin Formation, middle Kopan Formation), Pripjat Depression (middle Korenev Formation). Greenland (?), Xinjiang, Angola (?), Congo (?), India (?). In the last 4 areas, neither M. seideli nor V. tauricornis have been found. Only a few species such as C. germari are found there, and these occur also in the succeeding two zones. Age. Early Gandarian (early Dienerian) Substage of late Brahmanian (Induan) Stage. Remarks. The base of this zone in the Germanic Basin marks a time of immigration when conchostracans from outside the Germanic Basin appear after

the conchostracan-free upper but not uppermost Calvo¨rde Formation. Therefore, the zone could begin somewhat earlier outside the Germanic Basin, but not much earlier, because below the conchostracan-free upper (not uppermost) Calvo¨rde Formation Falsisca cf. F. verchojanica is still present. In Xinjiang (NW China) Molinestheria, Vertexia and real Cornia are not yet present in the Gangetian with Falsisca cf. F. verchojanica (Molin), as is the case in the Germanic Basin, though above the top of the Falsisca cf. F. verchojanica Zone and the contemporaneous top of the Lystrosaurus fauna at least Cornia is common. Thus, there is not a very long gap between the Euestheria –Falsisca fauna of the F. verchojanica– F. cf. F. verchojanica zones and the Vertexiidaedominated fauna of the Molinestheria seideli Zone.

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369

Fig. 14. Conchostracan zonation of the Middle (Volpriehausen and Solling formations) and Upper Buntsandstein (Ro¨t Formation) in the Germanic Basin. This standard conchostracan zonation applies to the Olenekian to lower Anisian of Eurasia and also the upper Spathian and lower Anisian of North and (partly) South America. The ranges and illustrations (except Hornestheria sollingensis Kozur & Lepper n. gen. n. sp.) of the index species and of some selected species are shown. Scale ¼ 1 mm. Updated after Kozur & Weems (2007). Vertical distances not to sale. QS, QuickbornSandstone; 13, Lioleaiina radzinskii Kozur & Seidel; 14, Magniestheria rybinskensis (Novozhilov); 15, Magniestheria mangaliensis (Jones); 16, Magniestheria deverta (Novozhilov); 17, Palaeolimnadia alsatica detfurthensis Kozur & Seidel; 18, Palaeolimnadia nodosa (Novozhilov); 19, Euestheria exsecta (Novozhilov); 20, Euestheria albertii mahlerselli Kozur & Lepper n. subsp., slender morphotyp; 21, Euestheria albertii mahlerselli Kozur & Lepper n. subsp., stout morphotyp; 22, Hornestheria sollingensis Kozur & Lepper n. sp., because of space limitations, only the range is shown; 23, Palaeolimnadia alsatica alsatica Reible; 24, Euestheria albertii albertii (Voltz); 25, Dictyonatella dictyonata (Reible).

Estheriella bachmanni Zone Definition. Occurrence of Estheriella bachmanni Kozur & Hauschke without radially ribbed Estheriella.

& Seidel; Cornia germari (Beyrich), common; Euestheria gutta (Lyutkevich), very rare; E. oertlii Kozur, generally rare; Magniestheria ? malangensis (Marlie`re), rare; M. lerichi (Marlie`re), very rare.

Lower boundary. FAD of Estheriella bachmanni Kozur & Hauschke.

Type locality. Large clay pit Beesenlaublingen near Alsleben, NW of Halle, Sachsen-Anhalt, Germany. At this locality all conchostracan zones from the Molinestheria seideli Zone up to the Estheriella nodosocostata Zone are exposed in superposition.

Upper boundary. FAD of the ribbed Estheriella marginostriata Kozur. Important conchostracan species. Estheriella bachmanni Kozur & Hauschke; Molinestheria seideli postera Kozur & Seidel; Vertexia tauricornis tauricornis Lyutkevich; V. tauricornis transita Kozur

Occurrence. Middle eccentricity cycle 4 to lower eccentricity cycle 5 of the central basin portion of the Bernburg Formation in central and northern

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Fig. 15. Correlation of the uppermost Permian to lower Smithian conchostracan zonation with the palaeomagnetic zones of Szurlies (2001), re-numbered by Bachmann & Kozur (2004), the short eccentricity Milankovitch cycles after Bachmann & Kozur (2004) and the carbon isotope record from lake limestones after Korte & Kozur (2005). After Kozur & Weems (2007). The biostratigraphic lower boundary of the Olenekian (B) at the base of the M. subcircularis A. Z. at the base of cycle 10 of the Bernburg Formation marks one of the two biggest conchostracan turnovers in the Triassic. According to the palaeomagnetic correlation with Chaohu in South China, the base of the Olenekian (A) lies deeper, within the Gandarian (Dienerian) conchostracan fauna.

Germany, for example, at Beesenlaublingen; borehole Halle-Su¨d, basal Kraftsdorf Sandstone of Kraftsdorf west of Gera, Thuringia, Germany; and at the Finne, northern Thuringia, Germany. Age. Early Gandarian (early Dienerian) of the Germanic Basin. Remarks. Estheriella bachmanni Kozur & Hauschke was originally assigned by Kozur & Seidel (1983a) to Polygrapta rybinskensis (Novozhilov). However, Goretzki (2003) pointed out that the type species of Polygrapta [Estheria (Polygrapta) chatangensis Novozhilov, 1946] has a free umbonal area (with a distinct sculpture) that is larger than the free umbonal area in Estheriella. Therefore, Estheriella bachmanni cannot belong to Polygrapta. Estheriella bachmanni lacks radial ribs, but in all other features it corresponds to Estheriella (see discussion in Kozur & Hauschke 2008). Since the radially-ribbed Estheriellacea Kobayashi clearly evolved from the unribbed species Estheriella bachmanni (as shown by

Kozur & Seidel 1983a) and not from radiallyribbed Late Paleozoic Leaiacea Raymond, the radially-ribbed Late Palaeozoic and Triassic conchostracans are not directly related to each other. Jones & Chen (2000) have argued that Estheriella is a bivalve, based on their study of the type material of Estheriella costata Weiss and E. nodosocostata (Giebel). The type material of these species is from the marginal part of the central basin facies, and it was found mainly in siltstones and sandstones that preserve only impressions of the shells and none of the original shell material. In Estheriella from claystones, the chitinous shell material is sometimes preserved, and it is just like the shell material of other conchostracans. Also, in these cases, there is no hinge present, as would be expected in bivalves. Estheriella marginostriata Zone Definition. Range of Estheriella marginostriata Kozur without Estheriella nodosocostata (Giebel). Lower boundary. FAD of Estheriella marginostriata Kozur & Seidel.

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Upper boundary. FAD of Estheriella nodosocostata (Giebel). Important conchostracan species. Estheriella bachmanni Kozur & Hauschke, common in the lower part of the zone, rare in the upper part of the zone and there mostly occurring as transitional forms to E. marginostriata Kozur; Estheriella marginostriata Kozur; Cornia germani (Beyrich), very common; Euestheria oertlii Kozur, very common; Molinestheria seideli seideli, rare; M. seideli postera Kozur, rare to common; Magniestheria ? malangensis Marlie`re, rare; Vertexia tauricornis tauricornis Lyutkevich, very rare; V. tauricornis transita Kozur & Seidel, common. Type locality. Outcrop at the railway station for Kraftsdorf. Kraftsdorf Sandstone of the middle Bernburg Formation (middle eccentricity cycle 5 of this formation). Gandarian (Dienerian). Occurrence. Lower Bernburg Formation (middle eccentricity cycle 5 of this formation, Gandarian ¼ Dienerian) of the Germanic Basin. Gandarian (Dienerian) of Russian Platform and of Greenland. Age. Early part of the late Gandarian (late Dienerian) Substage of Brahmanian (Induan) Stage. Remarks. Estheriella marginostriata is the oldest Estheriella species with radial ribs, and it occurs at the beginning of a very rapid period of evolution within Estheriella. Soon after its appearance, it evolved into E. costata Weiss, which in turn rapidly evolved into E. nodosocostata (Giebel). Therefore, the E. marginostriata Zone comprises only a very short time interval (duration less than half a short eccentricity cycle) within the middle part of eccentricity cycle 5 of the Bernburg Formation. Estheriella nodosocostata Zone Definition. Range of Estheriella nodosocostata (Giebel). Lower boundary. FAD of Estheriella nodosocostata (Giebel). Upper boundary. LOD of Estheriella nodosocostata (Giebel). As E. nodosocostata generally is rare in the upper part of the E. nodosocostata Zone, the boundary with the overlying Cornia germari – Mangiestheria subcircularis Zone is often difficult to recognize. Molinestheria and Vertexia do not occur as high as this zone, and Mangiestheria subcircularis does not appear until after the beginning of the overlying Cornia germari –Mangiestheria subcircularis Zone.

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Important conchostracan species. Estheriella nodosocostata (Giebel), in some beds as mass occurrences, especially in siltstones and fine sandstones, in claystones that have mass occurrences of Cornia germari rare or absent; E. costata Weiss (¼ E. sastryi Gosh), in some levels common; E. marginostriata Kozur, only in the lower part of the zone and rare; Cornia germari Beyrich, very common, except in beds that have mass occurrences of E. nodosocostata and there rare or absent; Euestheria gutta (Lyutkevich), very rare; Magniestheria lerichi (Marlie`re), very rare; M. ? malangiensis (Marlie`re), very rare; Molinestheria seideli postera Kozur & Seidel, common to rare; Vertexia tauricornis transita Kozur & Seidel, common to rare, but in the upper part of the zone mostly absent. Type locality. Large clay pit Beesenlaublingen near Alsleben, Sachsen-Anhalt, Germany. Uppermost eccentricity cycle 5 to lower eccentricty cycle 7 of the Bernburg Formation. Upper Gandarian (upper Dienerian). In case the beds with abundant conchostracans of the E. nodosocostata Zone eventually are removed during mining from the clay pit, the Lindenschlucht outcrop at Su¨sser See, west of Halle/Saale (Sachsen-Anhalt, Germany) is proposed as a parastratotype. E. nodosocostata is relatively rare in this outcrop, but there are several outcrops with rich occurrences of E. nodosocostata east of the Harz Mountains, such as the road cut from Unterrißdorf to Oberrißdorf, west of Halle, Sachsen-Anhalt, which has excellently preserved E. nodosocostata and all other species that occur in this zone. In all of these cases, the outcrops may disappear in the future. Occurrence. The central basin part of the central and northern Germanic Basin, uppermost eccentricity cycle 5 to lower eccentricity cycle 7 of the Bernburg Formation. The southernmost known occurrence in the Germanic Basin is in sandstones near Bad Salzungen, southern Thuringia. Russian Platform, northern Urals, Pricaspian depression, Hungary (thin brackish intercalation in the Claraia aurita Zone of the Seis Member (Aracs Marl Formation, Werfen Group), Timan (in northern Russia), Jakutia (northeastern Siberia), eastern Greenland, upper Feixianguan Formation, Langdai (Guizhou, China), Xinjiang, India (part of the Panchet Group), Gabon (Africa), eastern Africa and Australia. Outside of the Germanic Basin, Russian Platform and northern India, where occurrences of E. nodosocostata and (or) E. costata are widespread, the E. nodosocostata Zone is mainly recognizable by the occurrence of Cornia germari and less often by the occurrence of Magniestheria lerichi and M. ? malangiensis. All of these taxa may occur in the immediately underlying and

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overlying zones, so often only the relatively short interval from the upper Molinestheria seideli Zone up to the Cornia germari –Magniestheria subcircularis Zone can be recognised collectively. In the Germanic basin, this interval comprises 7 short eccentricity cycles (c. 700,000 years), shorter than the average duration of one Triassic ammonoid or conodont zone. Age. Late Gandarian (late Dienerian) Substage of Brahmanian (Induan) Stage. Remarks. The E. nodosocostata Zone can be readily recognised by the presence of the index species, but it also can be recognised by mass occurrences of Cornia germari (without Molinestheria seideli seideli and Vertexia tauricornis tauricornis). As Cornia germari has a world-wide distribution, the E. nodosocostata Zone can be recognised both in the northern hemisphere and in Gondwana. C. germari also occurs in brackish beds in marine deposits, where it often can be found in huge numbers on single bedding planes (e.g. in the Claraia aurita Zone of the Seis Member in Hungary, in the Vavilovites sverdrupi Zone in northeastern Siberia, and in brackish upper Gandarian beds in Greenland). These occurrences permit the E. nodosocostata Zone to be very closely correlated with the marine scale. As C. germari is very common in the interval from the upper M. seideli Zone up to the Cornia germari– Magniestheria subcircularis Zone, in marine beds and most of Gondwana, correlation can be made only collectively for these zones of short to very short duration. In China, Palaeolimnadia pusilla Shen is seemingly the same as C. germari. This species is accompanied there by Magniestheria ? lerichi (Marlie`re) (junior synonym: Euestheria langdaiensis Shen), and M. ? malangensis (Marlie`re) (junior synonym: Euestheria orbicula Shen). Euestheria leidayensis Shen from the same beds is based on deformed specimens, probably of M. lerichi. Cornia germari–Magniestheria subcircularis Zone Definition. Co-occurrence of Cornia germari (Beyrich) and Magniestheria subcircularis (Chernyshev). Lower boundary. Disappearance of Estheriella nodosocostata (Giebel). The appearance of Magniestheria subcircularis (Chernyshev) may be a little later. As E. nodosocostata is very rare in the uppermost part of the E. nodosocostata Zone and M. subcircularis is very rare in the lower C. germari –M. subcircularis Zone, the exact position of the boundary between these two zones is difficult to recognise and requires careful sampling,

though both zones are quite distinct away from this boundary zone. Important conchostracan species. Cornia germari (Beyrich), very common; Mangiestheria subcircularis (Chernyshev), in the lower part of the zone rare, in the upper part moderately common; Euestheria gutta (Lyutkevich), very rare; Magniestheria ? lerichi (Marlie`re), very rare. Type locality. Road cut at Marienburg, Niedersachsen, Germany, upper short eccentricity cycle 7 to top of cycle 9 of the upper Bernburg Formation. Occurrence. Widely distributed in the central basin facies of the central and northern Germanic Basin in central and northern Germany and Poland, upper Bernburg Formation, uppermost Gandarian (uppermost Dienerian). Claraia auritabearing Werfen Beds of Hungary (Kozur & Mock 1993), upper Gandarian (upper Dienerian). Age. Latest Gandarian (latest Dienerian). Remarks. The Cornia germari –Magniestheria subcircularis Zone is dominated by Cornia germari, which occurs often as monospecific faunas. In contrast to older Cornia germari-dominated faunas, other Vertexiidae (Vertexia, Molinestheria) and Estheriella are absent. Occasionally, especially in the upper part of the zone, some bedding planes with Magniestheria subcircularis are present. In the lower part of eccentricity cycle 7 of the upper Bernburg Formation, the youngest representatives of E. nodosocostata are found, but the species is very rare in this interval. In the upper part of eccentricity cycle 7, no Estheriella nodosocostata have been found and some transitional forms between Magniestheria ? malangiensis and M. subcircularis occur. This level currently is regarded as the lowest part of the C. germari –Magniestheria subcirularis Zone, but this assignment is not yet certain. Magniestheria subcircularis Zone Definition. Range of Magniestheria subcircularis (Chernyshev) without Cornia germari (Beyrich). Lower boundary. Disappearance of Cornia germari (Beyrich). Upper boundary. Appearance of Magniestheria truempyi (Kozur & Seidel). Important conchostracan species. Magniestheria subcircularis (Chernyshev), common; M. endybalica (Molin), very rare; Euestheria gutta (Lyutkevich), very rare; Lioleaiina triasiana (Chernyshev), rare; L. jakutica (Molin), very rare.

TRIASSIC CONCHOSTRACANS

Type locality. Road cut at Marienburg, Niedersachsen, Germany, short eccentricity cycle 10 of the upper Bernburg Formation. Occurrence. Upper but not uppermost Bernburg Formation of the Germanic Basin in the central basin facies (not on swells, such as the Eichsfeld Swell and its surroundings), southward to the Thuringian Basin. Kuzneck Basin of southern Siberia (part of the Malcev Formation), northeastern Siberia (Jakutia), eastern and central Africa. In northeastern Siberia, this zone can be firmly dated as basal Olenekian because Cornia germari is present there in the immediately underlying upper Gandarian (upper Dienerian) Vavilovites sverdrupi ammonoid zone. Age. Earliest Smithian Substage (earliest Olenekian Stage). Remarks. The sudden end of the Gandarian (Dienerian) conchostracan fauna, with its spined Vertexiidae, is one of the strongest extinction events that struck Triassic conchostracan faunas. The spined Vertexiidae (Cornia) disappeared suddenly in the upper Bernburg Formation within conchostracanrich sediments and without any gap between it and overlying faunas. In the Germanic Basin the M. subcircularis Zone is generally represented by monospecific faunas of the index species. Magniestheria truempyi Zone Definition. Range zone of Magniestheria truempyi (Kozur & Seidel). Lower boundary. FAD of Magniestheria truempyi (Kozur & Seidel). Upper boundary. LOD of Magniestheria truempyi (Kozur & Seidel). Important conchostracan species. Magniestheria truempyi (Kozur & Seidel). Type locality. Southern slope of railroad cut west of the Ertinghausen tunnel, about 2.75 km SE of Volpriehausen, Solling Mountains. TK 25 Nr. 4324; R: 35 52 960, H: 57 52 370. Rail km 43.974. Lower part of short eccentricity cycle 11 of the upper Bernburg Formation. Occurrence. In the Germanic Basin this zone occurs only in those parts of the central basin facies where the gap between the Bernburg Formation and the Volpriehausen Formation has a minimal duration, for example, around Halle (Sachsen-Anhalt, Germany), borehole Halle-Su¨d, west of Halle at Oberrißdorf (type locality of the index species), SW of Halle at Wangen where the

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M. truempyi Zone occurs in the sandy marginal facies of the uppermost Bernburg Formation. Solling Mountains (southern Niedersachsen, Germany) where the most complete exposure of the Bernburg Formation is found. Madagascar, in beds immediately below the Flemingites beds (which represent the second ammononid zone of the Olenekian) and above the Claraia beds (which represent the upper Gandarian). Age. Early Smithian Substage (Early Olenekian Stage). Remarks. Magniestheria truempyi is the type species of Magniestheria Kozur & Seidel. This genus, which is very characteristic of the Olenekian, and was originally established as a subgenus of Liograpta Novozhilov by Kozur & Seidel (1983b), but even then these authors pointed out that it had evolved from Euestheria Depere´t & Mazeran. Shen et al. (2002) have shown that Magniestheria is not closely related to Liograpta. They were able to demonstrate a close relationship with Euestheria, and regarded Magniestheria as a subgenus of Euestheria. Kozur & Bachmann (2004) later elevated Magniestheria to a genus. Magniestheria rybinskensis– Lioleaiina radzinskii Zone Definition. Occurrence of Magniestheria rybinskensis (Novozhilov) and Lioleaiina radzinskii Kozur & Seidel without Magniestheria mangaliensis (Jones). Lower boundary. FAD of Magniestheria rybinskensis (Novozhilov), which evolved from M. truempyi Kozur & Seidel. Upper boundary. FAD of Magniestheria mangaliensis (Jones), which evolved from M. rybinskensis (Novozhilov). Important conchostracan species. Magniestheria rybinskensis (Novozhilov), very abundant; Lioleaiina radzinskii Kozur & Seidel, mostly rare, in the Volpriehausen Sandstone sometimes common; Euestheria gutta gutta (Lyutkevich), very rare. Type locality. Southern slope of railroad cut west of the Ertinghausen tunnel, about 2.75 km SE of Volpriehausen, Solling Mountains, TK 25 Nr. 4324; R: 35 52 960, H: 57 52 370. Upper 5 m of short eccentricity cycle 11 of the upper Bernburg Formation and the overlying Volpriehausen Sandstone of the lower Volpriehausen Formation. Lower Smithian. Occurrence. Uppermost Bernburg Formation (upper half of short eccentricity cycle 11 of the Bernburg Formation in the centre of the Germanic Basin in central and northern Germany, Poland)

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and Volpriehausen Sandstone of the lower Volpriehausen Formation of the northern and central Germanic Basin, Russian Platform, Mangyshlak, Kuzneck Basin (southern Siberia). All lower (but not lowermost) Smithian (Lower Olenekian). Age. Early (but not earliest) Smithian Substage (Early Olenekian). Remarks. ‘Pseudestheria’ rybinskensis Novozhilov, 1960 originally was not well described, and the illustrations (drawings) could not be assigned to any Triassic genus with certainty. Goretzki (2003) re-studied the type material and has shown that ‘Pseudestheria’ rybinskensis Novozhilov, 1959 and ‘Lioestheria’ quellaensis Novozhilov, 1960 are synonyms. These forms belong to Magniestheria. For ‘Lioestheria’ quellaensis this was clear even from the original description and illustration, but not for ‘Pseudestheria’ rybinskensis, which has name priority. The base of the Magniestheria rybinskensis– L. radzinskii Zone can be readily recognised by the transition from M. truempyi to M. rybinskensis, which occurred in the Germanic Basin in the uppermost Bernburg Formation (but only in the central basin facies where the entire Bernburg Formation is preserved), in the Solling Mountains in the 5 m below the Volpriehausen Formation. In the Halle region the base of the M. rybinskensis–L. radzinskii Zone is located 1–3 m below the base of the Volpriehausen Sandstone, and in the Wangen section SW of Halle it is located only 0.5 m below the Volpriehausen Sandstone. In most regions of the Germanic Basin, the lower part of the M. rybinskensis –L. radzinskii Zone that is present in the uppermost Bernburg Formation has been removed by preVolpriehausen erosion. The Volpriehausen Sandstone also belongs to this zone and is widely distributed across the entire central and northern Germanic Basin. In the lower Magniestheria rybinskensis –L. radzinskii Zone of the Germanic Basin, a monospecific M. rybinskensis fauna generally is found, but in the Volpriehausen Sandstone Lioleaiina radzinkii may be common, for example, in the former clay pit Walpernhain, eastern Thuringia, Germany. Magniestheria mangaliensis Zone Definition. Stratigraphic range of Magniestheria mangaliensis (Jones). Lower boundary. FAD of Magniestheria mangaliensis (Jones). Upper boundary. FAD of Magniestheria deverta (Novozhilov). Important conchostracan species. Magniestheria mangaliensis (Jones), mass occurrences; M.

rybinskensis (Novozhilov), only in the lower part of the zone; Lioleaiina radzinskii Kozur & Seidel, only in the lower part of the zone and very rare. Type locality. Large clay pit Baalberge near Bernburg, NE of the Harz Mountains, Germanic Basin. Upper part of the Lower Olenekian (upper Smithian). Occurrence. Volpriehausen Formation above the Volpriehausen Sandstone of the Germanic Basin, Russian Platform, Mangyshlak, China, Mangli Beds of the Panchet Group in India, Angola. Upper Aracs Marl Formation of the Werfen Group in Hungary. Accompanying conodonts in the Werfen Group of Hungary (Pachycladina association) indicate Early Olenekian (late Smithian) age (the conchostracans were derived from brackish intercalations in the marine beds). Within the marine Werfen Beds of Hungary, in the Germanic Basin, on the Russian Platform and in Mangyshlak the M. mangaliensis Zone is overlain by Upper Olenekian (Spathian) deposits, confirming the late Smithian age of the M. mangaliensis Zone. Age. Late Smithian (late Early Olenekian). Remarks. M. mangaliensis is a very dominant form in this zone. If it is present, usually no other conchostracans occur. This is the case in such widely separated areas as India, the Germanic Basin and Hungary. This zonal index species occurs also in brackish beds (Germanic Basin, Werfen Group of Hungary), where it is the only species present and sometimes is found as mass assemblages. M. mangaliensis that have been reported from the Middle and Upper Triassic of South America (e.g. Geinitz. 1876; Gallego 1992) do not belong to this species; they are probably Euestheria buravasi (Kobayashi) which is also known from the lower Norian of Thailand. More than 20 years ago, Tasch (1987, p. 100) expressed doubts that these forms belong to M. mangaliensis. Magniestheria deverta Zone Definition. Stratigraphic range of Magniestheria deverta (Novozhilov), in the middle and upper part of the zone together with Euestheria exsecta (Novozhilov) and Palaeolimnadia alsatica detfurthensis Kozur & Seidel. Lower boundary. FAD of Magniestheria deverta (Novozhilov). Upper boundary. LOD of Magniestheria deverta (Novozhilov). Important conchostracan species. Magniestheria deverta deverta (Novozhilov), very common; Magniestheria deverta bogdoensis (Lopato), in the

TRIASSIC CONCHOSTRACANS

lower part common; Euestheria exsecta (Novozhilov), very common in the middle and upper part of the zone; Palaeolimnadia alsatica detfurthensis Kozur & Seidel, mostly common in the middle and upper part of the zone; P. ? cf. P. mecsekensis Nagy, very rare. Type locality. Beesenstedter Grund, southeastern foreland of Harz Mountains. Detfurth Clay. Lower Spathian (lower part of Upper Olenekian). Occurrence. Lower Spathian of the Germanic Basin (uppermost Volpriehausen Formation, Detfurth Formation, lowermost Hardegsen Formation), Pricaspian Depression (Bolshoe Bogdo), northern Siberia (eastern Taimyr Mountains), and in the lower Val Badia Member of the Csopak Marl Formation (Werfen Group) of Hungary. Age. Early Spathian (early Late Olenekian). Remarks. All occurrences of the deverta Zone can be clearly assigned to the lower Spathian. At the Bolshoe Bogdo, in the Balaton Highland and in the eastern Taimyr Mountains, marine intercalations with lower Spathian ammonoids are present. In the Germanic Basin, the type locality contains abundant Pleuromeia megaspores (Talchirella daciae Antonescu & Taugourdeau-Lantz) and miospores, including mass occurrences of Densoisporites neijburgii (Schulz) Balme without Cycloverrutriletes presselensis Schulz; this association is characteristic of the lower Spathian. The type locality Beesenstedter Grund is generally assigned to the Detfurth Formation. Euestheria exsecta Zone Definition. Co-occurrence of Euestheria exsecta (Novozhilov) and Palaeolimnadia ? nodosa (Novozhilov) without Magniestheria deverta (Novozhilov) and Euestheria albertii mahlerselli Kozur & Lepper n. subsp. Lower boundary. LOD of Magniestheria deverta (Novozhilov). Upper boundary. FAD of Euestheria albertii mahlerselli Kozur & Lepper n . subsp. Important conchostracan species. Euestheria exsecta (Novozhilov), very abundant; Palaeolimnadia alsatica detfurthensis Kozur & Seidel; P. ? nodosa (Novozhilov), common; P. ? cf. P. ? mecsekensis Nagy, rare. Type locality. Outcrop at the cemetery of Leißling, Sachsen-Anhalt (Germanic Basin), lower, but not lowermost Hardegsen Formation (Hardegsen 2?). Lower Spathian (lower part of Upper Olenekian).

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Occurrence. Lower Spathian of the Germanic Basin (Hardegsen 2 –4), Pricaspian Depression (Bolshoe Bogdo), northern Siberia (eastern Taimyr Mountains), middle and upper Val Badia Member of the Csopak Marl Formation (Werfen Group) of Hungary. Wupatki Member of Moenkopi Formation, near Meteor Crater, Arizona. Age. Early, but not earliest Spathian (early Late Olenekian). Remarks. Euestheria exsecta clearly dominates in the E. exsecta Zone, but it is a rather long-ranging form that is very abundant also in the middle and upper part of the underlying M. deverta Zone and ranges up to the uppermost Spathian Hornestheria sollingensis Zone. Therefore the E. exsecta Zone is an interval Zone. At the Bolshoe Bogdo and in Hungary, marine intercalations in the E. exsecta Zone contain Tirolites cassianus, which indicates an early Spathian age. Euestheria albertii mahlerselli – Palaeolimnadia ? nodosa Zone Definition. Co-occurrence of Euestheria albertii mahlerselli Kozur & Lepper n. subsp., E. exsecta (Novozhilov), Palaeolimnadia alsatica detfurthensis Kozur & Seidel and P. ? nodosa (Novozhilov). Lower boundary. FAD of Euestheria albertii mahlerselli Kozur & Lepper n. subsp. Upper boundary. FAD of Hornestheria sollingensis Kozur & Lepper n. sp. Important conchostracan species. Euestheria albertii mahlerselli Kozur & Lepper n. subsp.; E. exsecta (Novozhilov); Palaeolimnadia alsatica detfurthensis Kozur & Seidel; P. ? nodosa (Novozhilov). Type locality. Quarry Wu¨rgassen, TK 25 Nr. 4322; R: 3529350, H: 5723715, Solling Mountains, Trendelburg-Schichten of Solling-Formation. Occurrence. Lower part of Solling Formation up to Trendelburg Beds of the Solling Formation of the Solling Mountains (Germany). Age. Late Spathian (upper Upper Olenekian). Remarks. E. exsecta, P. alsatica detfurthensis and P. ? nodosa dominate in the lower part of the E. mahlerselli –P. ? nodosa Zone. E. albertii mahlerselli dominates in the upper part of the zone (upper Trendelburg Beds), but the other species of the zone are still present and sometimes even common. Hornestheria sollingensis Zone Definition. Range of Hornestheria sollingensis Kozur & Lepper n. sp.

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H. W. KOZUR & R. E. WEEMS

Lower boundary. FAD of Hornestheria sollingensis Kozur & Lepper n. sp. Upper boundary. LOD of Hornestheria sollingensis Kozur & Lepper n. sp., Euestheria exsecta (Novozhilov) and Palaeolimnadia ? nodosa (Novozhilov). Important conchostracan species. Hornestheria sollingensis Kozur & Lepper n. sp.; Euestheria exsecta (Novozhilov), rare; Palaeolimnadia alsatica detfurthensis Kozur & Seidel; P. ? nodosa (Novozhilov). Age. Latest Spathian. Karlshafen Beds of the Solling Formation. Remarks. Hornestheria is common in China and there restricted to the Anisian. Hornestheria species of China have been wrongly assigned to the unrelated Lower Permian genus Protomonocarina Tasch. Hornestheria ziguiensis (Shen) is very similar to H. sollingensis, but not identical to it. Because of the occurrence of Hornestheria, an Anisian age cannot be excluded for the H. sollingensis Zone, but the typical Late Olenekian species Euestheria exsecta, Palaeolimnadia alsatica detfurthensis and P. ? nodosa are still present in this zone, and all of these are missing in the overlying earliest Anisian Stammen Beds. Therefore, the conchostracans in the Karlshafen Beds of the Solling Formation indicate a somewhat ambiguous age, because guide forms of the Spathian occur together with Hornestheria, which is more typical of the Anisian. The palynological index species of the Anisian, Hexasaccites thiergartii (Ma¨dler) Kozur, does not occur in this zone and first appears considerably higher in the uppermost Solling Formation (Brugman 1986). However, this palynological index form has not been directly correlated with the pelagic Olenekian –Anisian boundary. Its lowest occurrence is in the lower Aszo´fo´´–Formation of the Balaton Highland, at the Spathian –Anisian boundary level, but these shallow marine beds have not been precisely correlated with the pelagic Olenekian –Anisian boundary. Even so, it is clear that Hexasaccites thiergartii begins close to the base of the pelagic Anisian, so the absence of this species in the Karlshafen Beds and its FAD in the overlying Stammen Beds of the uppermost Solling Formation strongly suggest a latest Spathian age for the Karlshafen Beds and the Hornestheria sollingensis-Zone within it.

1976). Modern taxonomic research has been conducted on the conchostracans in the lower Anisian of the Germanic Basin and in the Holbrook Member of the Moenkopi Formation of Arizona (e.g. Reible 1962; Gall 1971; Kozur 1982; Kozur & Seidel 1983b; Kozur, Lucas & Morales, in prep.), in the Anisian of northern Siberia (e.g. Novozhilov 1965) and in the upper Ladinian of the Germanic Basin (e.g. Warth 1969; Kozur & Seidel 1983a).The Middle Triassic conchostracans of China generally were not dated any more precisely than Middle Triassic, but Prof. Shen Yanbin (Nanjing) has sent us important data about marine intercalations that provide more detailed correlations for some of the Middle Triassic conchostracan faunas. The lower Anisian and upper Ladinian conchostracans of the Germanic Basin have been correlated in detail with marine faunas of the Ro¨t and Upper Muschelkalk through ammonoids, bivalves, etc. The upper Longobardian Erfurt Formation and Grabfeld Formation below the Estheria Beds of the upper Grabfeld Formation have only yielded a few marine and brackish intercalations, but all of these have yielded enough information to show that they are all Longobardian in age and that the base of the Carnian is at the base of the Estheria Beds. In northern Siberia the Anisian age of the conchostracan faunas has been determined by correlation with intercalated marine beds that contain ammonoids. For most of the Middle Triassic (the Pelsonian and Illyrian substages of the Anisian and the Fassanian substage of the Ladinian) no detailed correlations between Middle Triassic conchostracan faunas and the marine zonation can be made as yet. For the northern Siberian and Chinese faunas it is not possible to designate a type locality, and it is also difficult to establish which species except the index species also occur in these zones. For these reasons no type localities are designated, and accompanying species are only mentioned when they are important for the definition of the zone, for comparisons with other faunas or when the total fauna is known. In some cases, age and faunal composition of a zone are rather well known but the zone is only proven to exist within a limited area (e.g. within the Germanic Basin or China) and has yet to be found outside that area. For all of these reasons, we must regard the Middle Triassic conchostracan zonation as preliminary (Fig. 16).

Preliminary Middle Triassic conchostracan zonation

Euestheria albertii mahlerselli– P. alsatica alsatica Zone Definition. Co-occurrence of Euestheria albertii mahlerselli Kozur & Lepper n. subsp. and Palaeolimnadia alsatica alsatica (Reible).

The Middle Triassic conchostracan taxonomy has been well defined in China (e.g. Chang et al.

Lower boundary. Disappearance of Euestheria exsecta (Novozhilov), Hornestheria sollingensis

TRIASSIC CONCHOSTRACANS

Stage

Substage

Germanic Basin

China

Longobardian

Euestheria minuta

Euestheria minuta

Fassanian

Euestheria franconica

Euestheria franconica

Ladinian

Illyrian

Anisian

Conchostracan Zone

Pelsonian

Bithynian

Aegean

377

northern Siberia North America

Euestheria minuta

Xiangxiella bicostata

Xiangxiella bicostata

Diaplexa tigjanensis

Diaplexa tigjanensis

Diaplexa tigjanensis

Euestheria albertii albertii Euestheria albertii albertii Euestheria albertii Euestheria albertii mahlerselli-P. mahlerselli-P. alsatica alsatica alsatica alsatica

Euestheria albertii albertii

Euestheria albertii albertii

Euestheria a. mahlerselli-P. alsatica alsatica

Euestheria a. mahlerselli-P. alsatica alsatica

upper E. albertii albertii Zone

Fig. 16. Middle Triassic conchostracan zonation.

Kozur & Lepper n. sp., Palaeolimnadia alsatica detfurthensis Kozur & Seidel and P. ? nodosa (Novozhilov). Appearance of Palaeolimnadia alsatica alsatica (Reible). Upper boundary. Appearance of Euestheria albertii albertii (Voltz). Important conchostracan species. Euestheria albertii mahlerselli Kozur & Lepper n. subsp. and Palaeolimnadia alsatica alsatica (Reible). Occurrence. Upper Solling Formation (Stammen Beds in the Solling Mountains and correlative Thuringian Chirotherium Sandstone of Thuringia and the northern part of southern Germany) to middle Ro¨t Formation below the Dolomitische Grenzbank with Costatoria costata (Zenker) and Myophoria vulgaris (von Schlotheim). Solling Mountains, Thuringia, Franconia (Germany), all in the Germanic Basin. Lower Patacs Siltstone Formation of the Mecsek Mountains (Hungary), Aegean. Guizhou (China), lower Anisian. Age. Aegean (early part of early Anisian). Remarks. The E. albertii mahlerselli– P. alsatica alsatica Zone usually has a monospecific fauna of E. albertii mahlerselli. This species is rather longranging and occurs throughout the late Spathian and Aegean. Nevertheless, this zone is readily distinguished from the underlying uppermost Spathian H. sollingensis Zone by the absence of all Spathian

or late Spathian guide forms, such as E. exsecta, H. sollingensis, P. alsatica defurthensis and P. ? nodosa. Only P. alsatica alsatica first appears at the base of this zone; it is generally rare and ranges up to the Bithynian. The Anisian age of this zone is demonstrated by its co-occurrence with Hexasaccites thiergartii (Ma¨dler) Kozur. The first Bithynian faunal elements appear at the base of the overlying conchostracan zone. Euestheria albertii albertii Zone Definition. Occurrence of Euestheria albertii albertii (Voltz) without Diaplexa tigjanensis Novozhilov. Lower boundary. FAD of Euestheria albertii albertii (Voltz). Upper boundary. FAD of Diaplexa tigjanensis Novozhilov. Important conchostracan species. Euestheria albertii albertii (Voltz), very common; E. ? dactylis Shen, rare; Euestheria n. sp. C sensu Kozur et al. (1993), rare; Dictyonatella dictyonata (Reible), Lioleaiina n. sp. A (¼ Praeleaia sp. Warth 1969; Gall 1971), very rare; Magniestheria ? n. sp. A sensu Isaura sp. A Gall (1971), rare; Magniestheria n. sp. B sensu Kozur et al. (1993), rare; Palaeolimnadia alsatica alsatica Reible, common; P. mecsekensis Nagy, common; Palaeolimnadia n. sp. A sensu Kozur et al. (1993) ex gr. P. paucilinearis Shen, rare.

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Occurrence. Upper part of Upper Ro¨t Claystones (¼ upper part of Upper Variegated Member) and Myophoria Beds Member of Ro¨t Formation of Germany (E. albertii albertii has its FAD in the level with joint occurrence of C. costata and M. vulgaris at the base of the Bithynian), and Gre`s a` Voltzia of Lorraine (France), Germanic Basin, lower Bithynian. Upper Patacs Siltstone Formation and Magyaru¨ro¨g Anhydrite Member of the Mecsek Mountains (Hungary), lower and middle Bithynian. Holbrook Member of the uppermost Moenkopi Formation of Arizona, upper Bithynian. Lower Badong Formation, Hubei (China), Anisian. Lower Anisian of northern Siberia, there Euestheria albertii albertii was described as Estheria (Diaplexa) tigjanensis Novozhilov by Novozhilov (1946). Lower Anisian South America. Age. Bithynian (early Anisian). Remarks. The E. albertii albertii Zone was introduced by Kozur et al. (1993). The authors recognised the continuous increase in size of E. albertii albertii. This species first appears in the Germanic Basin in the middle part of the Upper Ro¨t Claystone of Franconia at a horizon where it co-occurs with Costatoria costata and Myophoria vulgaris. There, and in the upper part of the Upper Ro¨t Claystone, E. albertii albertii is very primitive, and its maximum length is only 4 mm. In the Myophoria Beds (Lower and Upper Dentritic Beds) of Franconia and in the time-equivalent Gre´s a` Voltzia of Lorraine, the maximum length of E. albertii albertii is 5 mm, and in the Holbrook Member (upper Bithynian) of Arizona the maximum length is 6–6.3 mm. In the upper Patacs Siltstone Member of the Mecsek Mountains in Hungary E. albertii albertii begins also with small forms that have a maximum length of 4 mm. In the uppermost Patacs Siltstone Member and in the Magyaru¨ro¨g Anhydrite Member, the maximum length increases to 5–5.5 mm. The FAD of M. vulgaris and E. albertii albertii within the Ro¨t Formation of the Germanic Basin coincides with the base of the Bithynian. Thus, during the Bithynian, the increase in the maximum length of E. albertii albertii was approximately 2 mm, or about 50% greater than the length of the most primitive forms. Diaplexa tigjanensis Zone Definition. Occurrence of Diaplexa tigjanensis Novozhilov without Xiangxiella bicostata Shen. Lower boundary. FAD of Diaplexa tigjanensis Novozhilov. Upper boundary. FAD of Xiangxiella bicostata Shen. Important conchostracan species. Diaplexa tigjanensis Novozhilov, 1946 (junior synoynym: Diaplexa ?

xuanhanensis Chen 1974), common; Palaeolimnadia triangularis Shen, common; Vileginia tuberculata (Novozhilov 1946) (junior synonym: Sedovia fecunda Novozhilov 1958); V. dorofeevi (Novozhilov) (junior synonym: Tigjanium borchgrevinki Novozhilov). Occurrence. Margin of the Laptev Sea in Siberia, upper Anisian (?middle Anisian). Badong Formation of Xuanhan (Sichuan, China), early Illyrian (late Anisian) or Pelsonian (middle Anisian). Age. ?Pelsonian, early Illyrian (late Anisian). Remarks. The much better documented Diaplexa ? xuanhanensis Chen, 1974 from the Anisian to lower Ladinian Badong Formation of Xuanhan, Sichuan (China) cannot be separated from D. tigjanensis Novozhilov, 1946 and, therefore, is regarded as a junior synonym of D. tigjanensis. Middle(?) and upper Anisian conchostracan faunas are characterized by coarsely reticulated Palaeolimnadiidae, which in the past have been assigned to different genera and even to different families. These are: Diaplexa Novozhilov 1946; Vileginia Novozhilov 1958; Tigjanium Novozhilov 1958; Sedovia Novozhilov 1958. Vileginia replaced Diaphora Novozhilov 1946, a homonym of Diaphora Stephens 1827 (in Southern & Nicklas 1827), Macquart, 1834, and Lo¨w, 1879; see Kobayashi 1954. Diaplexa has a small, flat, free umbonal area. Vileginia, Tigjanium and Sedovia have a large, flat, free umbonal area and cannot be separated from each other. Therefore, Sedovia and Tigjanium are regarded as junior synonyms of Vileginia. Whether reticulation appears as a reticulum or as convex tubercles depends on the preservation. Diaplexa and Vileginia may be separated on the base of the size of the flat umbonal area, but this is not a pronounced difference. The occurrences close to the Laptev Sea in northern Siberia are surely of Anisian age, as indicated by associated bivalves and a few ammonoids; the occurrence of Hungarites, Gervillia arctica Kiparisova, and two Trigonodus species probably indicates a late Anisian (Illyrian) age, but these marine faunal elements were not illustrated. Thus, a middle Anisian (Pelsonian) age cannot be excluded. The Chinese material is from the lower part of the Anisian to lower Ladinian (Fassanian) Badong Formation of Xuanhan, Sichuan. The probably late Illyrian or Fassanian genus Xiangxiella Shen was not reported from this level, and the lower Anisian guide form Euestheria albertii albertii also is missing. Thus, an early Illyrian or (and) Pelsonian age is probable for the occurrences of Diaplexa tigjanensis (¼ Diaplexa ? xuanhanensis Chen) from China.

TRIASSIC CONCHOSTRACANS

Xiangxiella bicostata Zone Definition. Range of Xiangxiella bicostata Shen. Lower Boundary. FAD of Xiangxiella bicostata Shen. Upper boundary. LOD of the genus Xiangxiella Shen. Important conchostracan species. Xiangxiella bicostata Shen (synonyms: Xiangxiella xilingxianensis Shen and probably also X. elongata Shen, based on a badly deformed specimen); X. acuta Shen; Euestheria hubeiensis Shen; E. shizibaoensis Shen; E. ? dactylis Shen; E. lepida Shen; Palaeolimnadia machaolingensis Shen. Occurrence. Upper Badong Formation, Hubei (China), upper Illyrian to Fassanian. Age. Upper Illyrian to Fassanian. Remarks. This zone can be readily recognized by the presence of Xiangxiella, which has on the free umbonal area two ridges that form an acute angle. Similar forms with only one prominent umbonal ridge were assigned to Protomonocarina Tasch by Shen in Chang et al. (1976), but they are not congeneric with the badly preserved Lower Permian genus Protomonocarina, in which the umbonal ridge is segmented and may represent an imprint of an appendage. The Middle Triassic forms of ‘Protomonocarina’ belong to Hornestheria Kozur & Lepper n. gen. The X. bicostata Zone is restricted to the upper Badong Fomation, Hubei (China). However, this probably does not mean that the X. bicostata Zone is an endemic fauna, because conchostracan faunas of late Illyrian–Fassanian age so far are unknown from any other part of the northern hemisphere or from Gondwana. In the Germanic Basin, this time interval is represented by marine or hypersaline beds without conchostracans. In North America this time interval is not represented by any known sediments. The age of the upper Badong Formation is rather well known. The conchostracans of the X. bicostata Zone occur only in the upper part of the Anisian to Fassanian Badong Formation, according to Prof. Shen (Nanjing, pers. comm.), above a marine horizon with Illyrian marine bivalves (Adygella illyrica). Thus, the X. bicostata Zone must be late Illyrian or Fassanian. Euestheria franconica Zone Definition. Range of Euestheria franconica (Reible).

379

Lower boundary. FOD of Euestheria franconica (Reible), Euestheria minuta (von Zieten) and Euestheria n. sp. Upper boundary. LOD of Euestheria franconica (Reible). Important conchostracan species. Euestheria franconica (Reible); E.minuta (von Zieten); Euestheria n. sp. Occurrence. This zone is restricted to brackish intercalations in the marine Upper Muschelkalk of the Germanic Basin above the Cycloides Bank. Age. Fassanian. Remarks. As this zone occurs only in brackish beds, it is possible that it overlaps with the Xiangxiella bicostata Zone and that Xiangxiella bicostata does not occur for facies reasons. However, it is also possible that the X. bicostata Zone is restricted to the upper Anisian. Euestheria minuta Zone Definition. Range of Euestheria minuta (von Zieten) without Xiangxiella and Laxitextella multireticulata (Reible). Lower Boundary. LOD of E. franconica (Reible). Upper boundary. FAD of Laxitextella multireticulata (Reible). Important conchostracan species. Euestheria minuta (von Zieten), dominant; Lioleaiina n. sp. A (¼ Praeleaia sp. Warth, 1969), very rare. Occurrence. Upper Ladinian of Eurasia from the Germanic Basin to China. Upper Ladinian of North Africa and Argentina. Upper Ladinian (lowermost Wolfville Formation) of the Fundy Basin, northeastern Canada. Age. Longobardian (late Ladinian). Remarks. The Euestheria minuta Zone is the youngest Triassic conchostracan Zone that occurs not only in the entire northern hemisphere, but also in parts of Gondawana (North Africa and Argentina). Despite this fact, it is not characterized by a very distinctive fauna, but rather by the absence of any distinctive taxa, such as Diaplexa, Vileginia, and Xiangxiella, which occur below E. minuta, and Laxitextella, which occurs above the E. minuta Zone. In most of these faunas, only E. minuta is present. The very rare Lioleaiina n. sp. also is known only from the middle Longobardian of the Germanic Basin (Erfurt Formation). As E. minuta is similar to other Euestheria species, often older or younger

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Euestheria species have been determined as E. minuta, such as the Spathian E. exsecta (by Chang et al. 1976), the lower Anisian E. albertii (e.g. by Reible, 1962) and some small Upper Triassic Euestheria. Real E. minuta have not been found below the E. franconica Zone and no true E. minuta occurs in the Anisian. E. hubeiensis Shen from the upper Illyrian to Fassanian X. bicostata Zone may be the forerunner of E. minuta. Euestheria winterpockensis (Bock 1953) (junior synonym: Cyzicus (Euestheria) minutus multicostatus Geyer 1987), the successor of E. minuta, as well as Laxitextella multireticulata (Reible), which also evolved from E. minuta, are common in the lower Carnian (Cordevolian). E. minuta occurs as well, at least a part of Cyzicus (Euestheria) minutus albertii (Voltz 1835) sensu Geyer (1987) persists into the Cordevolian of the Germanic Basin, and E. minuta occurs in the L. multireticulata Zone of the Newark Supergroup (Kozur & Weems 2007), where it is locally common.

now can be subdivided. The diagnostic species to do this, however, remain to be described. In addition to the problems cited above, the generic assignment of several species in the Tuvalian and Norian is not yet resolved because their phylomorphogenetic lineages have yet to be investigated. Thus, the Upper Triassic conchostracan zonation (Fig. 17) for now must remain preliminary, in some cases without type localities and often with incomplete knowledge of the species other than the index species that occur in these zones.

Cordevolian and Julian conchostracan zones and faunas Laxitextella multireticulata Zone Definition. Occurrence of Laxitextella multireticulata (Reible) without L. laxitexta (Sandberger). Lower Boundary. FAD of L. multireticulata and Euestheria winterpockensis (Bock).

Preliminary Upper Triassic and lower Hettangian conchostracan zonation Upper Triassic conchostracans are less well studied than Lower and Middle Triassic conchostracans. Only in a few intervals (e.g. the Cordevolian) is the taxonomy of Upper Triassic conchostracans well known, most forms described and their stratigraphic and regional ranges known. In the lower Julian, no conchostracans are known from anywhere in the world. During this time interval, a very significant sea-level drop produced a gap in all shallow water marine sedimentation areas (e.g. above the Wetterstein Formation and the Cassian Dolomite in the Northern and Southern Alps), and even continental areas with outflow to the sea show a break in deposition. Therefore no well dated lower Julian continental sediments with conchostracans are known. In contrast, very rich and diverse conchostracans faunas are known from the upper Julian Schilfsandstein, but all species have yet to be described. At least two different faunas are present. In the Tuvalian, four conchostracan zones can be recognized in the Germanic Basin and can be closely correlated with the marine scale. The upper two zones occur in direct superposition, but the other zones are separated by conchostracanfree hypersaline beds. In North America, the Newark Supergroup has yielded abundant conchostracan faunas in the Tuvalian interval, even some from between the conchostracan zones of the Germanic Basin, but most of these faunas either are not yet described or in need of revision, and several intervals remain to be sampled. In the Norian, several zones were established by Kozur & Weems (2005, 2007), but some of these zones

Upper Boundary. FAD of L. laxitexta. Important conchostracan species. Laxitextella multireticulata; Euestheria minuta (von Zieten); E. winterpockensis (Bock). Occurrence:. This zone occurs with the same species both in the lower variegated Estherienschichten of the upper Grabfeld Formation in the Germanic Basin and in the lowermost strata of the Newark Supergroup in the eastern United States. There, it occurs in the Irishtown Beds below the New Oxford Formation in the Gettysburg Basin (Maryland and Pennslyvania), in the Falling Creek Formation of the Taylorsville Basin (Virginia), in the Tuckahoe Formation of the Deep Run and Richmond basins (Virginia), and in the lacustrine unit in the Farmville, Briery Creek, and Scottsburg basins (Virginia). Age. Early Cordevolian. In the Germanic Basin, the FAD of L. multireticulata coincides with the FAD of the palynotaxon Patinasporites densus Leschik. Slightly higher is the FAD of Vallasporites ignacii Leschik. In the Carnian GSSP at Stuores Wiesen, these two species begin immediately above the base of the Carnian, as defined by the base of the Daxatina canadiensis Zone (Broglio Loriga et al. 1999). The beginning of this sporomorph association in the Germanic Basin, close to the base of the ‘Estheria’ Beds of the upper Grabfeld Formation, there defines the base of the Carnian (Schulz & Heunisch 2005). The correlation of this association with the base of the Carnian in the Germanic Basin was first recognized by Hauschke & Heunisch (1989). In the western part of the Southern

TRIASSIC CONCHOSTRACANS Stage

Substage

Conchostracan Zone Bulbilimnadia froelichi

Germanic Basin

Newark Supergroup Newark Basin other basins Bulbilimnadia Bulbilimnadia froelichi froelichi

Bulbilimnadia sheni

Hettangian

381 SW United States

Bulbilimnadia sheni

Bulbilimnadia killianorum

Bulbilimnadia killianorum

Euestheria brodieana

Euestheria brodieana

Gregoriusella polonica

Gregoriusella polonica

Bulbilimnadia killianorum uppermost E. brodieana Zone

B. killianorum uppermost E. brodieana Zone

Rhaetian Gregoriusella polonica

N. barnaschi - S. macdon N. barnaschi - S. macdon Shipingia olseni

Shipingia olseni

Shipingia olseni

Redondestheria grovetonensis

Redondestheria grovetonensis

R. grovetonensis

Sevatian

Norian

Shipingia hebaozhaiensis Alaunian

Lacian

Shipingia hebaozhaiensis

Shipingia hebaozhaiensis

Shipingia mcdonaldi

Shipingia mcdonaldi

Shipingia mcdonaldi

small Shipingia and large Euestheria

small Shipingia and large Euestheria

small Shipingia and large Euestheria

Euestheria buravasi – Euestheria n. sp. Palaeolimnadia schwanbergensis Laxitextella freybergiP. schwanbergensis Laxitextella freybergi

Euestheria buravasi – Euestheria n. sp. Palaeolimnadia Palaeolimnadia schwanbergensis schwanbergensis Laxitextella freybergiP. schwanbergensis Laxitextella freybergi Wannerestheria pennsylvanica H. ? ovata Laxitextella seegisi H. princetonensis

Euestheria buravasiEuestheria n. sp.

Tuvalian Laxitextella seegisi

Carnian

E. gallegoi

E. gallegoi Schilfsandstein faunas

Laxitexella laxitexta L. multireticulata

Laxitexella laxitexta L. multireticulata

Julian Cordevolian

Fulton site fauna/ W. pennsylvanica H. princetonensis

A. wingatella

L. multireticulata

Fig. 17. Conchostracan zones of the Upper Triassic and Hettangian.

Alps, transitional forms between Laxitextella multireticulata and L. laxitexta from the upper L. multireticulata Zone occur somewhat above the Ladinian – Carnian boundary, where another undescribed species of Laxitextella first appears (Tintori 1990; Tintori & Brambilla 1991; Kozur & Mock 1993). The L. multireticulata Zone represents the oldest conchostracan fauna in most of the Newark Supergroup of eastern North America. Only in the Fundy Basin (southeastern Canada) has a latest Middle Triassic conchostracan fauna been found in the lower Newark Supergroup. Remarks. In the lower L. multireticulata Zone, only typical specimens of the index species occur, with a reticulation composed exclusively of very small polygons that sometimes are arranged in irregular vertical or oblique stripes. Some of these polygons also may be elongated parallel to the growth

lines. In the upper part of this zone, in addition to typical forms, there also occur specimens that have some parts of the carapace covered with reticulations that are larger than normal and other parts covered with reticulations that are as small as in the type material. These advanced L. multireticulata are probably transitional forms to L. laxitexta. The upper part of the L. multireticulata Zone cannot be recognized in the Germanic Basin, because there the appropriate interval for the transition from the L. multireticulata Zone to the L. laxitexta Zone has soil horizons that have produced no conchostracans. At nearly all localities in the Newark Supergroup where the L. multireticulata Zone is found, these beds also produce the fish Dictyopyge (Kozur & Weems 2007). So far, only the Irishtown beds below the new Oxford Formation in the Gettysburg Basin (see Fig. 9) have failed to produce Dictyopyge or any other fish.

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Gregoriusella fimbriata–Laxitextella laxitexta Zone Definition. Occurrence of Gregoriusella fimbriata (Warth) and Laxitexella laxitexta (Sandberger). Lower boundary. FAD of Laxitextella laxitexta (Sandberger). Upper boundary. LOD of Gregoriusella fimbriata (Warth) s.s. Important conchostracan species. Gregoriusella fimbriata (Warth); Gregoriusella n. sp.; Laxitextella laxitexta (Sandberger). Occurrence. The Gregoriusella fimbriata– Laxitextella laxitexta Zone is known so far only from the Germanic Basin and England. Age. Late Cordevolian. Remarks. In the lower part of this zone, only Laxitextella laxitexta occurs along with some rare indeterminable small specimens of Euestheria. A small Gregoriusella n. sp. that is much less abundant than L. laxitexta first appears slightly higher, then in the middle and upper part of the zone the similarly small Gregoriusella fimbriata comes to dominate the assemblage. The upper boundary of this zone is difficult to define, because in the Germanic Basin it is followed by an interval barren of conchostracans that is particularly rich in pedogenic soil beds. A stratigraphic gap then follows. Above the gap, Gregoriusella cf. G. fimbriata and Laxitextella cf. L. laxitexta are present in the basal Schilfsandstein, but this fauna is poorly preserved. The lower Schilfsandstein above the basal beds is again dominated by Laxitextella of the L. laxitexta group, but Gregoriusella cf. G. fimbriata is no longer present or else very rare. A further change occurs in the upper Schilfsandstein, where in the Gaildorf Horizon a fauna occurs that contains many Laxitextella of the L. laxitexta group, among them a new species. Palaeolimnadia n. sp. is also common. This new species may be identical with Palaeolimnadia n. sp. from the lower Bluewater Creek Formation of northeastern Arizona. Forms similar to L. laxitexta persist upward to the middle Tuvalian. Two conchostracan zones can be established for the upper Julian after description of the conchostracan faunas of the Schilfsandstein in the Germanic Basin.

Tuvalian conchostracan zones of the Germanic Basin Eosolimnadiopsis gallegoi Zone Definition. Range Zone of Eosolimnadiopsis gallegoi Kozur.

Lower boundary. FAD of Eosolimnadiopsis gallegoi Kozur. Upper boundary. LOD of Eosolimnadiopsis gallegoi Kozur. Important conchostracan species. Eosolimnadiopsis gallegoi Kozur; Anyuanestheria n. sp. C Kozur; Laxitextella n. sp. C Kozur. Occurrence. Only known from the Germanic Basin, where it occurs in the basal Rote Wand of the Weser Formation (base of the Upper Gypsum Keuper) in southern Germany, and the equivalent Dolomie de Beaumont Horizon in eastern France. Age. Earliest Tuvalian, immediately above the middle Carnian wet intermezzo (MCWI) sensu Kozur & Bachmann (2008a). Remarks. The Dolomie de Beaumont Horizon of Eastern France contains the bivalve Costatoria vestita (von Alberti). Intercalated brackish marls and brackish to freshwater facies equivalent to the Dolomie de Beaumont yield conchostracans typical of the Eosolimnadiopsis gallegoi Zone. According to Prof. Renato Posenato, Ferrara, (pers. comm.) C. vestita occurs in latest Julian and earliest Tuvalian strata, and his specimens (sampled in situ) came from near the supposed Julian/Tuvalian boundary. Taking into consideration that the conchostracans of the E. gallegoi Zone have been derived from beds immediately above the middle Carnian wet intermezzo, this level can be assigned to the earliest Tuvalian. Laxitextella seegisi Zone Definition. Occurrence of Laxitextella seegisi Kozur. Lower boundary. FAD of Laxitextella seegisi Kozur. Upper boundary. LOD of Laxitextella seegisi Kozur and L. cf. L. laxitexta (Sandberger). Important conchostracan species. Gregoriusella bocki (Olempska); Krasiestheria parvula Olempska; Laxitextella cf. L. laxitexta (Sandberger); Laxitextella seegisi Kozur; n. gen. A n. sp. A. Occurrence. Lehrberg Beds of southwestern Germany (Seegis 1997); Krasiejo´w locality in Opole Silesia, southwestern Poland (Olempska 2004). Age. Later part of early Tuvalian to middle Tuvalian. The Lehrberg Beds have in marginal parts of the Germanic Basin a brackish to fresh water fauna, but in the deeper parts of the basin there also is a fauna from a slightly hypersaline –marine

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environment with dolomites and gypsum that includes Reubenella sp. Among euryhaline marine ostracods that also can live in brackish environments, Seegis (1997) mentioned and illustrated what he called Simeonella alpina (Bunza & Kozur 1971). However, all of his specimens represent Simeonella nostorica Monostori that clearly can be distinguished from the upper Julian Simeonella alpina. S. nostorica is common in the marine lower Tuvalian of Hungary and Austria, indicating a late early Tuvalian age for the Laxitextella seegisi Zone of the Lehrberg Beds. Remarks. Gregoriusella bocki (Olempska) and Krasiestheria parvula Olempska so far have been found only in the Krasiejo´w locality in Opole Silesia, southwestern Poland. The co-occurrence of Laxitextella cf. L. laxitexta and Laxitextella seegisi (¼ Laxitextella sp. A of Olempska 2004), which also are common in the Lehrberg Beds of southwestern Germany, indicates that both horizons have the same age. Kozur & Weems (2007) assigned the vertebrates from Krasiejo´w (Poland) to the lower Adamanian land vertebrate faunal chron (LVF) because the fauna includes Stagonolepis, the index taxon for the Adamanian, and also a few holdover Otischalkian taxa (e.g. Parasuchus and Metoposaurus). Lucas et al. (2007b) have since sought to define the base of the Adamanian by the LOD of Parasuchus and Metoposaurus, but this definition makes the Otischalkian –Adamanian boundary diachronous because these last Otischalkian elements occur only locally in the lower Adamanian where Stagonolepis is present. The timetransgressive nature of this boundary is readily demonstrated by conchostracan correlations, because the late Otischalkian (sensu Lucas et al. 2007b) conchostracan faunas of the L. seegisi Zone of Krasiejo´w correlate in North America with horizons that contain Adamanian land vertebrate faunas. Lucas (2010) concurs with this conclusion and has returned to defining the base of the Adamanian as the FAD of Stagonolepis and thus returning the famous Krasiejo´w fauna to the Adamanian as advocated by Kozur & Weems (2007) and removing the contradiction between conchostracan correlations and the land-vertebrate definition of the Otischalkian –Adamanian boundary in North America and Europe. According to Dr Sulej, (Warsaw, pers. comm.) the following vertebrates have been found at the Krasiejo´w site: Metoposaurus diagnosticus krasiejowensis Sulej; Cyclotosaurus intermedius Sulej & Majer; Teratosaurus silesiacus Sulej; Stagonolepis sp., Paleorhinus sp. (¼ Parasuchus sp.), and Silesaurus opolensis Dzik. The occurrence of Teratosaurus at a Tuvalian locality initially seemed surprising, because this genus was described

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originally from the Norian of southern Germany. However, according to Dr Sulej (pers. comm.), the Krasiejo´w Teratosaurus silesiacus is more primitive than the southern German species T. suevicus. In North America, the Anyuanestheria wingatella Zone from the ‘Lake Ciniza’ facies in New Mexico corresponds to the L. seegisi Zone. Very similar faunas occur also in the middle Cumnock Formation of the Deep River Basin (palaeoequatorial southern Newark Supergroup). Laxitextella freybergi Zone Definition. Range of Laxitextella freybergi Kelber & Kozur without Palaeolimnadia schwanbergensis Reible. Lower boundary. FAD of Laxitextella freybergi Kelber & Kozur. Upper boundary. FAD of Palaeolimnadia schwanbergensis Reible. Important conchostracan species. Laxitextella freybergi Kelber & Kozur; Laxitextella dorsorecta (Reible) emend.; Euestheria hausmanni (Schmidt); new species of Euestheria and Gregoriusella. Occurrence. Upper Tuvalian Coburg Sandstein of the Germanic Basin. Upper Tuvalian Fulton site of the Gettysburg Basin. Age. Late Tuvalian. The advanced Laxitextella species, transitional to Shipingia, indicate a late Tuvalian age for the Coburg Sandstein. According to Schulz & Heunisch (2005) and Heunisch (2005), the overlying Mainhardt Formation also contains Tuvalian sporomorphs. Thus, the Coburg Sandstein lies in the upper, but not the uppermost part of the Tuvalian Enzonalesporites vigens– Spiritisporites spirabilis Phase (van den Bergh 1987). This means that the Coburg Sandstein is late, but not latest Tuvalian in age (see also Bachmann & Kozur 2004). Remarks. The Laxitextella freybergi Zone has a diverse fauna. Most of the species are new and will be described by Kelber & Kozur (in prep.). Laxitextella freybergi– Palaeolimnadia schwanbergensis Zone Definition. Co-occurrence of Laxitextella freybergi Kelber & Kozur and Palaeolimnadia schwanbergensis Reible. Lower boundary. FAD of Palaeolimnadia schwanbergensis Reible. Upper boundary. LOD Laxitextella freybergi Kelber & Kozur and L. dorsorecta (Reible).

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Important conchostracan species. Laxitextella freybergi Kelber & Kozur; L. dorsorecta (Reible); Palaeolimnadia schwanbergensis Reible. Occurrence. So far only known from the Germanic Basin in gypsum-free intervals of the Heldburg Gypsum Member of the Weser Formation. Age. Latest Carnian (latest Tuvalian). According to Schulz & Heunisch (2005) and Heunisch (2005), the Mainhardt Formation contains the youngest Tuvalian sporomorphs. This formation is the marginal Germanic Basin equivalent of the Heldburg Gypsum Member. Remarks. The Laxitextella freybergi –Palaeolimnadia schwanbergensis Zone contains a typical upper Tuvalian fauna, the diversity of which is somewhat reduced compared with the Coburg Sandstone. Palaeolimnadia schwanbergensis also appears in this zone, and its monospecific faunas are characteristic of the immediately overlying lowermost Norian zone both in the Germanic Basin and in the Newark Basin (in the Warford Member of the Passaic Formation where the lowest Aetosaurus occurs in the Newark, Huber et al. 1993). At the upper boundary of the L. freybergi –P. schwanbergensis Zone, all late Tuvalian elements disappear. Correlatively, P. schwanbergensis becomes much more common.

Tuvalian conchostracan zones of North America Only zones are described from which rich material is present and the species present are at least partly described. Further zones will be described in a future paper, such as a lower Tuvalian Gregoriusella n. sp. Zone from the upper Stockton Formation. Anyuanestheria wingatella Zone Definition. Range of Anyanestheria wingatella (Tasch). Lower boundary. FAD of Anyuanestheria wingatella (Tasch) and Howellisaura princetonensis (Bock). Upper boundary. LOD of Anyuanestheria wingatella (Tasch) and Howellisaura princetonensis (Bock). Important conchostracan species. Anyuanestheria wingatella (Tasch); Howellisaura princetonensis (Bock); Laxitextella seegisi Kozur; Congestheriella elliptoidea (Bock). Type locality. Carnian ‘Lake Ciniza’ lake beds, Blue Water Creek Formation, McKinley County, New Mexico.

Occurrence. Carnian ‘Lake Ciniza’ lake beds, New Mexico, Adamanian LVF, middle Tuvalian. Middle Cumnock Formation, Deep River Basin (Sanford Subbasin), parts of the Dockum Group, Texas. Age. Adamanian LVF, middle Tuvalian. Remarks. This zone is very important for the correlation of different Tuvalian zones within North America and between North America and the Germanic Basin. Howellisaura princetonensis and Congestheriella elliptoidea allow a correlation with the H. princetonensis Zone of the lower Lockatong Formation in the Newark Basin. H. princetonensis is also common in the lower Cumnock Formation of the Durham Subbasin (Deep River Basin), though Anyuanestheria wingatella is not known yet from that horizon. It does, however, co-occur with H. princetonesis in the middle Cumnock Formation. The presence of Laxitextella seegisi allows correlation with the late early to middle Tuvalian L. seegisi Zone of the Germanic Basin. The following three zones occur in well-established superposition in the Tuvalian Lockatong Formation of the Newark Basin. The middle zone, the Howellisaura ? ovata Zone, is an endemic monospecific fauna, most probably related to a facies not normally well suited to conchostracans. According to van Houten (1962, 1964), parts of the Lockatong Formation were deposited in a natron lake, though this interpretation has been called into question by Smoot (2006). The low diversity conchostracan fauna of the lower quarter of the Lockatong Formation (up to the Princeton Member) obviously lived in a nearly normal fresh water lake. The species found there also occur outside of the Newark Basin. The succeeding monospecific Howellisaura ? ovata fauna, however, is only known from the Newark Basin. It appears abruptly in the Nursery Member of the Lockatong Formation and ranges up to the Skunk Hollow Member, where H. ? ovata (Lea) begins to co-occur with Wannerestheria pennsylvanica (Wanner). By the Smith Corner Member, W. pennsylvanica totally replaces H. ? ovata. W. pennsylvanica also is known outside the Newark Basin, but only from the adjacent and connected Gettysburg Basin. Thus, the fauna of this zone also appears to be strongly endemic. Howellisaura princetonensis Zone Definition. Occurrence of Howellisaura princetonensis Bock without Anyuanestheria wingatella (Tasch). Lower boundary. FAD of Howellisaura princetonensis Bock. Upper boundary. FAD of Howellisaura ? ovata (Lea) in the Newark Basin.

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Important conchostracan species. Howellisaura princetonensis Bock; Congestheriella elliptoidea (Bock). Occurrence. Lower fourth of Lockatong Formation up to the top of the Princeton Member of the Lockatong Formation in the Newark Basin. Lower Cumnock Formation of the Sanford Subbasin. Age. Adamanian LVF of middle Tuvalian age. Remarks. Kozur & Weems (2007) established originally a Euestheria princetonensis –E. ovata Zone because the exact range of these guide forms within the Lockatong Formation was not yet known. Howellisaura ? ovata Zone Definition. Occurrence of Howellisaura ? ovata (Lea) without Wannerestheria pennsylvanica (Wanner). Lower boundary. FAD of Howellisaura ? ovata (Lea) in the Newark Basin. Upper boundary. FAD of Wannerestheria pennsylvanica (Wanner). Important conchostracan species. Howellisaura ? ovata (Lea), often as monospecific faunas; Howellisaura princetonensis Bock, only in the lower part and rare. Type locality. Railroad cut at south end of Black Rock tunnel near Phoenixville, Pennsylvania (40.1438 N, 75.5129 W). Occurrence. Endemic fauna from the lower Lockatong Formation (Nursery Member to Byram Member) in the Newark Basin. Age. Middle Adamanian LVF. Middle Tuvalian. Remarks. A direct correlation with conchostracan faunas outside the Newark Basin is not possible for this endemic fauna. As the index species of the overlying Wannerestheria pennsylvanica Zone occurs together with L. freybergi in the late Tuvalian conchostracan fauna of the Fulton site in the Gettysburg Basin, and this site can be readily correlated with the L. freybergi Zone of the Germanic Basin, the H. ? ovata Zone must be older than these late Tuvalian conchostracan faunas and, of course, younger than the H. princetonensis Zone, which does occur outside of the Newark Basin. Wannerestheria pennsylvanica Zone Definition. Range of Wannerestheria pennsylvanica (Wanner). Lower boundary. FAD of Wannerestheria pennsylvanica (Wanner).

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Upper boundary. FAD of Palaeolimnadia schwanbergensis Reible. Important conchostracan species. Wannerestheria pennsylvanica (Wanner), strongly dominating or occurring as a monospecific fauna; Howellisaura ? ovata (Lea), only in the lower part of the zone (Skunk Hollow Member). Type locality. Lower Little Conewago Creek, York County, Pennsylvania, at and between locality 1 (40.0932 N, 76.7264 W) and locality 11 (40.0703 N, 76.7353 W) of Wanner (1926). Occurrence. Basal Gettysburg Formation of the Gettysburg Basin. Upper Lockatong Formation (Skunk Hollow Member to the top of the formation) in the Newark Basin. Age. Late Tuvalian. Remarks. Wannerestheria pennsylvanica (Wanner) first appears in the Skunk Hollow Member of the Lockatong Formation in the Newark Basin, where it co-occurs with Howellisaura ? ovata (Lea). Higher up in the upper Lockatong Formation, the conchostracan fauna changes to a monospecific W. pennsylvanica fauna, especially abundant in the Smith Corner Member. In the basal Gettysburg Formation of the Gettysburg Basin, rich monospecific W. pennsylvanica faunas are also common. Farther to the southwest, a much more diverse fauna occurs in the basal Gettysburg Formation at the Fulton site. This conchostracan fauna is important for correlation with other conchostracan faunas because it is diverse and contains the following species: Euestheria hausmanni (Schmidt), Laxitextella freybergi Kelber & Kozur, Wannerestheria pennsylvania (Wanner), and n. gen. A n. sp. B. This fauna probably can be assigned specifically to the lower part of the L. freybergi Zone, because n. gen. A occurs in the Germanic Basin only below the Coburg Sandstone (in which occurs the type L. freybergi Zone). The Fulton site fauna may well represent the upper part of the conchostracan-free interval between the L. seegisi Zone and the L. freybergi Zone in the Germanic Basin. Therefore, the Fulton site fauna, because it contains the index species for both the L. freybergi Zone and the W. pennsylvanica Zone, is important because it establishes a correlation (or at least an overlap) between these two zones.

Norian, Rhaetian and lower Hettangian conchostracan zonation Palaeolimnadia schwanbergensis Zone Definition. Occurrence of Palaeolimnadia schwanbergensis Reible without Laxitextella

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freybergi Kelber & Kozur and other upper Tuvalian guide forms. Lower boundary. LOD of Laxitextella freybergi Kelber & Kozur. Upper boundary. FAD of Euestheria buravasi Kobayashi and Euestheria n. sp. aff. E. hausmanni (Schmidt). Important conchostracan species. Palaeolimnadia schwanbergensis Reible. Type locality. Along Warford Creek, just east of New Jersey State Road 29 (40.4694 N, 75.0611 W), about 200 m above the base of the Warford member of the Passaic Formation. Occurrence. Warford Member of basal Passaic Formation, lowermost Norian, which is also the lowest reported occurrence of Aetosaurus in the Newark Basin. Lowermost 10 m of the Arnstadt Formation (Steinmergelkeuper), in its central basin facies in the Germanic Basin where there is no gap between it and the underlying Weser Formation (Upper Gypsum Keuper). Basal Norian. In more marginal facies, where there is a time gap between the Weser Formation and the Arnstadt Formation, the base of the Arnstadt Formation begins with the next younger Euestheria buravasi-Euestheria n. sp. Zone. Age. In the Germanic Basin and in the Newark Basin, earliest Norian. Remarks. In the Germanic Basin, Palaeolimnadia schwanbergensis first appears in the latest Tuvalian, but it occurs there together with a typical latest Tuvalian fauna (L. freybergi and other late Tuvalian guide forms). Euestheria buravasi–Euestheria n. sp. Zone Definition. Occurrence of Euestheria buravasi Kobayashi and Euestheria n. sp. aff. E. hausmanni (Schmidt). Lower boundary. FAD of Euestheria buravasi Kobayashi and Euestheria n. sp. aff. E. hausmanni (Schmidt). Upper boundary. FOD of Shipingia of the Shipingia ? mansuyi (Kobayashi) group. Important conchostracan species. Euestheria buravasi Kobayashi; Euestheria n. sp. aff. E. hausmanni (Schmidt); Euestheria bunopasi (Kobayashi); Anyuanestheria n. sp. A. Type locality. Triangle Brick Quarry in the Durham Subbasin of the Deep River Basin, Durham County, North Carolina (35.8694 N, 78.8960 W).

Occurrence. Triangle Brick Quarry in the Durham Subbasin of the Deep River Basin (NC), (‘lithofacies Association II,’ equivalent to the uppermost Cumnock Formation in the adjacent Sanford Subbasin). Lower Arnstadt Formation (lower Steinmergelkeuper) above its lowermost part in Thuringia and northern Germany. Lower Norian of Thailand. Age. Euestheria buravasi was originally described from the early Norian of Thailand (Kobayashi 1975). An earliest Norian age also is indicated by vertebrates from the Triangle Brick Quarry. Aetosaurus arcuatus (Marsh) in particular indicates an age no older than Norian (Lucas et al. 1998). The co-occurrence of a dicynodont at this quarry (Sues et al. 2001) in the past has been considered as evidence of a Carnian age for this deposit, but this is no longer true because a Norian (or possibly Rhaetian) dicynodont recently has been reported from Poland (Dzik et al. 2008). Thus, this particular tetrapod assemblage is completely compatible with an early Norian age. Interval with small Shipingia and large Euestheria Remarks. Between the lower Norian Euestheria buravasi–Euestheria n. sp. Zone and the middle Norian Shipingia hebaozhaiensis fauna, there is both in the Germanic Basin and in the Newark Supergroup a diverse conchostracan fauna that has not been investigated adequately. It consists of small Shipingia similar to Shipingia ? mansuyi (Kobayashi 1954), large Euestheria, in part previously described from the Norian of SE Asia but in need of taxonomic revision, and partly of new taxa. This fauna belongs to a new zone that can be only established after adequate taxonomic work is completed on this fauna. The age of this fauna is late Lacian to early Alaunian. Norestheria barnaschi–Shipingia mcdonaldi Zone Diagnosis. Range of Norestheria barnaschi Kozur & Weems n. sp. and Shipingia mcdonaldi Kozur & Weems n. sp. Lower boundary. FAD of Norestheria barnaschi Kozur & Weems n. sp. and Shipingia mcdonaldi Kozur & Weems n. sp. Upper boundary. FAD of large Shipingia hebaozhaiensis Shen. Important conchostracan species. Norestheria barnaschi Kozur & Weems n. sp.; Shipingia mcdonaldi Kozur & Weems n. sp.; Euestheria sp. Type locality. Borehole Malschenberg 1 (Etzold & Franz 2005). Basal Stubensandstein 3, uppermost Alaunian.

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Occurrence. Upper part of middle Arnstadt Formation in the central basin facies, lowermost Stubensandstein 3 in the more marginal facies of the Germanic Basin. In the Wachsenburg section (Thuringia), this zone occurs immediately below the Wachsenburg Sandstone of the middle (grey) Steinmergelkeuper (middle Arnstadt Formation). Heidlersburg Member of the Gettysburg Formation of the Gettysburg Basin (Pennsylvania). Sanford Formation of the Sanford Subbasin (Deep River Basin) (North Carolina).

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Supergroup, only the typical large S. hebaozhaiensis with well-developed microsculpture were known from the Newark Supergroup and China. We use the original concept of the S. hebaozhaiensis Zone. Kozur & Weems (2007) later assigned small to moderately large and more slender forms without distinct microsculpture to S. hebaozhaiensis, but in the present paper these forms are separated as a new species, S. mcdonaldi. They always occur below, and thus are older than, the large S. hebaozhaiensis s.s., which are restricted to the upper Alaunian, but they still are of Alaunian age.

Age. Middle Norian (Alaunian). Remarks. The N. barnaschi–S. mcdonaldi Zone comprises the lower part of the S. hebaozhaiensis Zone sensu Kozur & Weems (2007). The small to medium-sized S. hebaozhaiensis of the lower S. hebaozhaiensis Zone sensu Kozur & Weems (2007) are here considered to be an independent species, S. mcdonaldi n. sp. Often a monospecific fauna of S. mcdonaldi occurs, for example, in the Heidlersburg Member and in the middle Steinmergelkeuper immediately below the Wachsenburg Sandstone of western Thuringian Basin. Sometimes only N. barnaschi is present. In the Malschenberg borehole, Norestheria n. sp. occurs immediately above S. mcdonaldi. In North America N. barnaschi has not been found. Shipingia hebaozhaiensis Zone Definition. Range Zone of Shipingia hebaozhaiensis Shen s.s. (big forms). Lower Boundary. FAD of S. hebaozhaiensis Shen s.s. Upper Boundary. LOD of S. hebaozhaiensis Shen s.s. Important conchostracan species. Most localities yield monospecific faunas of big S. hebaozhaiensis, though occasionally undescribed species of Euestheria also occur. Occurrence. Lower middle part of the Groveton Member of the Bull Run Formation in the Culpeper Basin (Virginia), upper Alaunian; upper part of middle Arnstadt Formation (upper part of middle Steinmergelkeuper) of northern Germany, upper Alaunian; lowermost Stubensandstein 3 of Wu¨rttemberg (southwestern Germany), upper Alaunian; Morocco (material provided by Dr Paul E. Olsen, Columbia University), upper Alaunian; China, middle Norian (upper Alaunian). Age. Late Alaunian. Remarks. When Kozur & Weems (2005) established the S. hebaozhaiensis Zone in the Newark

Redondestheria grovetonensis Zone Definition. Range zone of Redondestheria grovetonensis Kozur & Weems. Lower boundary. FAD of Redondestheria grovetonensis Kozur & Weems. Upper boundary. LOD of Redondestheria grovetonensis Kozur & Weems. Important conchostracan species. Redondestheria grovetonensis Kozur & Weems; Shipingia olseni Kozur & Weems. Occurrence. Upper middle part of the Groveton Member of the Bull Run Formation at the Groveton locality and across an interval of at least nine lake beds east of Arcola, Culpeper Basin, Virginia, lower Sevatian. Age. Early Sevatian. Remarks. The R. grovetonensis Zone and the genus Redondestheria Kozur et al. have not been found outside of North America, probably because of provincialism. In North America, Redondestheria is found in the Newark Supergroup and in the Chinle Group of New Mexico. Shipingia olseni Zone Definition. Occurrence of Shipingia olseni Kozur & Weems in the absence of S. hebaozhaiensis Shen and Redondestheria Kozur, Weems & Lucas. Lower boundary. LOD of Redondestheria grovetonensis Kozur & Weems. Upper boundary. LOD of Shipingia olseni Kozur & Weems. Important conchostracan species. In the Newark Supergroup, mostly a monospecific S. olseni fauna occurs. Rarely, small Norian Euestheria are present (also known from China and Southeast Asia) that are not well studied in the Newark Supergroup. In the Germanic Basin, unlike in the Newark Supergroup, S. olseni is uncommon and several other species occur which have yet to be described.

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In the upper S. olseni zone this is also the case in the Newark Supergroup. In the southwestern United States several undescribed new species also are present in this zone, including a large Euestheria, a Shipingia and a new genus. Additionally, Redondestheria novomexicoensis Kozur, Weems & Lucas occurs, which is more advanced than R. grovetonensis. Type locality. Cedar Run section, Catharpin Creek Formation, Culpeper Basin. Occurrence. Uppermost Catharpin Creek Formation (Haymarket locality), middle Catharpin Creek Formation (Cedar Run locality), upper Groveton Member of the Bull Run Formation (northwest of Arcola), all in the Culpeper Basin (Virginia); Constitution Avenue section (Exeter, Pennsylvania), in the uppermost 21 –28 m of the Passaic Formation (according to measurements by Fowell & Olsen 1993), Boyertown Road section, middle Passaic Formation, both in the Newark Basin (Pennsylvania and New Jersey); upper Blomidon Formation in the Fundy Basin, (Nova Scotia, Canada); Duke Ranch Member of the upper Redonda Formation of New Mexico; upper Arnstadt Formation in northern Germany. Age. All occurrences of the S. olseni Zone have a Sevatian age, and this species occurs up to the top of the Sevatian. No Shipingia olseni have been found anywhere in the Rhaetian. Remarks. After the description of several new species that occur within the presently defined S. olseni Zone, this zone certainly can be divided into subzones. Gregoriusella polonica Zone Definition. Range of Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. Lower boundary. LOD of very large Shipingia olseni Kozur & Weems and the likewise very large Redondestheria novomexicoensis Kozur, Weems & Lucas. Upper boundary. LOD of Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. Important conchostracan species. Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. In the lower part of this zone, mostly monospecific faunas occur. In the upper part of the zone, Euestheria brodieana (Jones) becomes steadily more abundant and ultimately dominates the fauna. Type locality. Active clay pit of Lipie, Upper Silesia, Poland, Germanic Basin. Grey, plantbearing mudstones and siltstones with some

sandstone intercalations, lowermost Rhaetian. In the middle of the 4 m thick succession begins the Rhaetian sporomorph guide form Rhaetopollis germanicus Schulz. Occurrence. Lower Exter Formation of northern Germany, lower Rhaetian; southwestern Poland (Upper Silesia), lower Rhaetian. Upper Redonda Formation of New Mexico. Both in the Germanic Basin and in the upper Redonda Formation above the S. olseni Zone. Age. Early part of early Rhaetian. This zone possibly may begin in the uppermost Sevatian, but if so is not demonstrable at this time. Remarks. The similarly small E. brodieana can be distinguished from Gregoriusella polonica only by the microsculpture. Therefore, in faunas that are not well-preserved, the distinction between the G. polonica Zone and the E. brodieana zone may be difficult to discern. The base of the Olenekian and the base of the Rhaetian mark the two most sweeping turnovers in conchostracan faunas within the Triassic. The changes in the conchostracan faunas at the base of the Rhaetian are more striking in one way because the upper Norian faunas were dominated by very large conchostracans, while the Rhaetian (and Hettangian) conchostracan faunas are everywhere composed of very small forms. Euestheria brodieana Zone Definition. Range of E. brodieana (Jones), in the absence of both the lower Rhaetian Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. and the basal Hettangian genus Bulbilimnadia Shen. Lower boundary. LOD of Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. Upper Boundary. FAD of Bulbilimnadia killianorum n. sp. Important conchostracan species. Both in the Germanic Basin and in the Newark Supergroup, the E. brodieana Zone consists of monospecific E. brodieana conchostracan faunas. Occurrence. Upper Rhaetian of England, Germany, and France (Argiles de Levallois); upper part of lower Rhaetian in Germany; uppermost Rhaetian in uppermost Catharpin Creek Formation immediately below the first lava flow, Culpeper Basin. Age. In Europe, all published occurrences of the E. brodieana fauna have been in upper or uppermost Rhaetian strata. Recently, E. brodieana has shown up in the upper part of the lower Rhaetian. This

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earliest known occurrence is with Gregoriusella polonica, so this fauna is assigned to the upper Gregoriusella polonica Zone. Somewhat below the marine middle Rhaetian, E. brodieana totally replaces Gregoriusella polonica. Thus, the full range of the E. brodieana Zone is late early Rhaetian through late Rhaetian. Bulbilimnadia killianorum Zone Definition. Co-occurrence of Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp. and Euestheria brodieana (Jones) without Bulbilimnadia sheni Kozur & Weems or with rare B. sheni in the upper part of the zone. Lower boundary. FAD of Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp. Upper boundary. Strong decrease in the abundance of Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp. and a strong correlative increase in the abundance of Bulbilimnadia sheni Kozur & Weems, such that the number of B. sheni greatly exceed the number of B. killianorum Kozur, Weems & Lucas n. sp. Important conchostracan species. Euestheria brodieana Jones, dominant; Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp., rare in the lowermost part of the zone, above it common; B. sheni Kozur & Weems, very rare and only in the upper part of the zone. Type locality. Potter Canyon, northern Arizona, about 1 km west of the type section of the Whitmore Point Member. Occurrence. Lower third of Midland Formation in the Culpeper Basin. Middle and upper Whitmore Point Member of the Moenave Formation in southern Utah and NE Arizona. Early Hettangian of Sołtyko´w, Holy Cross Mountains, southeastern Poland (Pien´kowski 2004). Age. Earliest Hettangian. Remarks. The changeover from Rhaetian to Hettangian conchostracan faunas is very gradual. The monospecific E. brodieana fauna of the uppermost Rhaetian is followed by a basal Hettangian fauna that is still strongly dominated by E. brodieana, but additionally includes some B. killianorum. Somewhat later, E. brodieana still dominates, but by this point B. killianorum becomes common and in the upper B. killianorum Zone very rare B. sheni also occur. The successive, still lower Hettangian fauna also is dominated by E. brodieana, but B. sheni has become more common than B. killianorum among Bubilimnadia. This level is assigned to the B. sheni Zone. B. killianorum disappears in the upper

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part of the B. sheni Zone, but E. brodieana is still dominant. Only in the late early Hettangian B. froelichi Zone does Bulbilimnadia become dominant and E. brodieana disappear. Bulbilimnadia sheni Zone Definition. Co-occurrence of Bulbilimnadia sheni Kozur & Weems and other Liassic conchostracans along with dominant Euestheria brodieana (Jones). B. sheni is more common than B. killianorum within the genus Bulbilimnadia, and by the upper B. sheni Zone is the only Bulbilimnadia species present. Lower boundary. The base of the B. sheni Zone is placed at the level where B. sheni becomes more common than B. killianorum. Upper boundary. LOD of Euestheria brodieana (Jones), FAD of B. froelichi Kozur & Weems. Important conchostracan species. Euestheria brodieana (Jones), dominant; Bulbilimnadia sheni, common; Palaeolimnadia. cf. P. longmenshanensis Shen, very rare; P. cf. P. semicircularis Shen, very rare and only in the lower part of the zone. Bulbilimnadia killianorum n. sp. is still present but always distinctly rarer than B. sheni. Occurrence. Middle Midland Formation to lower Waterfall Formation in the Culpeper Basin (Virginia), lower Hettangian; middle East Berlin Formation in the Hartford Basin (Connecticut and Massachusetts), lower Hettangian. Age. Early, but not earliest Hettangian. Bulbilimnadia froelichi Zone Definition. Range Zone of B. froelichi Kozur & Weems. Lower boundary. LOD of Euestheria brodieana (Jones), FAD of B. froelichi Kozur & Weems. Upper boundary. LOD of B. froelichi. Important conchostracan species. Bulbilimnadia froelichi Kozur & Weems; Palaeolimnadia baitianbaesis Chen. Occurrence. Upper Waterfall Formation of the Culpeper Basin. Age. Late early Hettangian. Remarks. The fauna of this zone consists almost exclusively of B. froelichi, but this may be due to this faunal horizon only being sampled at a single locality. Except for B. froelichi, only a very few Palaeolimnadia cf. P. baitianbaensis Chen are known from this locality. Rhaetian holdovers

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(e.g. E. brodieana) are absent. At least one zone above this one is represented by material collected by Nicholas McDonald from the lower and middle Portland Formation in Connecticut. This material will have to be named and described before the next successive zone can be properly defined.

Summary Because of their lifestyle, their reproductive strategy, and their rapid rate of evolution, conchostracans are exceptionally well suited for establishing a widespread and detailed biozonation for continental rocks of the Triassic System. Many Triassic conchostracan species are distributed throughout the entire northern hemisphere, and in the Late Permian to Middle Triassic interval some of these forms also are found in Gondwana as well. Upper Triassic Gondwanan conchostracan faunas are different from conchostracan faunas of the northern hemisphere. In the late Carnian and Norian, some provincialism has been documented even within the northern hemisphere. For example, the late Carnian Howellisaura and the Sevatian Redondestheria seem to be restricted to North America, while Acadiestheriella n. gen. so far has been found only in the Sevatian deposits from the Fundy Basin of southeastern Canada. The tolerance by conchostracans of a variety of aqueous environments, ranging from temporary fresh water pools, permanent lakes, and alkaline lakes to brackish marine environments, makes them some of the most commonly found invertebrates in continental environments and also makes it possible to find them in marginal marine or intertonguing marine and marginal marine environments where they can be closely intercorrelated with the standard marine biozones of the Triassic. The stratigraphic resolution that can be achieved with conchostracan zones is often as high as for ammonoid and conodont zones found in pelagic marine deposits. In this paper, we establish a conchostracan zonation for the Changhsingian (Late Permian) to Hettangian (Early Jurassic) of the northern hemisphere. Within the Triassic alone, we recognize 34 conchostracan biozones that permit detailed correlation of strata within this period. For the most part, this zonation is very well correlated with the marine scale and is especially robust for the Changhsingian to early Anisian interval, for the late Ladinian to Cordevolian interval, and for the Rhaetian to Hettangian interval. Conchostracan faunas within the Changhsingian to early Anisian interval unequivocally establish that the Permian– Triassic boundary lies within the flood basalts of the Siberian Trap. Thus, the acme of this flood basalt event precisely matches the Permian–

Triassic boundary. For much of the Middle and Upper Triassic, our zonation is preliminary, in great measure because many of the taxa within this interval are undescribed. Also, because of some late Carnian endemism in North America, several zones are established for local usage in the Newark Supergroup of the eastern United States and in the Chinle Group of the American southwest. In the Late Triassic, there are two major extinctions among conchostracans. The first is in the early Norian (Lacian), when fairly diverse forms descended directly from Carnian forebearers mostly go extinct. This is followed by a brief interval within the Lacian of low conchostracan diversity, followed by a radiation of new forms including the characteristic Norian genera Redondestheria and Shipingia. Through the remainder of the Norian these forms tend to become larger overall. The second major extinction was at the end of the Norian, when both Redondestheria and Shipingia went extinct. Throughout the Rhaetian and the early Hettangian (earliest Jurassic), conchostracan faunas consist universally of small forms and generally are represented by low diversity faunas. It was not until after the major CAMP volcanic episode, in the region that was to become the Atlantic Ocean, that conchostracan faunas became larger and more diverse again. This prolonged Rhaeto-Hettangian interval of small and low diversity faunas probably reflect global environmental stress caused by the CAMP volcanic event. Some conchostracan taxa, of exceptional stratigraphic importance for the Triassic but not previously recognized, are described here in an accompanying appendix. These represent five new genera, six new species and a new subspecies. We wish to thank a large number of individuals and institutions that have been crucial to making this study truly comprehensive and thorough. Foremost, we would like to thank Prof. Shen Yanbin (Nanjing), who improved our paper greatly through very productive discussions, providing data about the taxonomy and stratigraphic ranges of Triassic conchostracans in China, by translating some of the relevant Chinese literature (especially regarding diagnoses and ranges), and by providing copies of relevant conchostracan literature that we were unable to find elsewhere. Prof. Gerhard H. Bachmann (Halle) helped greatly throughout this study by providing photography and other technical assistance, stratigraphic data, and by organising several excursions within the Germanı`c Basin to sample for conchostracans and improve the stratigraphic subdivision and correlation of the Germanic Triassic. Prof. Paul Olsen (Lamont-Doherty, Palisades, New Jersey) and Dr Spencer Lucas (New Mexico Museum of Natural History and Science, Albuquerque) kindly loaned us specimens, provided outcrop data, and helped us through discussions concerning the Triassic

TRIASSIC CONCHOSTRACANS stratigraphy and literature of the United States. Grzegorz Niedz´iedzki, PhD, Dipl. Geol. Tadeusz Ptaszyn´ski, Warsaw), Dr Tomasz Sulej (Warsaw), and Prof. Grzegorz Pien´kowski (Warsaw) provided specimens from the Late Triassic and Early Jurassic of Poland and also provided very useful discussions concerning the stratigraphy of that area. Prof. Mark Hounslow (Lancaster) provided specimens from England, helped us locate relevant literature concerning their stratigraphic setting, and provided enlightening discussions and excursion opportunities. Dr Jochen Lepper (Hannover) took the senior author on several very productive sampling excursions in the Solling Mountains area and provided very useful discussions. Prof. Nicholas McDonald (Simsbury, Connecticut) made his comprehensive collection of specimens from the Hartford Basin available to us, and Mr Richard Fulton (Gettysburg, Pennsylvania) made available his collection of stratigraphically important specimens from the Gettysburg Basin. We also thank Dipl. Geol. Andreas Etzold (Freiburg) for making important Germanic Basin material available to us and for providing significant outcrop data and discussions concerning the stratigraphy of the Keuper in southern Germany and the sampling of the Malschenberg-1 borehole. Drs Jens Barnasch (Kassel) and Matthias Franz (Freiberg) provided us with material, outcrop data, and discussions concerning the Keuper stratigraphy of the Germanic Basin. Klaus-Peter Kelber (Wu¨rzburg) made available to us exceptionally wellpreserved material from the Keuper of southern Germany. Dr Thomas Martens (Gotha) provided specimens and valuable data concerning outcrops in the Germanic Basin. Drs Edgar Nitsch (Freiburg) and Dieter Seegis (Schorndorf) provided very useful discussions concerning the stratigraphy of the Germanic Keuper and also important outcrop data. Dr Manfred Warth (Stuttgart) provided important outcrop data and material from the Germanic Basin. Drs Joachim Iffland (Gu¨strow) and Ju¨rgen Kopp (Kleinmachnow) provided significant help in locating conchostracan samples from boreholes in northern Germany. Dr Marc Durand (Nancy) provided valuable material from the Keuper of France and important data concerning outcrops and literature. Prof. G. N. Sadovnikov (Moscow) helped us to find relevant literature and data on the distribution of conchostracans in the Tunguska Basin and northeastern Siberia. Prof. Sankar Chatterjee (Texas Tech University, Lubbock) kindly allowed us to borrow stratigraphically important conchostracan specimens from the Museum of Texas Tech and Bill Mueller (Texas Tech University, Lubbock) helped greatly by taking us out to collect new material and by providing valuable discussions concerning the stratigraphy of the Texas Triassic. Dr Wolfram M. Ku¨rschner, Utrecht, provided us with material from the lower Whitmore Point Memmber of Potter Canyon. Dr Andrew R. C. Milner (St. George, Utah) provided us with stratigraphically important specimens from the basal Jurassic of the Colorado Plateau region, Prof. Shekhar Chakdra Ghosh (Calcutta) provided us with literature and discussions concerning the Lower Triassic conchostracans of India, and Dr Oscar F. Gallego (Corrientes) provided us with literature and discussions concerning South American occurrences of Triassic conchostracans. We also wish to thank Dr Gu¨nter Schweigert of the Naturkunde Museum (Stuttgart), Justin Spielmann of the New

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Mexico Museum of Natural History and Science (Albuquerque), Janet and David Gillette of the Museum of Northern Arizona (Flagstaff), Cope MacClintock of the Yale Peabody Museum (New Haven), Matthew Carrano, Michael Brett-Surman, Dan Levin, and Charyl Ito of the United States National Museum (Washington, D.C), Vincent Schneider and Patricia Weaver of the North Carolina Museum of Natural Sciences (Raleigh), and Deborah Skilliter of the Nova Scotia Museum (Halifax) for making material from their museums available to us for study and providing relevant museum numbers and locality information. Finally, we wish to thank Joseph Smoot (USGS), John Repetski (USGS), Justin Spielmann (NMMNH&S), Spencer Lucas (NMMNH&S), and James Ogg (Purdue University) for very helpful reviews of the manuscript of this paper, and Lynn Wingard (USGS) for kindly making her photographic equipment available to us for high-resolution photographs of many of the specimens shown herein.

Appendix 1 – Description of some new, stratigraphically important taxa Family Euestheriidae Defretin-Lefranc, 1965 Genus Euestheria Depe´ret & Mazeran, 1912 Type species. Posidonia minuta von Zieten, 1833 Euestheria albertii mahlerselli Kozur & Lepper n. subsp. 1993 Euestheria n. sp B – Kozur, Mahler & Sell, p. 258, figs 3–4 1993 Euestheria albertii n. subsp. A – Kozur, Mahler & Sell, p. 258, figs 3– 6

Derivatio nominis. In honour of Horst Mahler, Veitsho¨chheim, Franconia (Germany) and Ju¨rgen Sell, Euerdorf, Franconia (Germany), for their excellent work on the fauna of the Solling and Ro¨t formations in Franconia, Germany. Holotype. The specimen illustrated by Kozur et al. (1993) in figs 3 –6, rep.-no. SMTE 5826/13-18, from the collection of Horst Mahler/Ju¨rgen Sell. Type locality. Aura, southwest of Bad Kissingen (Franconia, Germany), TK 25: 5826 Bad Kissingen Su¨d, R: 35 72 750, H: 55 60 475. Stratum typicum. Greenish-grey claystones in the uppermost Thuringian Chirotherium Sandstone. Diagnosis. Very small, strongly convex Euestheria with commonly stout but also more rarely slender morphotypes. The strongly convex umbo is situated at the end of the anterior third of the shell and somewhat overreaches the dorsal margin. In front of the umbo, the dorsal margin is short, directed obliquely downward and only

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slightly separated from the anterior margin. Behind the umbo, the dorsal margin is straight, directed slightly downward and separated from the posterior margin by a posterodorsal corner. Anterior margin higher and somewhat less convex than the posterior margin. Strongest curvature of the anterior and posterior margins somewhat above the mid-line, at the posterior margin sometimes partly along the mid-line. Ventral margin convex, with strongest curvature somewhat anterior to the mid-length. The 9– 20 growth bands are narrow and rather uniform; only in the umbonal field are they somewhat narrower and not so pronounced. Free umbonal field very small. Very wellpreserved specimens show an indistinct, very small reticulation.

Measurements. Stout morphotype: l ¼ 2.4–3.3 mm, h ¼ 1.6–1.9 mm, l/h ¼ 1.11 –1.44 Slender morphotype. l ¼ 2.35 –3.2 mm, h ¼ 1.5– 1.8 mm, l/h ¼ 1.57–1.74

Age. Late Spathian and Aegean. Euestheria albertii mahlerselli– Palaeolimnadia alsatica alsatica Zone. Occurrence. Euestheria albertii mahlerselli occurs rarely in the lower Solling Formation, where the Spathian guide forms Euestheria exsecta (Novozhilov), Palaeolimnadia nodosa (Novozhilov) and P. alsatica detfurthensis Kozur & Seidel are still dominant. In younger parts of the Solling Formation, the frequency of E. albertii mahlerselli increases upward as the frequency of the other three mentioned species decreases. In the basal Anisian upper Solling Formation Stammen Beds of the Solling Mountains and in the contemporaneous Thuringian Chirotherium Sandstone of Thuringia and Franconia, E. albertii mahlerselli is either the only known species (Franconia, Solling Mountains) or it co-occurs with P. alsatica alsatica Reible (Thuringia). The Spathian guide forms E. exsecta, P. nodosa and P. alsatica detfurthensis are no longer present in that level. In the lower and middle Ro¨t Formation, below the basal Bithynian ‘Dolomitische Grenzbank’ of the middle Ro¨t Formation with Costatoria costata (Zenker) and Myophoria vulgaris (von Schlotheim), the same conchostracan fauna is present as in the upper Solling Formation. This level can be correlated with the Aegean by marine faunas (Kozur 1999). The upper Solling Formation Stammen Beds/Thuringian Chirotherium Sandstone interval also belongs to the Aegean as indicated by the first occurrence of the sporomorph Hexasaccites thiergartii (Ma¨dler) Kozur (Brugman 1986; Bachmann & Kozur 2004). Thus, the total range of E. albertii mahlerselli is late Spathian to Aegean. Remarks. The only significant difference between E. albertii mahlerselli and E. albertii albertii (Voltz) is the consistently smaller size of E. albertii mahlerselli. Additionally, typical specimens of E. albertii albertii are proportionately higher (i.e. their l/h index is smaller).

Typical E. albertii albertii from the Gre`s a` Voltzia in Lorraine are 5 mm long. The most primitive representatives of E. albertii albertii come from the basal Bithynian middle Ro¨t Formation (‘Dolomitische Grenzbank’). They reach a maximum length of 4 mm and are found together with Costatoria costata and Myophoria vulgaris (Kozur et al. 1993). Advanced E. albertii albertii, from the upper Bithynian Holbrook Member of the Moenkopi Formation of Arizona, are up to 6– 6.3 mm long.

Genus Gregoriusella n. gen. Derivatio nominis. In honour of Grzegrorz Niedz´wiedzki PhD, Warsaw, for his excellent palaeontological work on the continental Triassic of Poland. Grzegorz (Latinized) ¼ Gregorius. Type species. Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. Diagnosis. Strongly convex, generally oval, very small to small carapace. The strongly convex umbo is situated in the anterior third of the carapace and overreaches the short dorsal margin. The 12–25 narrow growth bands have a uniform width. The microsculpture on the outer growth bands consists of often indistinct, short radial lirae; on the inner growth bands there are short radial lirae, fine reticulations or a pitted surface. If the outer layer of the shell is present, then between the radial lirae or within the reticulations a pitted surface can be recognized. Assigned species. Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp.; Palaeestheria fimbriata Warth, 1969; Menucoestheria bocki Olempska, 2004; Gregoriusella spp. (several yet to be described new species, mainly from the Tuvalian). Age. Upper Cordevolian to lower Rhaetian. Remarks. Two species, here assigned to Gregoriusella, in the past were classified as Palaeestheria (P. fimbriata Warth 1969) and Menucoestheria (M. bocki Olempska 2004). Palaeesteria Barnard (1929) was not adequately defined and was used by him in a broad sense for all fossil conchostracans without diagnosis. He listed three species under this genus, and from these Raymond (1946) selected Estheria anomala Jones as the type species. This Lower Cretaceous species is not well known, but it does have a distinct free umbonal area and therefore is probably a limnadiid conchostracan. Microsculpture is not reported from this species. This genus therefore is quite different from Gregoriusella and can be distinguished by its larger size, larger free umbonal area and absence of microsculpture, though the last characteristic may be an artefact because the type species is not well known and no modern description has been made. The type species of Menucoestheria Gallego & Covacevich, 1998 (M. ternaraensis Gallego 1998) is a large

TRIASSIC CONCHOSTRACANS conchostracan with reticulation on the inner growth bands and radial lirae on the outer growth bands. It may be a junior synonym of Anyuanestheria Zhang & Chen 1976 but this is uncertain because Menucoestheria is a Gondwanan genus and it cannot be proven as yet that it evolved from Laxitextella. Gregoriusella has a rather similar microsculpture, but the space between the radial lirae is pitted. During the upper Cordevolian in the Germanic Basin, Gregoriusella evolved from small Euestheria, while Anyuanestheria was evolving from large Laxitextella of the L. laxitexta group. In this case, the difference in size between Anyuanestheria and Gregoriusella can be used to differentiate the two genera. The same difference exists between Gregoriusella and Menucoestheria. Euestheria Depe´ret & Mazeran, 1912 is very similar, but it has either no microsculpture or a microsulpture that consists of mostly indistinct fine reticulations. Gregoriusella gradually evolved during the late Cordevolian from Euestheria by development of radial lirae, at first only on the outer growth bands though in later forms the entire shell may have radial lirae. The oldest species is Gegoriusella fimbriata (Warth). This species evolved during the Cordevolian from Euestheria by development of rather indistinct radial lirae. The microsculpture of this species is rather misleadingly documented in Warth (1969, pl. 3). On this plate the holotype is shown together with two SEM photographs of microsculpture. These SEM pictures are neither from the holotype nor from Gegoriusella fimbriata. They are of shell fragments of transitional forms between Laxitextella Kozur and Anyuanestheria Zhang & Chen which are much larger and have much more distinct microsculpture.

Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp. (Fig. A1, Photographs 1– 10)

Derivatio nominis. In reference to its occurrence in Poland.

Holotype. The specimen in Figure A1: 1, 2; ZPAL V33/543 in the collection of the Institute of Paleobiology, Polish Academy of Science, Warsaw.

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Anterior and posterior margins symmetrically rounded, anterior margin higher than posterior margin and curvature less than in posterior margin. Posterodorsal margin somewhat bevelled. Ventral margin symmetrically rounded. The 12–25 growth bands are narrow and rather uniform. The microsculpture on the outer growth bands consists of short radial lirae, but the middle and inner growth bands are finely reticulated or pitted. If the outer layer of the shell is preserved, the space between the lirae or within the reticulae are densely pitted (Fig. A1: 10).

Measurements. l ¼ 2.2–3.3 mm, h ¼ 1.3–2.3 mm, l/h ¼ 1.25–1.44

Age. Uppermost Sevatian and lower Rhaetian. Occurrence. In grey claystones, siltstones, and rarely sandstones, all rich in plant detritus, Lipie, Upper Silesia, uppermost Norian to lowermost Rhaetian. In the basal Exter Formation (also rich in plant detritus) in the Tarnow borehole from northern Germany, uppermost Norian to lowermost Rhaetian. Upper Redonda Formation (Duke Ranch Member) in Apache Canyon, New Mexico, lowermost Rhaetian. Remarks. At Lipie, Rhaetipollis germanicus first appears in the middle of the exposed four metres of sediments, where G. polonica is present. In general, the FAD of R. germanicus is regarded as a marker for the base of the Rhaetian, but it is very rare or perhaps even absent in the basal-most part of the Rhaetian. As the very large (often 10–12.5 mm long) Shipingia olseni and Redondestheria novomexicoensis that characteristize the upper Sevatian are no longer present in the G. polonica Zone, the entire G. polonica Zone is assigned to the lowermost Rhaetian, but it remains possible that this zone could begin slightly before the Norian– Rhaetian boundary. Gregoriusella bocki (Olempska) is larger than G. polonica (up to 4.8 mm long) and also is the largest known species of Gregoriusella. The ventral margin is only slightly convex. Gregoriusella fimbriata (Warth) is perfectly oval. The anterior and posterior margins have almost the same height and the posterodorsal margin is not distinctly bevelled.

Locus typicus. Clay quarry at Lipie, Upper Silesia, Poland.

Stratum typicum. Within a 4 m succession of plant-bearing grey mudstones and siltstones with some sandstones. Upper conchostracan horizon with first Rhaetipollis germanicus Schulz. Lowermost Rhaetian.

Material. More than 100 specimens. Diagnosis. Small, oval, strongly convex carapace. Umbo situated in the anterior third of the carapace, strongly convex, overreaches the dorsal margin. Dorsal margin short, straight to very slightly convex, with gradual transition into the anterior and posterior margins.

Family Xiangxiellidae Shen 1976 The Xiangxiellidae were established by Shen Yan-Bin (in Chang et al. 1976) as a subfamily of the Vertexiidae Kobayashi 1954. The Xiangxiellidae, with one or two rather broad radial elements on the free umbonal area, are only distantly related to the Vertexiidae, which have a hollow spine on the free umbonal area and often also possess long spines on the dorsal margin and sometimes a long posterodorsal spine. The umbonal spine is often deformed and may be squeezed onto the umbonal area to give the appearance of a conical umbonal sculpture or a radial element. Where undeformed, however, it stands

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Fig. A1. 1– 10: Gregoriusella polonica Kozur, Niedz´wiedzki & Sulej n. sp., lower Rhaetian G. polonica Zone; 1, 2: Holotype, outer layer of the shell corroded and radial lirae therefore not preserved in their full length, SEM picture. Lipie (Upper Silesia), upper conchostracan horizon, material from G. Niedz´wiedzki, Warsaw. 1: total view; 2: detail of shell surface with partly preserved radial lirae. 3, 4: Specimen with deformed anterior margin, shell partly preserved. From borehole Tarnow 1 –65, sample KO 7-02-10 at 956.7– 956.8 m, 1.0–1.1 m above the base of the Exter Formation; 3: total view; 4: detail. 5: Microsculpture of the outer layer of the shell with very indistinct radial lirae and reticulation, both distinctly pitted, SEM picture, from borehole Tarnow 1 –65, sample KO 7-02-12, lower Exter Formation at 954.15 m, 3.55 m above the base of the Exter Formation, 6 –8: Lipie (Upper Silesia), specimen with preserved different layers of the shell, SEM pictures, material from G. Niedz´wiedzki, Warsaw; 6: Total view; 7: Preserved outer layer of shell in the middle part of the carapace, with

TRIASSIC CONCHOSTRACANS perpendicular or oblique to the shell surface. Most probably, the Xiangxiellidae evolved from such representatives of the Palaeolimnadiidae Tasch, in which a node is present on the free umbonal area. If this node became elongated, it easily could have changed into an umbonal radial element. Palaeolimnadia with an umbonal node, such as P. alsatica detfurthensis Kozur & Seidel, are present in the Spathian. In some specimens, the umbonal node already has become oval and these or similar forms could be the forerunners of the Xiangxiellidae.

Genus Hornestheria Kozur & Lepper n. gen. Derivatio nominis. In honour of the late Dr Manfred Horn (1934– 1999, Wiesbaden). Type species. Hornestheria sollingensis Kozur & Lepper n. gen. n. sp. Diagnosis. Shell small to moderately large, with rather low convexity. Dorsal margin moderately long, straight, with a distinct posterodorsal corner but often only an indistinct anterodorsal corner. The anterior margin is slightly rounded, with its greatest curvature generally distinctly above the mid-line. In most species, the lower part of the anterior margin is strongly oblique. The posterior margin is strongly rounded and often in the upper part oblique. Its greatest curvature is somewhat below the mid-line. The ventral margin is moderately to strongly convex. The free umbonal area is distinct, but not very large. It bears a swollen radial element that is often quite distinct, but for preservational reasons it is sometimes not recognisable. There are 7–17 growth bands; the inner bands have a dense, fine reticulation, but in adult forms the outer growth bands bear only fine, short radial lirae. Age. Uppermost Spathian and Anisian. Occurrence. Karlshafen Beds of the Solling Formation (uppermost Spathian) of the Solling Mountains. Anisian of China.

Assigned species. Hornestheria sollingensis Kozur & Lepper n. gen. n. sp.; Protomonocarina sinensis Shen 1976; Protomonocarina carinata Shen 1976; Protomonocarina oblonga Shen 1976; Protomonocarina ziguiensis Shen 1976; Protomonocarina hubeiensis Shen 1976; Protomonocarina buzhuangheensis Shen 1976; Protomonocarina xiangxiensis Shen 1976.

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Remarks. Chang et al. (1976) assigned the Anisian species that we now assign to Hornestheria to the Lower Permian genus Protomonocarina Tasch 1962. Only one specimen of one species shows what might be interpreted as a characteristic feature of this latter genus, namely a radial element on the free umbonal area that is distinctly segmented. This radial element, however, clearly is not a sculptural element, but rather a soft part element (appendage) that is visible through the chitinous shell. Even if the radial ridge on the free umbonal area is a sculptural radial element, Protomonocarina Tasch, 1962 still would be a junior synonym of Lioestheria Depe´ret & Mazeran 1912. The type species of this genus, Estheria (Lioestheria) lallyensis, also has a radial element on the free umbonal area and no reticulation on the growth bands, just as in Protomonocarina. Some species of Curvacornutus, which Novozhilov (1970) described from the Tatarian of Russia, are similar, but they have a distinctly curved radial element, and the shell is not reticulated. The type species of Curvacornutus, the Lower Permian C. primus Tasch from North America, is based on a single specimen that probably is not closely related to the species that Novozhilov (1970) described under this genus. The Permian Megasitum Novozhilov 1970 is distinctly different in that it has a larger free umbonal area and a much wider and larger radial element.

Hornestheria sollingensis Kozur & Lepper n. gen. n. sp. (Fig. A2, Photographs 1 –7)

Derivatio nominis. In reference to its occurrence in the Solling Formation of the Solling Mountains. Holotype. Figure A2: 1, 2, rep. no. Kozur-Lepper 2003.53 MLU (in the collection of the Institut fu¨r Geologische Wissenschaften of the Martin-Luther-Universita¨t Halle). Type locality. Quarry Frohrieper Berg (Fa. Bunk), TK 25 Nr. 4322; R: 35 31 300, H: 57 24 570. Stratum typicum. Karlshafen Beds of the Solling Formation, in a greenish-grey claystone with a rich conchostracan fauna, about 4 m above its basal boundary with the Trendelburg/Karlshafen Beds. Uppermost Spathian.

Fig. A1. (Continued) fine reticulation; 8: preserved outer layer of the shell on the outer growth bands with radial lirae, space between the radial lirae pitted. 9: Carapace surface with partly preserved shell, including outer layer of the shell, distinct radial lirae, from borehole Tarnow 1– 65, sample KO 7-07-9 at 956.95 m, 0.85 m above the base of the Exter Formation; 10: Shell fragment from the surface of a complete carapace, outer layer well-preserved, SEM picture, radial lirae with pitted surface between the radial lirae on the outer growth bands, in the upper part of the picture (growth bands from the middle part of the shell) pitted surface without radial lirae. From borehole Tarnow 1– 65, sample KO 7-02-7 at 957.8 m, base of Exter Formation. The material illustrated in 3– 5, 9, and 10 is deposited in the collection of the Institut fu¨r Geologische Wissenschaften, Martin-Luther-Universita¨t Halle.

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Fig. A2. 1– 7: Hornestheria sollingensis Kozur & Lepper n. gen. n. sp., from quarry Frohrieper Berg (Fa. Bunk), TK 25 Nr. 4322; R: 35 31 300, H: 57 24 570. Karlshafen Beds (Solling Formation), greenish-grey claystone with rich conchostracan fauna, about 4 m above the basal boundary on the Trendelburg/Karlshafen Beds. Uppermost Spathian.

TRIASSIC CONCHOSTRACANS

Material. More than 100 specimens in different preservation.

Diagnosis. With the characteristics of the genus. The small, only slightly convex shell has a straight dorsal margin with an indistinct anterodorsal corner and a distinct posterodorsal corner. The anterior margin is slightly to moderately rounded, with the greatest curvature somewhat above the mid-line. The lower part of the anterior margin is only slightly or occasionally moderately bevelled. The posterior margin is more strongly rounded than the anterior margin, with its greatest curvature situated in the middle. It is slightly (or, rarely, moderately) bevelled in its upper part. The ventral margin is moderately convex. The length of the free umbonal area is 25–30% of the total length of the shell. It has a distinctly swollen but rather narrow radial element that is about half as long as the height of the free umbonal area. The 13–17 growth bands and the free umbonal area all have a fine and dense reticulation. On the outer growth bands, the reticulation is replaced by delicate, short, radial lirae. Measurements. l ¼ 3.0–3.8 mm, h ¼ 1.95–2.85 mm, l/h ¼ 1.3– 1.66

Age. Uppermost Spathian. Occurrence. Karlshafen Beds of the Solling Formation in the Solling Mountains, Germany. Remarks. The Anisian Hornestheria species of China are smaller (2.0–2.8 mm long), and only in H. ziguiensis (Shen, 1976) are short indistinct radial lirae present on the outermost growth bands of adult forms. This latter species is most similar to H. sollingensis in its outline, but it is consistently smaller (2.3– 2.8 mm long), and the reticulation and radial lirae are weaker than in H. sollingensis. The other Anisian Hornestheria species have a different outline in addition to their smaller size and weaker sculpture. In some specimens the lower part of the anterior margin and the upper part of the posterior margin are strongly bevelled, as in H. sinensis (Shen), or else they have a subrectangular outline, as in H. carinata (Shen).

Family Estheriellidae Kobayashi 1954 emend. Kozur & Seidel 1983 Genus Acadiestheriella n. gen. Derivatio nominis. The genus is named in honour of Acadia University in Wolfville, Nova Scotia. Type species. Acadiestheriella cameroni n gen. n. sp.

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Diagnosis. Carapace small to moderately large, moderately convex. Dorsal margin long, straight, without distinct corners at the anterior and posterior margins. Free umbonal area distinct but small, situated terminally (anterodorsally). The 8 –12 growth bands have a dense, fine reticulation and bear small to large, roundish nodes arranged in four to seven (rarely eight) radial lines. Assigned species. Acadiestheriella cameroni n. gen. n. sp.

Age. Sevatian. Occurrence. So far only known from the Fundy Basin (southeastern Canada). Remarks. Most similar is Lioleaiina Novozhilov 1952 emend. Kozur & Seidel 1983. This genus has two to five radial ridges on the lateral shell surface that bear nodes at crossing points with the growth lines. Exceptionally, the radial ridges are broken up into radial lines of nodes. However, these nodes are situated on elevated radial line segments, so even in these forms low radial ridges are present. Characteristic of Lioleaiina is a dorsal ridge with nodes parallel and very close to the dorsal margin. This dorsal ridge or line of nodes at the dorsal margin and parallel to it is not present in Acadiestheriella. The youngest Lioleaiina are found in the Anisian and Ladinian of the Germanic Basin and consistently have three nodose radial ribs on the lateral shell and one along the dorsal margin. Earlier, lower Olenekian Lioleaiina have up to five nodose lateral ribs and one dorsal rib. Thus, the trend in this genus is toward reduction of the nodose radial ribs. As Acadiestheriella has up to eight radial lines of nodes, it seems unlikely that it was derived directly from Liolaiina. More probably, Acadiestheriella evolved directly from Estheriella Weiss 1875. This would explain why four to seven, and sometimes even eight, radial lines of nodes are present in Acadiestheriella. It is not likely that the trend toward reduction of the number of radial nodose ribs within the Lioleiina from the Smithian to the Longobardian would be reversed above the Ladinian. Estheriella Weiss 1875 is most readily distinguished from Acadiestheriella by the presence of nodose ridges in Estheriella instead of the radial lines of isolated nodes that are characteristic of Acadiestheriella. As in Estheriella, the nodose sculpture is very variable in Acadiestheriella. In Lioleaiina, the nodose sculpture is more uniform.

Fig. A2. (Continued) 1, 2: Holotype; 1: total view; 2: Detail of the anterior part, showing microsculpture of very fine reticulation and indistinct radial lirae; 3: Radial element readily visible; 4 –7: Specimens with preserved distinct, narrow radial elements. Scale: 1 mm. All material is deposited in the collection of the Institut fu¨r Geologische Wissenschaften, Martin-Luther-Universita¨t Halle.

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Fig. A3. 1– 8: Acadiestheriella cameroni n. gen. n. sp., Blomidon Provincial Park, Fundy Basin, Nova Scotia, southeastern Canada, upper Blomidon Formation, Sevatian. 1: Dorsal view of a double-valved specimen, anterior margin to the left. Nodes arranged in seven distinct radial lines; 2: Specimen with a few small nodes arranged in six indistinct radial lines; 3: Specimen with a few nodes, arranged in four indistinct short radial lines; 4, 5: Holotype, nodes

TRIASSIC CONCHOSTRACANS

Acadiestheriella cameroni n gen. n. sp. (Fig. A3, Photographs 1– 8)

Derivatio nominis. The species name is in honour of Professor Barry Cameron, recently retired from Acadia University, who first reported (but did not describe) this form from the Fundy Basin. Holotype. NSM007GF023.017, part and counterpart, Figure A3: 4, 5. Type locality. 45.2343 N, 64.3546 W, about 3 km SE along the shoreline from the bottom of steps down to beach at White Water picnic area in Blomidon Provincial Park, Cape Blomidon, Nova Scotia, Canada. Stratum typicum. Near the middle of the Blomidon Formation in the Fundy Basin. Material. More than 20 specimens on 17 small slabs of rock (4 with counterparts), deposited in the collections of the Nova Scotia Museum, Halifax, Nova Scotia (NSM007GF023.001through NSM007GF023.17). Diagnosis. As for the genus. Description. Carapace small to moderately large, moderately convex. Dorsal margin long, straight, without distinct corners at the anterior and posterior margin. Anterior margin slightly rounded, anteroventral margin slightly oblique. Posterior margin stronger and symmetrically rounded, greatest curvature either somewhat higher than in anterior margin or else in anterior and posterior margins at the same height. Ventral margin convex. Free umbonal area distinct but small, located anterodorsally. The 8 –12 growth bands are densely and finely reticulated and bear small to large roundish nodes that are arranged in four to seven (rarely eight) radial lines. The nodes are roundish, very variable, distinct to indistinct, and small to large. They are widely separated and situated at the boundary between adjacent growth bands. The six to eight radial lines of nodes sometimes reach from the outermost growth band to the innermost growth band adjacent to the free umbonal area, but often they are considerably shorter and restricted to only the middle growth bands. In this case, only two or three nodes may be present in a single radial line. Occasionally, the nodes may be totally absent.

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Remarks. A. cameroni may be an endemic form because it is only known from the Sevatian middle Blomidon Formation of the Fundy Basin, where it is locally common.

Family inc. Genus Wannerestheria n. gen. Derivatio nominis. In honour of H. E. Wanner, who discovered and named this species (Wanner 1926). Type species. Estheria mangaliensis pennsylvanicus Wanner 1926.

Diagnosis. Medium-sized, sub-ovate to sub-oblong carapace, moderately convex. Dorsal margin long and straight but without anterodorsal and posterodorsal corners. Anterior and posterior margin convex, nearly of the same height. Ventral margin convex. Free umbonal area small, situated in the anterodorsal corner. Growth bands widest posteroventrally, otherwise rather uniform and strongly elevated at the contact between two adjacent growth bands. When the shell is preserved, contact between two adjacent growth bands possess small, densely-spaced roundish nodes; these are not visible or only indistinctly visible when the shell is not preserved. Assigned species. Estheria mangaliensis pennsylvanicus Wanner 1926.

Occurrence. Basal Gettysburg Formation of the Gettysburg Basin and upper Lockatong Formation (Skunk Hollow Member to the top of the formation) of the Newark Basin. Upper Tuvalian. Remarks. The densely-spaced small nodes on the growth lines clearly distinguish Wannerestheria n. gen. both from Euestheria Depe´ret & Mazeran 1912 and Howellisaura Bock 1953. Euestheria is further distinguished by a stronger convexity of the shell, especially in the umbonal area, and by a shorter straight part of the dorsal margin.

Wannerestheria pennsylvanica (Wanner 1926) (Fig. A4, Photographs 5, 7, 8) 1926 Estheria mangaliensis pennsylvanicus new subspecies – Wanner, p. 25, text-figure 2. Neotype. USNM 539462, Figure A5: 5,7. Paratype: USNM 539461, Figure A5: 8.

Measurements. l ¼ 3 – 5 mm; h ¼ 2.1 – 3.3 mm; l/h ¼ 1.32 – 1.7

Age and occurrence. As for the genus.

Type locality. 40.0932 N, 76.7264 W, along abandoned road bed on north side of Little Conewago Creek, York Haven quadrangle, York County, Pennsylvania.

Fig. A3. (Continued) arranged in four distinct radial lines; photographs taken under different directions of light; 6: Cluster of three specimens, one specimen with very distinct nodes, arranged in six distinct radial lines, another specimen with indistinct nodes, arranged in five radial lines and a third specimen without nodes; 7: Specimen with large nodes arranged in five distinct radial lines; 8: Specimen with both large and small distinct nodes arranged in seven distinct radial lines. Scale: 1 mm. All material is deposited in the collection of the Nova Scotia Museum, Halifax, Nova Scotia.

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Fig. A4. 1, 6: Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp., Potter Spring Canyon, NMMNH locality number 7735, upper Whitmore Point Member of Moenave Formation, purple mudstone, bed 43 of Lucas, basal Hettangian B. killianorum Zone; 1: Holotype; 6: specimen with damaged posterior part. 2, 3: Euestheria brodieana

TRIASSIC CONCHOSTRACANS

Stratum typicum. The basal lake bed sequence of the Gettysburg Formation.

Material. More than 100 specimens, including perhaps a dozen specimens on two slabs of shale that have the neoholotype and neoparatype specimens of this species. Diagnosis. Same as for genus. Remarks. Wanner (1926) described this conchostracan as a new subspecies, Estheria mangaliensis pennsylvanicus, but he failed to designate a holotype and, so far as we can determine, he did not place any of his described material in a museum. Therefore, Estheria mangaliensis pennsylvanicus, as originally described, must be considered a nomen dubium in the absence of any designated type material. Since Bock (1953a) declared Wanner’s subspecies to be a junior synonym of Howellisaura ? ovata (Lea), it has been largely forgotten. We have relocated Wanner’s type locality, however, and gathered numerous new specimens that are indistinguishable from his published material. Close examination of this new material shows that it is actually a distinct species that does not even belong to the same genus as H. ? ovata. It is smaller (3.7–5.5 mm against 6 –8 mm for H. ? ovata), more slender, and the growth lines are more pronounced (distinctly elevated at the contact between two adjacent growth bands). The growth lines between adjacent growth bands also bear densely-spaced small nodes. These nodes are readily recognisable in specimens with well-preserved shells, but they are only indistinctly recognisable in forms without preserved shells. As we have now placed some of our new material in the Smithsonian collections and designated a neotype and a paraneotype, Wanner’s taxon is now re-established with valid type material.

Family Shipingiidae Kozur & Weems 2005 Genus Shipingia Shen 1976 Type species. Shipingia hebaozhaiensis Shen 1976

Shipingia mcdonaldi n. sp. 2007 Shipingia hebaozhaiensis Shen, pars – Kozur & Weems, p. 146, P. 6, figures 3– 5, ? 6

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Derivatio nominis. Species named in honour of Professor Nicholas McDonald of Westminster School, for his many years of research on the Newark Supergroup basins and his many years of collecting important specimens from them. Holotype. The specimen illustrated by Kozur & Weems, 2007, in pl. 6, figure 5, USNM 538458. Type locality. 39.9393 N, 77.1432 W, on the west side of U.S. Highway 15, south-southeast of Heidlersburg, Pennsylvania. Stratum typicum. Heidlersburg Member of Gettysburg Formation, Gettysburg Basin. Material. More than 100 specimens, including the specimen illustrated by Kozur & Weems 2007, in pl. 6, figure 3 (USNM 538459), the specimen illustrated by Kozur & Weems, 2007, in pl. 6, figure 4 (USNM 538460), the top specimen illustrated by Kozur & Weems, 2007, in pl. 6, figure 6 (USNM 538726), and the bottom specimen illustrated by Kozur & Weems, 2007, pl. 6, figure 6 (USNM 538727). Diagnosis. Medium-sized slender carapace with low convexity. Dorsal margin long, straight, distinctly separated from the anterior and posterior margin. Anterior margin rounded in its upper part, strongly oblique and only slightly rounded in its lower part. The oblique part ranges higher up than the mid-line, so that the strongest curvature is in the upper third. Posterior margin higher than the anterior margin, strongly and symmetrically rounded, with its strongest curvature somewhat above the mid-line and its lower part slightly oblique. Ventral margin convex. The umbo is situated in the anterior third, but still distinctly behind the anterodorsal corner. The 10– 12 growth lines are wide in the posterior region and especially in the posteroventral part. The free umbonal area is distinct but small. Growth bands smooth or with an indistinct fine reticulation. Measurements. l ¼ 3.3– 5.5 mm; h ¼ 1.8– 3.0 mm; l/h ¼ 1.7–2.0

Age. Alaunian. Occurrence. Heidlersburg Member of the Gettysburg Formation of the Gettysburg Basin. Upper part of middle

Fig. A4. (Continued) (Jones), different morphotypes, Potter Spring Canyon, NMMNH locality number 7735, northern Arizona, upper Whitmore Point Member of Moenave Formation, purple mudstone, bed 43 of Lucas, basal Hettangian B. killianorum Zone. 4: Euestheria brodieana (Jones), Potter Spring Canyon, NMMNH locality number 7734, northern Arizona, lower Whitmore Point Member of Moenave Formation, greenish-grey to black shales, bed 22 of Lucas, upper E. brodieana Zone, upper Rhaetian. 5, 7, 8: Wannerestheria pennsylvanica (Wanner), lower Little Conewago Creek, York County, Pennsylvania, from locality 1 (40.09328N, 76.72648W) of Wanner (1926), type locality of W. pennsylvanica. Basal Gettysburg Formation of the Gettysburg Basin; 5: total view of neotype, 7: Detail of the posterior third, small nodes on the growth lines readily visible; 8: Paratype specimen without preserved shell, small nodes only indistinctly visible. Scale ¼ 1 mm. The specimens illustrated in 1–4 and 6 are deposited in the collection of the New Mexico Museum of Natural History & Science in Albuquerque. The specimens illustrated in 5, 7 and 8 are deposited in the collections of the Museum of Natural History of the United States National Museum.

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Fig. A5. 1– 4: Norestheria barnaschi n. gen. n. sp., holotype, from borehole Morsleben 52 A, northern Germany, Arnstadt Formation, sample 132 at 162.9– 163.0 m (119.6 m above the base of the Arnstadt Formation), uppermost Alaunian. 1: Total view; 2: Detail of the anterior part of the shell; 3: Detail of the anteroventral part, showing the abrupt

TRIASSIC CONCHOSTRACANS Arnstadt Formation of Thuringia and northern Germany. Basal part of ‘Stubensandstein 3’ of southern Germany.

Remarks. Kozur & Weems (2007) initially regarded S. mcdonaldi n. sp. as juvenile specimens of Shipingia hebaozhaiensis Shen, but S. mcdonaldi occurs both in the Germanic Basin and in the Newark Supergroup earlier than true S. hebaozhaiensis, which is consistently larger (6– 8 mm), not so slender (l/h ¼ 1.5–1.6) and has an anterior margin that is not so strongly oblique in its lower part. This new species also differs from Shipingia baxinensis Shen in that it lacks a distinct microsculpture of radial lirae.

Genus Norestheria n. gen. Derivatio nominis. In reference to the occurrence of this genus in the Norian.

Type species. Norestheria barnaschi n. gen. n. sp. Diagnosis. Carapace moderately large to large, only slightly convex. Free umbonal area distinct but rather small, situated between the anterodorsal and mid-dorsal field. Dorsal margin straight to slightly convex, rather short, with a gradual transition to the anterior margin and a normally indistinct posterodorsal corner. Anterior margin higher than posterior margin, in its lower part with a long oblique region that is not rounded and in its upper part slightly rounded. Posterior margin strongly rounded in its middle region, but in its upper part obliquely directed and straight or only slightly rounded. The lower part of the posterior margin is long, oblique and either straight or only slightly rounded or else rounded and only slightly bevelled. The ventral margin is straight in its anterior and middle region and slightly convex in its posterior region. There is an abrupt angular change going from the ventral to the anterior margin, but the transition from the ventral to the posterior margin is gradual. There are 8 –13 growth bands with a peculiar shape. Ventrally or anteroventrally they are straight, but then they tilt abruptly to similarly straight but obliquely oriented growth bands in the lower part of the anterior quarter. The microsculpture consists of a dense reticulation, partly accompanied by radial lirae. In the abruptly tilted anterior quarter, the elements of the microsculpture are distinctly larger.

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Assigned species. Norestheria barnaschi n. gen. n. sp.; Norestheria n. sp.

Age. Upper Alaunian. Occurrence. Upper part of middle Arnstadt Formation in northern Germany. Basal Stubensandstein 3 of southern Germany. Remarks. Laxitextella Kozur, 1982 does not have the abrupt change in the direction of the growth bands at the end of the anterior quarter of the shell. Although we have this abrupt tilt of the growth bands at the same place in two different species, we cannot quite exclude that this tilting is at least enhanced by deformation. The tilting is accompanied either by an abrupt change in the size of reticulation or of the radial lirae, so it also may be accompanied by an abrupt change in shell thickness at this point. In that case, it would be prone to deformation localized at this place. In Laxitextella the reticulation and other microsculpture elements do not become abruptly larger in the anterior quarter of the shell. Additionally, the shell is more convex, especially in the umbonal area.

Norestheria barnaschi n. gen. n. sp. (Fig. A5, Photographs 1 –4)

Derivatio nominis. In honour of Dr Jens Barnasch, Kassel, for his excellent stratigraphic work in the Keuper of northern and middle Germany. Holotype. The specimen illustrated in Figure A4: 1– 4. Type locality. Borehole Morsleben 52 A, northern Germany

Stratum typicum. Arnstadt Formation, sample 132 at 162.9–163.0 m (119.6 m above the base of the Arnstadt Formation), uppermost Alaunian. Material. Four specimens. Diagnosis. Moderately large flat carapace. Free umbonal area distinct, but rather small, situated between the anterodorsal and mid-dorsal fields. Dorsal margin relatively short, straight to very slightly convex. Anterior margin in the upper part slightly rounded, in the lower part

Fig. A5. (Continued) change from the straight ventral part to the oblique lower part of the anterior margin, accompanied by a distinct change of the type and size of the microsculpture (see also 2); 4: Microsculpture of the ventral part; 5: Norestheria n. sp., detail of the anteroventral part, showing the abrupt change from the straight ventral part to the oblique lower part of the anterior margin, which is accompanied by a distinct increase of the size of the reticulation. From borehole Malschenberg 1 (southern Germany) at 60.1– 60.2 m, basal Stubensandstein 3, upper Alaunian; 6: Norestheria n. sp., fragmentary specimen with its posterior margin in the upper part and its dorsal margin to the right. Posterodorsal corner readily recognisable. From borehole Malschenberg 1 (southern Germany) at 60.0 m, basal Stubensandstein 3, upper Alaunian; 7: Bulbilimnadia killianorum n. sp., Killian site, Fauquier County, Virginia, Culpeper Basin, middle Midland Formation, lower B. sheni Zone in which B. sheni is more abundant than B. killianorum (lower Hettangian). Scale of Figs 1, 6, 7 ¼ 1 mm; Figs 2–5 ¼ 100 mm. The material illustrated in 1 –6 is deposited in the collection of the Institut fu¨r Geologische Wissenschaften, Martin-Luther-Universita¨t Halle. The specimen illustrated in 7 is deposited in the collections of the Museum of Natural History of the United States National Museum.

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straight and strongly oblique. Transition to the ventral margin abrupt and angular. Posterior margin distinctly lower than anterior margin, in the middle part strongly convex, in the upper part oblique and straight or only slightly convex, in the lower part convex. Anterior and middle part of the ventral margin straight or only very slightly convex. Posterior part of the ventral margin convex. 10– 12 growth bands are present. At the border of the anterior quarter they tilt upward abruptly at a sharp angle from horizontal and parallel in the anterior part of the ventral margin to obliquely upward in the anterior quarter. The microsculpture consists of fine dense reticulation with radial lirae. Where the growth lines turn abruptly upwards, the microsculpture in the anterior quarter changes to distinctly larger elements.

Measurements. l ¼ 4.5– 5.6 mm; h ¼ 3.0– 3.3 mm; l/h ¼ 1.47– 1.55

Age. Upper Alaunian. Occurrence. Upper part of middle Arnstadt Formation in the Germanic Basin. Remarks. There is another undescribed species of Norestheria from the basal Stubensandstein 3 (Fig. A4, photographs 5 and 6). It has a microsculpture that consists only of a dense fine reticulation that abruptly changes into a coarse reticulation in the anterior quarter, at the point where the direction of the growth lines abruptly changes.

Family Bulbilimnadiidae Kozur & Weems 2005 Genus Bulbilimnadia Shen 1976 Type species. Bulbilimnadia bullata Shen 1976

Bulbilimnadia killianorum Kozur, Weems & Lucas n. sp. (Fig. A4, Photographs 1, 6; Fig. A5, Photograph 7)

Derivatio nominis. In honour of Jim and Bobbie Killian, who graciously let us dig on their property to recover numerous well-preserved lower Hettangian conchostracans, among them supplementary specimens of B. killianorum. Holotype. New Mexico Museum of Natural History specimen NMMNH P-58455 (Fig. A4: 1). Paratype. United States National Museum specimen USNM 538728 (Fig. A5: 7). Type locality. NMMNH locality 7735 (36.8801N, 112.8473W), Potter Canyon, Mohave County, northern Arizona.

Stratum typicum. 0.5 m-thick bed 43 of Spencer Lucas, purple mudstone in the Whitmore Point Member

of the Moenave Formation, about 3.5 m below the base of the Springdale Member of the Kayenta Formation.

Material. More than 100 specimens. Eight specimens from the Midland Formation of Virginia are included under the lot number USNM 538729. Specimens from the Whitmore Point Member of Utah and Arizona are on New Mexico Museum of Natural History slabs numbered P51076, P-51080, P-51082, P-51085, P-51086, and P-58440. Diagnosis. Small to very small, strongly convex carapace. Free umbonal area large (but for the genus rather small), with small, roundish node on the end of the anterior third of the free umbonal area, close to the dorsal margin. Umbo located at the end of the anterior third to the midlength of the carapace, strongly convex, somewhat overreaching the straight to slightly convex dorsal margin. Convex anterior and posterior margins develop gradually from the dorsal margin. Anterior margin in the lower part somewhat bevelled, with its strongest curvature somewhat above the mid-line. Posterior margin symmetrically rounded with its strongest curvature in the mid-line. Ventral margin strongly convex. 9– 17 rather narrow and uniform growth lines are present that bear an indistinct fine reticulation. Measurements. l ¼ 1.5–3.1 mm; h ¼ 1.28–2.1 mm; l/h ¼ 1.24–1.75

Age. Basal Hettangian B. killianorum Zone to lower middle part of lower Hettangian B. sheni Zone. Occurrence. Lower and middle Midland Formation of the Culpeper Basin. Upper Whitmore Point Member of the Moenave Formation of northern Arizona and southern Utah. Early Hettangian of Sołtyko´w, Holy Cross Mountains, southeastern Poland. Remarks. The successor species Bulbilimnadia sheni Kozur & Weems, 2005 has an elongated node on the free umbonal area. The small round node on the free umbonal area of this new species is fairly robust but still small for a member of the genus Bulbilimnadia. Although B. killinorum is not a typical Bulbilimnadia, it is connected by intermediate transitional forms to Bulbilimnadia sheni Kozur & Weems and thus is the direct forerunner of the first typical species of a lineage restricted to the Hettangian. For this reason, it would be pointless to establish a new genus for this single species, which is probably transitional between Euestheria brodieana (Jones) and Bulbilimnadia sheni Kozur & Weems.

Note added in proof After completion of this paper, some new radiometric data have been published and presented for the late Carnian and Norian part of the continental Chinle Group. Ramezani et al. (2009 abstract and 2009 GSA talk in Portland, Oregon) presented new radiometric data and new stratigraphic data pertaining to previously reported radiometric dates from the Chinle section in Petrified Forest National

TRIASSIC CONCHOSTRACANS Park (Albuquerque, May, 2009). According to these new data, the 219.4 Ma age date was not derived from the basal Blue Mesa Member as initially reported, but from the middle of the overlying Sonsela Member. Thus, the estimated age of c. 219 Ma for the basal Blue Mesa no longer can be regarded as accurate. For the upper Blue Mesa Member, 223.1 Ma is now given. This upper Tuvalian value is not in contradiction with the 230.91 + 0.33 Ma age reported for the upper part of the lower Tuvalian section of the Lagonegro Basin (Italy) (Furin 2006) and also is not in contradiction with the 225 + 3 Ma age reported for the basal Norian of Alaska (Gehrels et al. 1987). As a whole, the radiometric data from the continental Late Triassic of the southwestern USA cannot be used reliably as yet for refining the numeric ages of the late Carnian and Norian. First, as shown by Dickinson & Gehrels (2008), the detrital zircons of the Chinle yield a wide range of age estimates for most of the members. Second, the lithostratigraphic assignment of some data is not well established (e.g. the 219.4 Ma age date reported for the basal Blue Mesa has been reassigned to the middle Sonsela, see above). Finally, the bio- and chronostratigraphic assignment of some members has not yet been determined. Thus, at present the Sonsela Member has a vertebrate fauna which could be found in either the uppermost Adamanian or in the Revueltian land vertebrate stages. According to Spencer Lucas (pers. comm. 2010), the Norian base could be within the Sonsela Member, but such an assumption can be neither confirmed nor rejected as yet. Until such time as well preserved conchostracans are found in the Sonsela Member or until either the Norian vertebrate guide form Aetosaurus or undoubtedly Tuvalian guide forms are found within it, the correct age assignment of the Sonsela will remain uncertain.

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Tetrapod footprints – their use in biostratigraphy and biochronology of the Triassic HENDRIK KLEIN1,* & SPENCER G. LUCAS2 1

Ru¨bezahlstraße 1, D-92318 Neumarkt, Germany

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New Mexico Museum of Natural History, 1801 Mountain Road NW, Albuquerque, NM 87104-1375 USA *Corresponding author (e-mail: [email protected]) Abstract: Triassic tetrapod footprints have a Pangaea-wide distribution; they are known from North America, South America, Europe, North Africa, China, Australia, Antarctica and South Africa. They often occur in sequences that lack well-preserved body fossils. Therefore, the question arises, how well can tetrapod footprints be used in age determination and correlation of stratigraphic units? The single largest problem with Triassic footprint biostratigraphy and biochronology is the nonuniform ichnotaxonomy and evaluation of footprints that show extreme variation in shape due to extramorphological (substrate-related) phenomena. Here, we exclude most of the countless ichnospecies of Triassic footprints, and instead we consider ichnogenera and form groups that show distinctive, anatomically-controlled features. Several characteristic footprint assemblages and ichnotaxa have a restricted stratigraphic range and obviously occur in distinct time intervals. This can be repeatedly observed in the global record. Some reflect distinct stages in the evolutionary development of the locomotor apparatus as indicated by their digit proportions and the trackway patterns. Essential elements are archosaur tracks with Rotodactylus, the chirotherian ichnotaxa Protochirotherium, Synaptichnium, Isochirotherium, Chirotherium and Brachychirotherium, and grallatorids that can be partly linked in a functional-evolutionary sequence. Non-archosaur footprints are common, especially the ichnotaxa Rhynchosauroides, Procolophonichnium, Capitosauroides and several dicynodont-related or mammal-like forms. They are dominant in some footprint assemblages. From the temporal distribution pattern we recognize five distinct tetrapod-footprint-based biochrons likened to the known land-vertebrate faunachrons (LVFs) of the tetrapod body fossil record: 1. Dicynodont tracks (Lootsbergian ¼ Induan age); 2. Protochirotherium (Synaptichnium), Rhynchosauroides, Procolophonichnium (Nonesian ¼ Induan–Olenekian age); 3. Chirotherium barthii, C. sickleri, Isochirotherium, Synaptichnium (‘Brachychirotherium’), Rotodactylus, Rhynchosauroides, Procolophonichnium, dicynodont tracks, Capitosauroides (Nonesian – Perovkan ¼ Olenekian– early Anisian); 4. Atreipus–Grallator (‘Coelurosaurichnus’), Synaptichnium (‘Brachychirotherium’), Isochirotherium, Sphingopus, Parachirotherium, Rhynchosauroides, Procolophonichnium (Perovkan –Berdyankian ¼ Late Anisian– Ladinian); 5. Brachychirotherium, Atreipus–Grallator, Grallator, Eubrontes, Apatopus, Rhynchosauroides, dicynodont tracks (Otischalkian–Apachean ¼ Carnian–Rhaetian). Tetrapod footprints are useful for biostratigraphy and biochronology of the Triassic. However, compared to the tetrapod body fossil record with eight biochrons, the five footprint-based biochrons show less resolution of faunal turnover as ichnogenera and ichnospecies at best reflect biological families or higher biotaxonomic units. Nevertheless, in sequences where body fossils are rare, footprints can coarsely indicate their stratigraphic age.

Tetrapod footprints of Triassic age are known from North America, South America, Europe, North Africa, China, Australia, Antarctica and South Africa (Figs 1 & 2). The Triassic footprint record is archosaur-, lepidosauromorph/ archosauromorph- (Rhynchosauroides) and synapsiddominated (Haubold 1971b, 1984; Klein & Haubold 2007), and it includes the oldest dinosaur tracks. Much has been written about Triassic tetrapod footprint biostratigraphy, especially based on the European and North American records (see below).

Our goal here is to present a Pangaea-wide Triassic biostratigraphy and biochronology based on tetrapod footprints. To do so, we briefly discuss some problems of footprint ichnotaxonomy and their bearing on footprint biostratigraphy (see Lucas 2007 for a more extensive review of these issues). We follow with a review of the principal Triassic tetrapod footprint assemblages. We conclude with a synopsis of Triassic tetrapod footprint biochronology that recognizes five biochrons and compare that biochronology to Triassic tetrapod biochronology based on body fossils.

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 419– 446. DOI: 10.1144/SP334.14 0305-8719/10/$15.00 # The Geological Society of London 2010.

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Fig. 1. Distribution of principal Triassic tracksites on Triassic Pangaea. Locations are: 1, Sydney basin, Australia; 2, Karoo basin, South Africa; 3, Antarctica; 4, western Europe, 5, Italy; 6, Chinle basin, western United States; 7, Newark basin, New Jersey; 8, Argentina; and 9, Yangtze basin, China. Base map after Wing & Sues (1992).

Fig. 2. Principal Triassic footprint horizons and footprint localities. German section and numerical age according to Menning & German Stratigraphic Commission (2002) and Bachmann & Kozur (2004).

TETRAPOD FOOTPRINTS

Triassic tetrapod footprint ichnotaxa The use of tetrapod footprints in biostratigraphy and biochronology depends on the following criteria: (1) a stable and uniform ichnotaxonomy; (2) ichnotaxa with limited stratigraphic (temporal) ranges; and (3) wide geographical distribution of the ichnotaxa used for correlation. Triassic tetrapod ichnotaxonomy began with the binominal naming of Chirotherium barthii and C. sickleri from the Bunter (Lower – Middle Triassic) of Hildburghausen, Germany, by Kaup (1835a, b). At present, a large number of ichnogenera and ichnospecies from the Triassic have been introduced by various authors (see Haubold 1971b, 1984; Klein & Haubold 2007 for reviews). Some are considered invalid, whereas others are still under discussion and may be extramorphological (substrate-related) variations. In Triassic ichnotaxonomy (and in tetrapod ichnotaxonomy in general), the influence of the substrate on footprint shape is still not adequately considered, nor fully understood. Work on this topic is in progress (e.g. Laporte & Behrensmeyer 1980; Cohen et al. 1991; Manning 2004; Diedrich 2005; Mila`n & Bromley 2006, 2008; Mila`n et al. 2004). Indeed, even the differentiation of ichnotaxa based on mathematical, including statistical, approaches (Demathieu 1998) may often reflect substrate rather than anatomical signals. Nevertheless, it can be helpful to quantify and demonstrate different morphologies of footprints by landmark analysis or other methods that are independent of the subjective view of the observer (Karl & Haubold 1998; Klein & Haubold 2003). Tetrapod ichnotaxa should be based on anatomical rather than substrate- or facies-controlled features. In contrast to invertebrate trace fossils, tetrapod footprints are identified by their relationship to the locomotor anatomy of supposed trackmakers. This focuses attention on the number, shape and proportions of digits, the shape of the sole surface and the trackway pattern. For example, different evolutionary developments are distinct in archosaur footprints and can be likened to the evolution in some lineages of dinosaurs as well as of the crocodylian stem-group and osteological taxa close to the crown-group of Archosauria (Haubold & Klein 2000, 2002; Haubold 2006; Lucas 2003, 2007; Hunt & Lucas 2007b; Klein & Haubold 2007). The distinct evolutionary stages reflected in some archosaur footprints in these lineages have a restricted stratigraphic/temporal range within the Triassic, as do some osteological taxa (Lucas 1998, 1999). This makes them useful for biostratigraphy and biochronology. Nevertheless, ichnospecies cannot be attributed to species based on body fossils (Peabody 1955a). They are, instead, equivalent to osteological genera or families that have lower evolutionary

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turnover rates (Lucas 2007). Moreover, evolution may have only affected parts of the body other than the limbs. For example, tridactyl theropod footprints of the Grallator–Eubrontes type are of similar shape from the Late Triassic to the Jurassic. Thus, the foot morphology of theropod trackmakers was obviously consolidated in the Carnian –Norian and subsequently underwent no essential change. Therefore, a zonation based on tetrapod footprints is possible at only a relatively low resolution. Nevertheless, the intercontinental distribution of Triassic ichnotaxa such as Rotodactylus, Chirotherium, Grallator or Eubrontes and their abundance (Peabody 1948, 1955b; Haubold 1967, 1971a, b; Tresise & Sarjeant 1997; Haubold & Klein 2000; Lu¨ et al. 2004; King et al. 2005; Lucas et al. 2006a; Klein & Haubold 2007) enables the correlation of sequences with an otherwise poorly known body fossil record. Attempts to outline a tetrapod ichnostratigraphy of Triassic sequences were made by Haubold (1969, 1971b, 1984, 1986), Demathieu & Haubold (1972, 1974), Olsen (1980, 1983), Lockley & Hunt (1995), Hunt & Lucas (2007a, b), Lucas (2003, 2007) and Klein & Haubold (2007). Further contributions can be found in Ellenberger (1970, 1972, 1974), Demathieu (1984), Olsen & Galton (1984), Olsen & Baird (1986), Silvestri & Olsen (1989), Silvestri & Szajna (1993), Lockley et al. (1996), Szajna & Silvestri (1996), Avanzini et al. (2001), Lucas & Hancox (2001), Olsen et al. (2002), Szajna & Hartline (2003), Lucas & Huber (2003), Lucas & Tanner (2004), Gand et al. (2005), Lucas & Sullivan (2006) and Lucas et al. (2006a, b, c). Presently, the non-uniform ichnotaxonomy of Triassic tetrapod ichnotaxa hinders a generally accepted and conclusive concept. Many of the described ichnospecies are dubious, except some well-substantiated ichnotaxa such as Chirotherium barthii and C. sickleri, among others. Therefore, ichnogenera are considered by us to be the operational taxonomic units of a Triassic footprint biostratigraphy and biochronology.

Triassic tetrapod footprint assemblages Lowest Triassic Tetrapod footprints from this interval have been described from the Karoo basin (South Africa), the southern Sydney basin of Australia, and from Antarctica (Watson 1960; MacDonald et al. 1991; Retallack 1996). At these localities, the sequences largely straddle the Permian–Triassic boundary, and the footprints are in strata that in South Africa and Antarctica can be assigned to the Lootsbergian land –vertebrate faunachron (LVF) of latest Changshingian –Induan age (Lucas 1998, 2007;

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Hunt and Lucas 2007b). The poorly-preserved footprints from Australia were assigned to Dicynodontipus and attributed to dicynodont trackmakers, in particular Lystrosaurus (Retallack 1996); however, this relationship cannot be demonstrated conclusively. From the Bunter of Poland (Wio´ry, Holy Cross Mountains), extensive surfaces with footprints are known (Fig. 3). They occur within a fluvial sequence, the ‘Labyrinthodontidae beds’, which can be considered to be of late Induan to Olenekian age (Ptaszynski 2000; Niedzwiedzki & Ptaszynski 2007). Archosaur tracks of the chirothere morphological group dominate these assemblages. Indeed, Fuglewicz et al. (1990), Ptaszynski (2000) and Niedzwiedzki & Ptaszynski (2007) identified Synaptichnium, Brachychirotherium and Isochirotherium (Fig. 3b– d). The presence of the two latter forms, however, is dubious because the preserved footprint morphology does not match the diagnoses of these ichnotaxa (Beurlen 1950; Haubold 1971b; Karl & Haubold 1998, 2000). Instead, all these imprints from Poland reflect conservative foot structures because of their long pedal digits IV and V. In their overall shape, they resemble Protochirotherium wolfhagense, a chirotherian ichnotaxon described by Fichter & Kunz (2004) from the Detfurth Formation (Middle Bunter, Olenekian) of northern Hessen, Germany (Figs 3a & 4). The material from Poland was therefore referred to Protochirotherium by Klein & Haubold (2007). Possible trackmakers are archosauriforms close to the base of the crown-group Archosauria (Sereno 1991). Further characteristic footprints described from the Polish Bunter are Rhynchosauroides (lepidosauromorph and/or archosauromorph), Procolophonichnium (primitive amniote), Capitosauroides (amphibian) and a purported new ichnogenus, Prorotodactylus (Fuglewicz et al. 1990; Ptaszynski 2000; Fig. 3e–f ). Footprint surfaces in the Hardegsen Formation (Middle Bunter, Olenekian) of Northern Hessen, Germany, record Synaptichnium, ‘Isochirotherium’ and Rhynchosauroides (Demathieu & Haubold 1982).

Upper Lower Triassic to Lower Middle Triassic In this stratigraphic interval, the diversity of imprint forms increases. A large number of archosaur footprint taxa can be discriminated by different digit proportions and trackway pattern. Thus, chirotheres are represented by four ichnogenera: Synaptichnium, Brachychirotherium, Isochirotherium and Chirotherium (Haubold 1971b). Synaptichnium (Fig. 5a, b) reflects primitive foot morphology in its long pedal digits IV and V. It continues from the Induan and Olenekian into

the Anisian. It is part of the assemblages in the Moenkopi Formation of Arizona (Peabody 1948), and the Middle Triassic of Great Britain (Tresise & Sarjeant 1997; King et al. 2005). The presence of Brachychirotherium in the Lower and Middle Triassic is dubious. Forms that have been described under this name from the Bunter (Olenekian–Anisian) of Germany and the Middle Triassic (Anisian –Ladinian) of France and Argentina (Haubold 1971b; Demathieu & Gand 1973; Courel & Demathieu 1976; Demathieu & Demathieu 2004; Melchor & De Valais 2006) are mostly extramorphological variants of Synaptichnium (Fig. 5c–d). The substrate-controlled transition to Synaptichnium has been demonstrated by material from the German Middle Triassic (Anisian) by Klein & Haubold (2004). Furthermore, the analysis and revision of the type material of Brachychirotherium from the Upper Triassic (Carnian) of Germany by Karl & Haubold (1998, 2000) revealed that the inclusion of these Lower to Middle Triassic forms in the ichnogenus is problematic. The forms from the Lower and Middle Triassic are thus referred to here as ‘Brachychirotherium’ (Klein & Haubold 2007). Isochirotherium (Fig. 5e–g) is a characteristic ichnotaxon of the Middle Triassic. Characteristic features are the dominance of digits II and III in the pes and the comparatively small manus (Haubold 1971b). This ichnogenus is present in footprint assemblages of the German Bunter (‘Thu¨ringischer Chirotheriensandstein’), the Middle Triassic of Great Britain (Tresise & Sarjeant 1997, King et al. 2005) and the Moenkopi Formation of Arizona (Peabody 1948). Possible trackmakers are crurotarsans. In Chirotherium (Figs 5h –j & 6), more advanced foot structures are visible in the reduction and posterior shift of pedal digit I, and in the dominance of the digit group II-IV. Other characteristic features are the reduction of manual digits IV and V and the narrow trackway pattern (Walther 1917; Soergel 1925; Haubold 1971a, 2006). In particular, Chirotherium barthii (Figs 5h, j & 6a, c –e) displays an initial trend toward the tridactyl pattern later seen, for example, in the autopodia of dinosaurs. Chirotherium sickleri (Figs 5i & 6b) shows a more conservative pattern, with long pedal and manual digits IV. The taxonomic status of some large chirotherians from the Moenkopi Formation is questionable, such as Chirotherium moquiense and C. rex (Peabody 1948, 1956). The Moenkopi material is presently under revison by the authors. Assemblages with Chirotherium are known from the Bunter of Germany (‘Thu¨ringischer Chirotheriensandstein’, Haubold 1971a, 2006), Great Britain (Tresise & Sarjeant 1997; Clark et al. 2002; King et al. 2005), France (Demathieu 1984), Spain

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Fig. 3. Characteristic footprints from the lowest Triassic. (a) Protochirotherium wolfhagense. (b, c) Protochirotherium (Brachychirotherium, Isochirotherium after Fuglewicz et al. 1990; Ptaszynski 2000). (d) Synaptichnium. (e) Rhynchosauroides. (f) Procolophonichnium. (a) Detfurth Formation (Olenekian), northern Hessen, Germany. (b– f) Labyrinthodontidae beds (Induan), Wio´ry, Poland. After Fuglewicz et al. (1990), Klein & Haubold (2007), Ptaszynski (2000).

(Calderon 1897), Arizona (Moenkopi Formation: Peabody 1948), Argentina (Rusconi 1951; Peabody 1955b; Melchor & De Valais 2006), and China (Lu¨ et al. 2004).

Another characteristic ichnotaxon of this stratigraphic interval is Rotodactylus (Fig. 5k –l). These are small pes and manus impressions, often in dense concentrations on track surfaces. The pes

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Fig. 4. Protochirotherium wolfhagense Fichter & Kunz, 2004 (a) holotype; (b) additional specimen from surface type. Pes and manus imprints with skin structure from the Detfurth Formation (Olenekian) of northern Hessen, Germany. Photos: S. Voigt, Freiberg. Scale bars in cm.

shows a dominance of digits II, III and IV, with IV the longest. Digit V is in an extremely posterior position and only visible as a small circular impression. The manus is of similar shape, and in digit IV it may be shorter. Trackways show wide lateral overstep of the manus by the pes and a long stride length, depending on the velocity of movement. Peabody (1948) was the first to posit a cursorial, dinosaur-like animal (pseudosuchian) as the trackmaker of Rotodactylus. Haubold (1967, 1971a, b, 1999) described Rotodactylus from surfaces of the German Bunter (‘Thu¨ringischer Chirotheriensandstein’) and attributed it to a dinosauromorph producer. Rotodactylus is known in particular from assemblages of the Moenkopi Formation of Arizona (Peabody 1948), the German Bunter (Haubold 1967, 1971a, b, 1999) and the Middle Triassic of France (Demathieu 1984). Non-archosaur footprints co-occur with Rhynchosauroides (Fig. 7a). Rhynchosauroides is a small, pentadactyl and lacertoid imprint form with slender, inwardly curved digits I –IV. Digit IV is longest, and digit V is short and postero-laterally positioned. The pes and manus are of similar shape. Trackways mostly show lateral overstep of the manus by the pes. Rhynchosauroides is present in most assemblages of this stratigraphic interval

(e.g. Peabody 1948; Tresise & Sarjeant 1997; Demathieu 1984; Haubold 1971a; Melchor & De Valais 2006). Therapsid footprints (Fig. 7c, d) are known from the ichnogenera Dicynodontipus from the Bunter of Germany (Haubold 1966, 1971a; Demathieu & Fichter 1989) and Therapsipus from the Moenkopi Formation (Holbrook Member) of Arizona (Hunt et al. 1993). The morphology is characterized by a pentadactyl pes and manus of similar shape and size with short and straight digits. In the smaller Dicynodontipus (Fig. 7c), they are of subequal length, and digit V is shifted backward. Some trackways show the manus overstepped by the pes, dependent on the velocity of movement. Therapsipus (Fig. 7d) is a large form (25 cm pes length). Following Hunt et al. (1993), digit III in the pes and digits III and IV in the manus are longest. Procolophonichnium (Fig. 7b) encompasses the footprints of small, primitive amniotes. The pes and manus imprints are of similar morphology. They are semi-plantigrade to digitigrade and pentadactyl. Digit length increases from I to III, digits IV and III are subequal in length, and digit V is shorter. Trackways are broad, with the smaller manus

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Fig. 5. Archosaur footprints from the Upper Lower Triassic –Lower Middle Triassic. (a, b) Synaptichnium; (c, d) ‘Brachychirotherium’; (e – g), Isochirotherium; (h, j) Chirotherium barthii; (i) C. sickleri; (k, l) Rotodactylus. (c, e, h, i, k) ‘Thu¨ringischer Chirotheriensandstein’ (Bunter), southern Thuringia. (d) Ro¨t Formation (Bunter), northern Bavaria. (a, f) Bunter of Great Britain. (b, g, j, l) Moenkopi Formation, Arizona. After Baird (1954), Haubold (1971a, b), Peabody (1948), Soergel (1925).

anterior to the pes. They are known, for example, from the German Bunter (Ru¨hle von Lilienstern 1939; Haubold 1971a). Capitosauroides (Fig. 7e) refers to semiplantigrade, pentadactyl pes and manus imprints

with short, blunt and clawless digits that diverge widely. They have been attributed to amphibians (Haubold 1971a). The Bunter of Germany and the Moenkopi Formation of Arizona have provided examples (Haubold 1971a; Peabody 1948).

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Fig. 6. Archosaur footprints from the Upper Lower Triassic– Lower Middle Triassic. (a, c – e) Chirotherium barthii; (b) Chirotherium sickleri; (a, b) Moenkopi Formation, Arizona; (c) ‘Thu¨ringischer Chirotheriensandstein’ (Bunter), southern Thuringia; (d) Cerro de Las Cabras Formation, Argentina; (e) Guanling Formation, Guizhou Province, China. Photos: (c) D. Hildebrand; (e) H. Lu¨; others, H. Klein.

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Fig. 7. Non-archosaur footprints from the Upper Lower–Lower Middle Triassic. (a) Rhynchosauroides; (b) Procolophonichnium; (c) Dicynodontipus; (d) Therapsipus; (e) Capitosauroides; (a–c, e) Bunter Germany. (d) upper Moenkopi Formation, Arizona (after Haubold 1966, 1971a, b, 1984; Hunt et al. 1993).

Upper Middle Triassic Anisian –Ladinian footprint assemblages are well known from the eastern border of the Massif Central in France (Demathieu 1970; Courel & Demathieu 1976; Gand 1976, 1979a, b), Germany (Haubold & Klein 2002; Klein & Haubold 2004) and Italy (Avanzini 1999, 2000; Avanzini & Lockley 2002; Avanzini & Neri 1998; Mietto 1987). Characteristic ichnogenera are Isochirotherium, Synaptichnium and ‘Brachychirotherium’ (Figs 8a– f & 9a–d), continuing from the lower Middle Triassic. Sphingopus and Parachirotherium (Figs 8g–i & 9e–g) are characteristic upper Middle Triassic ichnogenera. They refer to pentadactyl footprints that follow the trend seen in Chirotherium

barthii: (1) reduction and posterior shift of pedal digit I and development of the mesaxonic tridactyl pattern II, III, IV, with digit III longest; and (2) decrease of the palmar surface, indicating a tendency toward bipedality (Haubold & Klein 2000, 2002). Sphingopus ferox was named and described by Demathieu (1966) from the Massif Central (France). The ichnogenus was also documented in strata of similar age in Northern Bavaria, Germany (Haubold & Klein 2002). Parachirotherium postchirotherioides is known from surfaces of the Benker Sandstein (upper Ladinian) of Northern Bavaria (Rehnelt 1950; Kuhn 1958; Haubold & Klein 2000). Significant is the appearance of the first truly tridactyl pes imprints [Atreipus–Grallator (‘Coelurosaurichnus’)] in trackways showing the transition

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Fig. 8. Upper Middle Triassic archosaur footprints. (a, b) Isochirotherium; (c, d) ‘Brachychirotherium’; (e, f) Synaptichnium; (g, h) Sphingopus; (i) Parachirotherium; (j) Rotodactylus; (k) Rigalites. (a–e, g, j) Anisian–Ladinian, Massif Central, France. (f, h, i) marginal facies of the Muschelkalk and Benker Sandstein (Anisian– Ladinian), northern Bavaria. K, Los Rastros Formation (Ladinian), Argentina. After Courel & Demathieu (1976); Demathieu (1966, 1970); Demathieu & Gand (1973); Gand (1976, 1979a); Haubold (1971b, 1984); Haubold & Klein (2000); Klein & Haubold (2004, 2007).

from quadrupedal to bipedal gait, thus indicating facultative bipedality of probable dinosauriform trackmakers in the Benker Sandstein (upper Ladinian) of Northern Bavaria, Germany (Weiss 1976; Haubold & Klein 2000). Similar trackways appear in the Anisian-Ladinian of France (Gand 1976, 1979a; Gand & Demathieu 2005; Demathieu 1989; Figs 10 & 11a –c).

Rotodactylus occurs in dense concentrations on surfaces of the upper Middle Triassic of France (Demathieu & Gand 1973; Gand 1976, 1979a; Fig. 8j). Rigalites (Fig. 8k) is a problematic ichnotaxon of archosaur affinity from the Los Rastros Formation (?Ladinian) of San Juan Province, Argentina (Huene 1931; Marsicano et al. 2004; Melchor &

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Fig. 9. Upper Middle Triassic archosaur footprints. (a) ‘Brachychirotherium’; (b) Synaptichnium; (c, d) Isochirotherium; (e, f) Sphingopus; (g) Parachirotherium. (a, d, e) Anisian–Ladinian, Massif Central, France. (c) Anisian, northern Italy. (b, f, g) Anisian–Ladinian, northern Bavaria. Photos: (c) Marco Avanzini; (d, e) Georges Gand; others, H. Klein.

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Fig. 10. Tridactyl pes and partly associated manus imprints of dinosauromorph archosaurs from the upper Middle Triassic, Atreipus-Grallator (‘Coelurosaurichnus’). (a) Benker Sandstein (upper Ladinian), northern Bavaria. (b–e) Anisian–Ladinian, Massif Central, France. After Gand & Demathieu (2005), Haubold (1984), Haubold & Klein (2000).

De Valais 2006). It encompasses large tetradactyl to pentadactyl pes imprints (more than 35 cm long) associated with a smaller, pentadactyl manus imprint. It shows similarities to large Parachirotherium or other chirotherians. Other bipedal to quadrupedal trackways of possible dinosauriform affinity are described from this unit by Marsicano et al. (2007). From Alpine regions in Switzerland, Demathieu & Weidmann (1982) describe an extensive surface from the Ladinian –Carnian boundary. It shows chirotherians and hundreds of undeterminable archosaur footprints and trackways. Also present in upper Middle Triassic strata are Rhynchosauroides and Procolophonichnium (Figs 11d & 12a –d). They are dominant in carbonate tidal flat deposits along the margin of the Bohemian and Rhenish massifs in Germany and The Netherlands (Demathieu & Oosterink 1983, 1988; Diedrich 2007), where archosaur footprints are rare.

Upper Triassic Upper Triassic tetrapod footprint assemblages are archosaur dominated. Most common are tracks of Brachychirotherium (sensu stricto) and tridactyl imprints of the Grallator-Eubrontes type.

Brachychirotherium (sensu stricto; Figs 13 & 16a –c) was first named and described by Beurlen (1950) from the Middle Keuper (‘Coburger Sandstein’, upper Carnian) of northern Bavaria, Germany. Diagnostic features are the broad, pentadactyl pes impression with short, thick digits and tiny claws. Digit V is represented only by an oval basal pad in a position postero-lateral to the rest of the foot. The smaller manus has a similar shape. Karl & Haubold (1998, 2000) analysed the type material and designated a lectotype of the type species Brachychirotherium hassfurtense and also confirmed the validity of a second ichnospecies in the Keuper, B. thuringiacum. Brachychirotherium is widespread in the Newark Supergroup of New Jersey and Pennsylvania (Baird 1957; Silvestri & Olsen 1989; Silvestri & Szajna 1993; Szajna & Hartline 2003; Szajna & Silvestri 1996; Olsen et al. 2002; Lucas & Sullivan 2006) and in the Chinle Group of the North American Southwest (Lockley & Hunt 1993, 1995; Lucas et al. 2001, 2006b, c; Klein et al. 2006; Hunt & Lucas 2007a). It is also known from the Los Colorados Formation (Norian) of Argentina (Arcucci et al. 2004; Melchor & De Valais 2006), and the Lower Elliot Formation (Norian) of South Africa (Ellenberger 1970, 1972, 1974; Olsen & Galton

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Fig. 11. Footprints from the upper Middle Triassic. (a– c) Atreipus–Grallator (‘Coelurosaurichnus’). (a, b) Anisian– Ladinian of the Massif Central (France). (c) Benker Sandstein (upper Ladinian), northern Bavaria. (d) Rhynchosauroides, Anisian of Winterswijk (The Netherlands). Photos (d) Henk W. Oosterink; others, H. Klein.

1984; D’Orazi Porchetti & Nicosia 2007). Probable trackmakers are crurotarsans. There is some controversy among tetrapod ichnologists concerning the presence and validity of other ichnogenera used for Upper Triassic tracks, especially Pseudotetrasauropus, Tetrasauropus, Evazoum, Eosauropus, and Otozoum (Fig. 14). The names Pseudotetrasauropus and Tetrasauropus were originally given to material from the Lower Elliot Formation of Lesotho, southern Africa that was attributed to sauropodomorphs (Ellenberger 1970, 1972, 1974; D’Orazi Porchetti & Nicosia 2007; Fig. 14a–b). Later, footprints

from the Chinle Group of the American Southwest, from Great Britain and from Poland were also assigned to these ichnotaxa (Lockley & Hunt 1993, 1995; Lockley & Meyer 1999; Lockley et al. 1993, 1996, 2000, 2001; Lucas et al. 2001; Fig. 14f –h; Gierlinski 2007). Some were finally re-assigned to Evazoum, an ichnotaxon named from the Carnian of Italy (Nicosia & Loi 2003; Lockley et al. 2006b; Fig. 14e) and to Eosauropus, another ichnotaxon based on specimens from the American Southwest (Lockley et al. 2006a; Fig. 14c). From footprint surfaces of the Norian of France, Gand et al. (2000) described Otozoum, a

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Fig. 12. Non-archosaur footprints from the upper Middle Triassic. (a) Procolophonichnium (‘Circapalmichnus’). (b, d) Rhynchosauroides. (c) Procolophonichnium. (a) Anisian– Ladinian of Massif Central (France); (b) Anisian of northern Italy. (c, d) Anisian of Winterswijk (The Netherlands). After Avanzini & Renesto (2002), Demathieu & Oosterink (1983), Gand (1977), Haubold (1984).

taxon known from the Lower Jurassic (Fig. 14d). Rainforth (2003) referred the trackways from France to Pseudotetrasauropus. Klein et al. (2006), however, considered all these forms to be extramorphological variants of Brachychirotherium. They demonstrated the substrate-controlled transition between these forms on track surfaces of the Upper Triassic Redonda Formation (Chinle Group) of New Mexico. Thus, because of disagreements over the ichnotaxonomy of these forms, as

well as their sporadic (though geographically widespread) record, they are not of significance to Triassic footprint biostratigraphy at present. Grallator and Eubrontes (Figs 15d–i & 16d, f ) are tridactyl, mesaxonic footprints of different sizes (pes length of Eubrontes . 25 cm). They were first described from the Lower Jurassic of the Newark Supergroup by Hitchcock (1836, 1845, 1847, 1858) and can be assigned to theropod trackmakers [Weems (2003) arguments for a prosauropod

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Fig. 13. Pes and manus imprints of Brachychirotherium. (a) B. hassfurtense (type species); (b, c) B. thuringiacum from the upper Carnian of Germany. (d) B. parvum, Newark Supergroup (Passaic Formation, Norian), North America. (e, g) Chinle Group (Redonda Formation, Norian), North America. (f) Los Colorados Formation (Norian), Argentina. After Baird (1957), Haubold (1971b), Karl & Haubold (1998), Klein et al. (2006).

trackmaker of Eubrontes have not been accepted by other ichnologists]. The type material has been re-located and re-examined by Olsen et al. (1998). These theropod footprint ichnogenera are widespread in the Upper Triassic of Pangaea. There is evidence of Grallator and/or Eubrontes from

countless surfaces in North America (Chinle Group, Newark Supergroup), Greenland (Fleming Fjord Formation), Great Britain (Mercia Mudstone Group), Sweden (Ho¨gana¨s Formation), France (Carnian –Norian of Arde`che and Ale`s), Germany (Keuper, Carnian–Norian of Bavaria), Switzerland

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Fig. 14. Imprint forms evaluated differently by various authors either as distinct ichnotaxa or extramorphological (substrate-related) variations of Brachychirotherium. (a) Pseudotetrasauropus. (b) Tetrasauropus. (c) Eosauropus. (d) Otozoum. (e – h) Evazoum. (a, b) Lower Elliot Formation (Norian), southern Africa. (c, f– h) Chinle Group (Norian), North America. (d) Norian, France. (e) Carnian, Italy. After D’Orazi Porchetti & Nicosia (2007), Gand et al. (2000), Klein et al. (2006), Lockley et al. (2006a), Nicosia & Loi (2003).

(Diavel Formation, Norian), Southern Africa (Lower Elliot Formation) and Argentina (Carnian) (e.g. Weiss 1934; Ellenberger 1970, 1972; Furrer 1993; Gierlinski & Ahlberg 1994; Jenkins et al. 1994; Lockley & Hunt 1995; Lockley et al. 1996; Gatesy et al. 1999; Haubold & Klein 2000, 2002; Courel & Demathieu 2000; Gand et al. 2000, 2005; Olsen et al. 2002; Gaston et al. 2003; Szajna & Hartline 2003; Marsicano & Barredo 2004; Mila`n et al. 2004; Gand & Demathieu 2005; Klein et al. 2006; Lucas et al. 2006a). Occasionally, a manus imprint is observed in association with footprints of the Grallator morphology (Olsen & Baird 1986; Courel & Demathieu 2000; Gand & Demathieu 2005; Figs 15a, b, c & 16e). For these tracks, Olsen & Baird (1986) introduced a

distinct ichnogenus and named it Atreipus. Haubold & Klein (2000) combined grallatorid imprints of facultative bipedal trackways in the plexus Atreipus –Grallator. Some authors still use the name ‘Coelurosaurichnus’ for these forms (Courel & Demathieu 2000; Gand & Demathieu 2005; Gand et al. 2005), and it can be found in the older literature (e.g. Huene 1941; Beurlen 1950; Heller 1952; Kuhn 1958). The following archosaur-footprint ichnotaxa are also present but generally less widespread and abundant in the Upper Triassic: (1) Chirotherium lulli (Fig. 17a), a small chirotherian known from trackways from the Newark Supergroup (Baird 1954); (2) Apatopus (Fig. 17b) occurs in the Newark Supergroup and the Chinle Group of North America (Baird

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Fig. 15. Tridactyl archosaur tracks from the Carnian through Lower Jurassic. Atreipus–Grallator with occasionally associated manus and Grallator– Eubrontes (bipedal theropod footprints). (a) Carnian of southern France. (b) Ansbacher Sandstein (Carnian) of northern Bavaria. (c) Norian of Newark Supergroup, North America. (d) Coburger Sandstein (upper Carnian) of northern Bavaria. (e) Chinle Group (Norian), North America. (f) Norian, France. (g) Rhaetian, Sweden. (h, i) Lower Jurassic of Newark Supergroup, North America (types of Eubrontes and Grallator). After Courel & Demathieu (2000), Gand & Demathieu (2005), Gaston et al. (2003), Gierlinski & Ahlberg (1994), Haubold & Klein (2000), Hunt et al. (2000), Olsen & Baird (1986), Olsen et al. (1998), Hunt et al. (2000).

1957; Foster et al. 2003). It shows a conservative morphology of long pedal digits IV and V and a large manus. Apatopus has been widely attributed to phytosaurs, however, there is no conclusive evidence supporting this interpretation; (3) Batrachopus, an ichnotaxon known from the Jurassic by quadrupedal trackways with tetradactyl to

pentadactyl pes- and manus-imprints, might be also present in the Triassic of the Newark Supergroup. It is attributed to crocodylomorphs. Bird-like footprints have been described from the Santo Domingo Formation (Upper Triassic –Lower Jurassic) of northwestern Argentina and referred to the ichnogenus Gruipeda (Melchor et al. 2002; De

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Fig. 16. Characteristic archosaur footprints from the Upper Triassic. (a) pes of Brachychirotherium hassfurtense (type species). (b) pes and manus of B. thuringiacum. (c) pes and manus of B. parvum. (d, e) Atreipus–Grallator. (f) Grallator. (a, b, d, e) Coburger Sandstein (upper Carnian) of northern Bavaria. (c, f) Chinle Group (Redonda Formation, Norian) of North America.

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Fig. 17. Imprint forms and trackways from the Upper Triassic. (a) Chirotherium lulli. (b) Apatopus. (c) Rhynchosauroides. (d) Procolophonichnium. (e) Gwyneddichnium. (f) probable mammal-like (synapsid) trackway. (a, b, d, e) Norian, Newark Supergroup, North America. (c, f) Chinle Group and Wingate Sandstone (Norian– Rhaetian), North America. After Baird (1954, 1957), Haubold (1971b), Lockley et al. (2004), Silvestri & Szajna (1993).

Valais & Melchor 2008). Similar forms are known from the Lower Elliot Formation of southern Africa (Trisauropodiscus: Ellenberger 1972). Surfaces with grallatorids and other undetermined archosaur footprints have been described

from the Carnian– Norian of Italy (Dalla Vecchia 1996; Dalla Vecchia & Mietto 1998). Non-archosaur footprints from Upper Triassic strata are Rhynchosauroides, Procolophonichnium, Gwyneddichnium, and mammal-like forms (Bock

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1952; Baird 1957, 1964; Kuhn 1958; Haubold 1971b, 1984; Olsen 2002; Silva et al. 2008; Fig. 17c –f ). Mammal-like imprints similar to Brasilichnium occur on some surfaces of the Chinle Group in the American Southwest (Lockley et al. 2004; Klein et al. 2006). From South America, an assemblage dominated by trackways of small therapsids is known from the Carnian of the Vera Formation in Rio Negro Province of Argentina (Casamiquela 1964; Domnanovich & Marsicano 2006; Melchor & De Valais 2006).

Triassic tetrapod footprint biochronology The composition and distribution of Triassic tetrapod footprint assemblages reflect ecological/taphonomical peculiarities as well as different directions and stages in the evolutionary development of the locomotor apparatus of some tetrapod groups. In particular, some archosaur footprints show a limited vertical (stratigraphic) range. Their occurrences are restricted to distinct time intervals, thus demarcating distinct biochronological units (Lucas 2003, 2007; Hunt & Lucas 2007b; Klein & Haubold 2007; Fig. 18). Between the late Olenekian/Anisian and the Norian the development of the tridactyl mesaxonic foot and bipedal gait of dinosaurs is reflected by the footprint record and can be followed in a functional evolutionary succession: Chirotherium–Sphingopus–Parachirotherium– Atreipus– Grallator (Haubold & Klein 2000, 2002). This has been used for biostratigraphy and biochronology by Klein & Haubold (2007). Thus, Chirotherium spans the Olenekian –Anisian, Sphingopus the Anisian –Ladinian, Parachirotherium the Ladinian, Atreipus the Carnian– Norian and Grallator the Norian –Rhaetian interval. Klein & Haubold (2007) discriminated six biochrons (I–VI) by the range of archosaur footprint assemblages. The beginning of each is marked by the first appearance datum (FAD) of a characteristic index ichnotaxon (in bold): I. Protochirotherium, Late Induan–Olenekian; II. Chirotherium, Rotodactylus, Isochirotherium, Synaptichnium (‘Brachychirotherium’), Late Olenekian –Anisian; III. Sphingopus –Atreipus– Grallator, Rotodactylus, Isochirotherium, Synaptichnium (‘Brachychirotherium’), Late Anisian – Ladinian; IV. Parachirotherium– Atreipus–Grallator, Synaptichnium (‘Brachychirotherium’), Late Ladinian; V. Atreipus–Grallator, Brachychirotherium, Carnian –Norian and VI. Grallator –Eubrontes, Brachychirotherium, Norian –Rhaetian. Lucas (2003, 2007) recognized five Triassic footprint assemblages: 1. Dicynodont tracks, earliest Triassic; 2. Chirothere, Olenekian –Anisian; 3. Procolophonichnium–Rhynchosauroides, Anisian– Ladinian; 4. Dinosauromorph, Ladinian –Carnian;

and 5. Dinosaur, Carnian– Rhaetian. In this scheme, 2 corresponds to II and III, 3 to III, 4 to IV, and 5 to V and VI of Klein & Haubold (2007). Hunt & Lucas (2007b) propose five assemblages: 1. Dicynodont tracks, earliest Triassic; 2. Chirothere, Olenekian–early Anisian; 3. Dinosauromorph, late Anisian –Ladinian; 4. Tridactyl dinosaur, Carnian – early Norian; and 5. Sauropodomorph, late Norian–Rhaetian. In addition to Lucas (2003, 2007), Hunt & Lucas (2007b) recognize a sauropodomorph track assemblage in the Late Norian–Rhaetian based on the taxa Evazoum and Eosauropus (see above) purportedly first appearing in the late Norian. This is contrary to Klein et al. (2006) and Klein & Haubold (2007), who considered the footprints of Evazoum to be extramorphological variants of Brachychirotherium, a crurotarsan and characteristic of the entire Late Triassic. Furthermore, Evazoum was first described from the Carnian of Italy by Nicosia & Loi (2003), thus indicating an earlier appearance. Independent from further subdivisions proposed by various authors, we can recognize five tetrapod footprint biochrons of Triassic age that can be identified across the Pangaean footprint record: 1.

2.

3.

Earliest Triassic dicynodont footprints. These tracks are from strata of the Lystrosaurus assemblage zone and thus are of Lootsbergian (¼ latest Changshingian– Induan) age (Lucas 1998; Lucas et al. 2007). However, there are only a few records of this assemblage and they are restricted to Gondwana, so it needs further documentation before its Pangaea – wide significance can be established. Protochirotherium is characteristic of strata of Nonesian age (¼ Olenekian). Morphologically, and based on its temporal distribution, it can be considered as the hypothetical ‘root’ of later locomotory developments in archosaurs. Associated forms are Rhynchosauroides, Procolophonichnium and footprints of temnospondyls. The appearance of Chirotherium barthii and C. sickleri, Rotodactylus, Isochirotherium and Synaptichnium (‘Brachychirotherium’) roughly demarcates the Nonesian –Perovkan (late Olenekian– Anisian) transition. Chirotherium barthii and C. sickleri disappear during the Anisian. The range of the other ichnotaxa spans most of the Middle Triassic (Perovkan – Berdyankian ¼ Anisian –Ladinian) together with Rhynchosauroides, Procolophonichnium, dicynodont and temnospondyl footprints that continue from the Nonesian. Rotodactylus and Isochirotherium disappear before the end of the Berdyankian (Ladinian).

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Fig. 18. Stratigraphic distribution of tetrapod footprint taxa and form groups in the Triassic with the biochrons and characteristic assemblages recognized by different authors. Range chart of archosaur footprints after Klein & Haubold (2007).

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The appearance of tridactyl footprints and quadrupedal to bipedal trackways of the Atreipus–Grallator type (‘Coelurosaurichnus’) demarcates the late Perovkan –Berdyankian (¼ late Anisian –Ladinian) as do pentadactyl footprints of Sphingopus and Parachirotherium. Other ichnotaxa continue from the Nonesian (see above). Brachychirotherium (sensu stricto) appears at the beginning of the Otischalkian (¼ early Carnian). It is a characteristic ichnotaxon of the Late Triassic, together with Atreipus– Grallator (quadrupedal to bipedal trackways), Grallator and Eubrontes (bipedal trackways). The stratigraphical upper limit of Brachychirotherium is the Triassic –Jurassic boundary (end of the Apachean); there is no evidence of Brachychirotherium in post-Triassic strata (Lucas & Tanner 2007a, b). The same is true for other chirotheres, and for Apatopus, Procolophonichnium and Gwyneddichnium. The range of Rhynchosauroides continues into the Jurassic (Olsen et al. 2002), and the same is true of Batrachopus and the mammal-like forms, as might be expected.

Rhynchosauroides and Procolophonichnium, as well as some dicynodont and temnospondyl footprints, have a long stratigraphic range. They span the complete Triassic Period. Therefore, they are of less utility for biostratigraphy as long as their taxonomy is unclear, as is the case presently. Their dominance in some assemblages (see above) is extremely facies-controlled and biased by ecological and taphonomical effects. Rhynchosauroides and Procolophonichnium trackmakers obviously frequented some Anisian –Ladinian carbonate tidal flats (assemblage 3 of Lucas 2003, 2007; Demathieu & Oosterink 1983, 1988; Diedrich 2007), an environment that archosaurs (chirothere trackmakers) mostly avoided. However, a few chirotheres are present as well (Demathieu & Oosterink 1988) and, on the other hand, Rhynchosauroides is common at least on some fluvial-lacustrine surfaces of the Early through Late Triassic (Demathieu 1966; Haubold 1971a, b). Evolutionary rather than facies-controlled signals from footprints are suitable to demarcate distinct time intervals in the Triassic and to outline a coarse biostratigraphy and biochronology of the Triassic. This footprint biochronology identifies five intervals of Triassic time, which is less resolution than the eight land-vertebrate faunachrons of Triassic age based on tetrapod body fossils (Fig. 18).

Conclusions Tetrapod footprints are useful in biostratigraphy and biochronology of the Triassic. However, compared

to body fossils, especially vertebrate bones, they provide lower temporal resolution. Footprints are the result of the interaction between animals and different substrates. Their shape can thus vary immensely, even if left by a single individual. This has produced a complicated and confusing ichnotaxonomy that reflects the subjective ichnotaxonomic evaluation of footprints by different researchers. Many of the purported ‘diagnostic’ features of tetrapod ichnospecies, but also of some ichnogenera, may be related to extramorphological (substrate-controlled) phenomena. Anatomical signals, such as the number of digits, digit proportions, or the trackway pattern allow discrimination of a number of characteristic form groups and ichnotaxa. In the Triassic, archosaur footprints show a distinct stratigraphic distribution pattern (limited temporal ranges) that can be ascribed to different evolutionary developments of the locomotor apparatus. Several functional-evolutionary sequences and characteristic assemblages can be recognized in the Triassic. They are the basis for the definition of five biochronological units based on tetrapod footprints (Lucas 2003, 2007; Hunt & Lucas 2007b; Klein & Haubold 2007). They can be used to determine the ages of sequences that lack body fossils or can be further used to test and refine Triassic correlations. The authors thank Marco Avanzini, Gerard Gierlinski, Claudia Marsicano and Andrew R. C. Milner for their reviews and critical comments that helped to improve the manuscript. The use of some photos was possible due to the kind permission of Sebastian Voigt, Marco Avanzini, Georges Gand, Henk W. Oosterink and Friedrich Leitz. For support and hospitality during studies in the collections of the University of California Museum of Paleontology, Berkeley, we like to thank Patricia A. Holroyd, Mark B. Goodwin, Randall B. Irmis, Kevin Padian and Mathew J. Wedel.

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W EISS , W. 1934. Eine Fa¨hrtenschicht im Mittelfra¨nkischen Blasensandstein. Jahresberichte und Mitteilungen des Oberrheinischen Geologischen Vereins, N. F., 23, 5– 11. W EISS , W. 1976. Ein Reptilfa¨hrten-Typ aus dem BenkerSandstein und untersten Blasensandstein des Keupers um Bayreuth. Geologische Bla¨tter fu¨r Nordost-Bayern, 26, 1 –7. W ING , S. L. & S UES , H.-D. 1992. Mesozoic and early Cenozoic terrestrial ecosystems. In: B EHRENSMEYER , A. K., D AMUTH , J. D., D I M ICHELE , W. A., P OTTS , R., S UES , H.-D. & W ING , S. L. (eds) Terrestrial Ecosystems Through Time. University of Chicago Press, Chicago, 327 –416.

The Triassic timescale based on nonmarine tetrapod biostratigraphy and biochronology SPENCER G. LUCAS New Mexico Museum of Natural History and Science, 1801 Mountain Road NW, Albuquerque, NM 87104-1375 USA (e-mail: [email protected]) Abstract: The Triassic timescale based on nonmarine tetrapod biostratigraphy and biochronology divides Triassic time into eight land-vertebrate faunachrons (LVFs) with boundaries defined by the first appearance datums (FADs) of tetrapod genera or, in two cases, the FADs of a tetrapod species. Definition and characterization of these LVFs is updated here as follows: the beginning of the Lootsbergian LVF ¼ FAD of Lystrosaurus; the beginning of the Nonesian ¼ FAD Cynognathus; the beginning of the Perovkan LVF ¼ FAD Eocyclotosaurus; the beginning of the Berdyankian LVF ¼ FAD Mastodonsaurus giganteus; the beginning of the Otischalkian LVF ¼ FAD Parasuchus; the beginning of the Adamanian LVF ¼ FAD Rutiodon; the beginning of the Revueltian LVF ¼ FAD Typothorax coccinarum; and the beginning of the Apachean LVF ¼ FAD Redondasaurus. The end of the Apachean (¼ beginning of the Wasonian LVF, near the beginning of the Jurassic) is the FAD of the crocodylomorph Protosuchus. The Early Triassic tetrapod LVFs, Lootsbergian and Nonesian, have characteristic tetrapod assemblages in the Karoo basin of South Africa, the Lystrosaurus assemblage zone and the lower two-thirds of the Cynognathus assemblage zone, respectively. The Middle Triassic LVFs, Perovkan and Berdyankian, have characteristic assemblages from the Russian Ural foreland basin, the tetrapod assemblages of the Donguz and the Bukobay svitas, respectively. The Late Triassic LVFs, Otischalkian, Adamanian, Revueltian and Apachean, have characteristic assemblages in the Chinle basin of the western USA, the tetrapod assemblages of the Colorado City Formation of Texas, Blue Mesa Member of the Petrified Forest Formation in Arizona, and Bull Canyon and Redonda formations in New Mexico. Since the Triassic LVFs were introduced, several subdivisions have been proposed: Lootsbergian can be divided into three sub-LVFs, Nonesian into two, Adamanian into two and Revueltian into three. However, successful inter-regional correlation of most of these sub-LVFs remains to be demonstrated. Occasional records of nonmarine Triassic tetrapods in marine strata, palynostratigraphy, conchostracan biostratigraphy, magnetostratigraphy and radioisotopic ages provide some basis for correlation of the LVFs to the standard global chronostratigraphic scale. These data indicate that Lootsbergian ¼ uppermost Changshingian, Induan and possibly earliest Olenekian; Nonesian ¼ much of the Olenekian; Perovkan ¼ most of the Anisian; Berdyankian ¼ latest Anisian? and Ladinian; Otischalkian ¼ early to late Carnian; Adamanian ¼ most of the late Carnian; Revueltian ¼ early– middle Norian; and Apachean ¼ late Norian– Rhaetian. The Triassic timescale based on tetrapod biostratigraphy and biochronology remains a robust tool for the correlation of nonmarine Triassic tetrapod assemblages independent of the marine timescale.

Triassic tetrapod (amphibian and reptile) fossils have long been used in biostratigraphy, a tradition extending back to at least the 1870s. Lucas (1990) advocated developing a global Triassic timescale based on tetrapod evolution, and subsequently Lucas (1998a) presented a comprehensive global Triassic tetrapod biochronology (Fig. 1). This biochronological timescale divides the Triassic into eight time intervals (land-vertebrate faunachrons, LVFs) based on successive changes in faunal composition driven by tetrapod evolution. This model has been tested and refined for more than a decade. Here, I present the current status of the Triassic tetrapod-based timescale, incorporating new data, analyses and modifications published since 1998.

In this paper: FAD ¼ first appearance datum; HO ¼ highest occurrence; LO ¼ lowest occurrence; LMA ¼ land–mammal ‘age’; LVA ¼ land-vertebrate ‘age’; LVF ¼ land–vertebrate faunachron; and SGCS ¼ standard global chronostratigraphic scale (the marine timescale).

Previous studies Although tetrapods have been used to correlate nonmarine Triassic strata since the 1800s, before the 1990s few attempts were made to establish a global tetrapod biostratigraphy or biochronology of the Triassic (Fig. 2). In the late 1800s, some workers did use tetrapod fossils to correlate

From: LUCAS , S. G. (ed.) The Triassic Timescale. Geological Society, London, Special Publications, 334, 447– 500. DOI: 10.1144/SP334.15 0305-8719/10/$15.00 # The Geological Society of London 2010.

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J

Hettangian

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Rhaetian

Wassonian Protosuchus Apachean Redondasaurus

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Typothorax coccinarum Late

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Eocyclotosaurus 250 252

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Cynognathus Lystrosaurus

Fig. 1. The Triassic timescale based on tetrapod biostratigraphy and biochronology. Restoration of Typothorax by Matt Celeskey.

nonmarine Triassic strata on a broad scale, for example, Cope (1875) who correlated part of the German Keuper to the Upper Triassic strata of the American Southwest based on shared taxa of fossil reptiles such as the phytosaur ‘Belodon’. Broom (1906, 1907, 1909) introduced the earliest, and perhaps the most influential, Triassic

tetrapod biostratigraphy, for the Lower Triassic of the Karoo basin in South Africa. He identified three successive biostratigraphic intervals, the Lystrosaurus, Procolophon and Cynognathus ‘beds’. Watson (1914a, b) later termed these ‘zones’ and, since Kitching (1970), the Lystrosaurus and Procolophon zones have been combined into a

Period

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Ordosian lvf Fuguan lvf

Nonesian lvf

Jimsarian lvf

Lootsbergian lvf

Fig. 2. Previous tetrapod-based subdivisions of Triassic time.

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single, Lystrosaurus zone (e.g. Rubidge et al. 1995; Botha & Smith 2007). Recognition elsewhere of the Lystrosaurus and/or Cynognathus ‘beds’ or ‘zones’ has long been possible in Antarctica, South America, India, China and/or Russia because some Early Triassic tetrapod taxa are virtually cosmopolitan, especially the genera Lystrosaurus and Cynognathus (Lucas 1998a). Romer (1975; also see Cox 1973) presented the first global Triassic tetrapod biochronology, by identifying three successive Triassic land– vertebrate ‘faunas’: A, Early Triassic; B, Middle Triassic; and C, Late Triassic (Fig. 2). Cosgriff (1984) divided Romer’s division A into A1 (¼ Lystrosaurus biochron) and A2 (¼ Cynognathus biochron). Ochev & Shishkin (1989; also see Anderson & Cruickshank 1978) recognized the same intervals as Romer, but chose to name them the: A, proterosuchian epoch; B, kannemeyerioidean epoch; C, and dinosaurian epoch. Cooper (1982) proposed a more detailed global tetrapod biostratigraphy of the Triassic than did Romer and other workers of the 1970s and 1980s (Fig. 2). In this, he recognized a succession of six Triassic zones based largely on a perceived stratigraphic succession of dicynodonts (Lucas & Wild 1995 later presented a revized Triassic dicynodont biozonation). Subsequent workers have not adopted Cooper’s zonation. Indeed, prior to Lucas (1998a), the concept of a global Triassic tetrapod biostratigraphy and biochronology had not progressed beyond Romer (1975). Tetrapod-based subdivisions of Triassic time have been proposed as local, provincial biochronologies for Argentina, North America and China. Bonaparte (1966, 1967, 1982) introduced a set of ‘provincial ages’ for the Triassic of Argentina, but he never defined these terms (Fig. 2). However, since then Lucas & Harris (1996) have defined the Chanarian as a LVF, and Langer (2005b) has defined the Ischigualastian as a LVF. Lucas (1993a) proposed a succession of LVFs for the Chinese Early– Middle Triassic tetrapod record. At about the same time, Lucas & Hunt (1993a) proposed Late Triassic LVFs based on the Chinle Group tetrapod record from the western United States, and Huber et al. (1993b) proposed Middle–Late Triassic LVFs based on the Newark Supergroup record of eastern North America (Fig. 2). Lucas et al. (1997a) since then have presented revized definitions of some of the Late Triassic LVFs proposed by Lucas & Hunt (1993a). Lucas & Huber (2003) reviewed global Late Triassic tetrapod biochronology and demonstrated the broad applicability of the LVFs proposed by Lucas and Hunt (1993a; also see Lucas 1997a). Lucas et al. (2007e) reviewed the status of the Triassic timescale based on patterns of tetrapod evolution

and made some necessary modifications that are incorporated and elaborated upon here.

Vertebrate biostratigraphy and biochronology The term LMA has long referred to intervals of geological (mostly Cenozoic) time characterized by distinctive mammalian fossil assemblages. LMAs have been defined to encompass Cenozoic time intervals on most of the world’s continents (Savage & Russell 1983), and for the Late Cretaceous of western North America (Cifelli et al. 2004). However, more broadly-based LVA or LVF have been introduced for parts of the Mesozoic record of Asia, South America and North America (Lucas 1997b, 2008). Thus, LVAs or LVFs have been proposed for the Triassic and Jurassic of China (Lucas 1993a, 1996); the Triassic of Argentina (Bonaparte 1966); the Late Triassic of western North America (Lucas & Hunt 1993a); the Middle Triassic –Early Jurassic of eastern North America (Huber et al. 1993a; Lucas & Huber 2003; Lucas & Tanner 2007a, b); the Late Jurassic–Early Cretaceous of western North America (Lucas 1993e); the Late Cretaceous of western North America (Russell 1964, 1975; Sullivan & Lucas 2003, 2006); the Late JurassicCretaceous of Mongolia and China (Jerzykiewicz & Russell 1991; Lucas & Estep 1998; Lucas 2006a); and the Cretaceous of Argentina (Leanza et al. 2004). Russell (1993) proposed marine vertebrate ages for the Cretaceous of western North America. Mammals are not the only tetrapods that can be used to recognize intervals of geologic time. In the Mesozoic, especially prior to the Late Cretaceous, when mammal fossils are very rare, nonmammalian tetrapods can be biochronologically useful. For this reason, some workers use the term LVA. Because LMAs and LVAs are not formal ages in stratigraphy, Lucas (1993a) introduced the term faunachron (essentially the same concept as Dunbar & Rodgers’ [1957] ‘faunichron’) to refer to the time interval that is equivalent to the duration of a ‘fauna’. I, thus, use the more precise term LVF instead of LMA or LVA. LVFs are biochronological units, and I define their beginnings by biochronological events. Each LVF begins with the FAD of a tetrapod index taxon, usually a genus, though species are used if they provide greater biostratigraphic resolution. In so doing, the end of an LVF is defined by the beginning of the succeeding LVF, which is the FAD of another tetrapod index taxon. This is a precise way to define LVF boundaries. LVFs thus are interval biochrons. A distinctive assemblage of vertebrate fossils characterizes each LVF. The name of the LVF is a

TRIASSIC TETRAPODS

geographical name taken from the place where (or very close to where) the characteristic example of the vertebrate fossil assemblage was collected. Many LMA and LVA names have been taken from the rock formation in which the fossils are found, and the rock formation name is based on a place name. However, using the rock formation name may cause confusion because it can imply that the LMA or LVA refers to the entire duration of deposition of the formation and not just to the duration of interval in which the vertebrate fossil assemblage is found, which is often much shorter. It is less confusing to choose another place name for the LMA or LVA. For example, the Late Triassic Ischigualastian LVA of Argentina (Bonaparte 1966) was named for the Ischigualasto Formation, but the Ischigualastian LVF vertebrates do not occur throughout the Ischigualasto Formation, which is potentially confusing. In contrast, the Late Triassic Adamanian LVF of western North America (Lucas & Hunt 1993a) is named after Adamana, where the fossils occur, not after the Blue Mesa Member of the Petrified Forest Formation, which contains the characteristic fossil assemblage. This prevents confusion between the concept of a formation and the concept of a LVF. The characteristic tetrapod assemblage is the primary basis for characterization of the LVF. Index fossils identified here meet the criteria of true index fossils (temporally restricted, common, widespread, easily identified) and do not include endemic or rare taxa that happen to be restricted to a LVF, usually as single records. Principal

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correlatives of the characteristic tetrapod assemblage of each LVF are listed in this article. These are tetrapod assemblages that are reasonably well studied, diverse and unambiguously correlated. Although I make a strong effort here to correlate the LVFs to the SGCS, the tetrapod biochronology of the Triassic is a timescale independent of the SGCS. It is also important to keep in mind that, although global LVF’s could not be defined today due to the wide separation of most of the continents, in the Triassic Pangaean world it was possible for at least some of the land vertebrates to spread across most of the world’s land area. Some degree of endemism is apparent, but it was not so great as to prevent definition of global or near-global faunachrons.

Triassic land – vertebrate faunachrons Introduction The Triassic tetrapod timescale is based on tetrapod assemblages from the Karoo basin in South Africa (Early Triassic: Lootsbergian–Nonesian), the Ural foreland basinin Russia (MiddleTriassic:Perovkan – Berdyankian) and the Chinle basin of the western USA (Upper Triassic: Otsichalkian–Apachean) (Fig. 3). The Karoo basin contains the tetrapod assemblages characteristic of the Lootsbergian and Nonesian LVFs. These assemblages are stratigraphically superposed and are thus demonstrably time successive; they are the classic Lystrosaurus assemblage zone and most of the Cynognathus

Ural basin

Chinle basin

Karoo basin

Fig. 3. Map of Triassic Pangaea showing the three areas that provide the fossils and strata that form the standards for the Triassic tetrapod timescale: Karoo basin, South Africa (Lootsbergian and Nonesian), Russian Urals basin (Perovkan and Berdyankian) and Chinle basin (Otischalkian, Adamanian, Revueltian and Apachean). Base map drawn by Matt Celeskey.

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assemblage zone (e.g. Rubidge et al. 1995; Groenewald & Kitching 1995; Kitching 1995; Hancox & Rubidge 1997; Hancox 2000; Smith & Botha 2005; and Botha & Smith 2007 provide an overview). These assemblages include amphibians, parareptiles, dicynodonts and cynodonts particularly useful for broad correlation. The South African Triassic tetrapod record contains a long hiatus between the uppermost strata of the Lower Triassic Cynognathus assemblage zone and southern African rocks that contain tetrapods of certain Late Triassic age (notably the lower Elliot Formation: Lucas & Hancox 2001). This forces the tetrapod biochronological standards for Middle Triassic time to be moved elsewhere. For this part of the standards, Lucas (1998a) used two superposed tetrapod assemblages from the Russian Ural foreland basin (e.g. Shishkin et al. 1995b, 2000a, b; Ivakhnenko et al. 1997; Novikov et al. 2000; Battail & Surkov 2000; Gower & Sennikov 2000; Spencer & Benton 2000; Ivakhnenko 2008a, b, c; Sennikov 2008; Tatarinov 2008) as the basis for the Middle Triassic Perovkan and Berdyankian LVFs. The presence of some temporal overlap between the top of the South African section (upper Cynognathus Zone) and the Urals foreland basin section makes correlation between these sections considerably easier. The Russian assemblages yield amphibians, archosaurs and dicynodonts of value for broad correlation. No Upper Triassic tetrapod assemblages are known from the Russian Ural foreland basin (e.g. Shishkin et al. 2000b), so the tetrapod biochronology standard for Late Triassic time again must be moved elsewhere. The Chinle Group strata of the American Southwest provide the best studied and most complete record for defining the Late Triassic LVFs: Otischalkian, Adamanian, Revueltian and Apachean. Of great importance, tetrapod assemblages from Texas (Otischalkian characteristic assemblage), Arizona (Adamanian characteristic assemblage) and New Mexico (Revueltian and Apachean characteristic assemblages) are stratigraphically superposed and thus are time successive (e.g. Lucas 1993c, 1997a; Lucas et al. 2001; Heckert & Lucas 2002a, b, 2003; Heckert 2004; Heckert et al. 2005a, b; Parker et al. 2006). The Chinle assemblages yield phytosaurs, aetosaurs and metoposaurs useful for broad correlation, and a burgeoning microvertebrate biostratigraphy also supports ths macrovertebrate-based correlation (Heckert 2004; Heckert & Lucas 2006).

Lootsbergian LVF Definition. Lucas (1998a) introduced the term Lootsbergian LVF for the time between the FAD of the dicynodont Lystrosaurus and the FAD of

the cynodont Cynognathus (Fig. 1). Its characteristic tetrapod assemblage is the Lystrosaurus Assemblage Zone found in the Balfour (Palingkloof Member), Katberg and Burgersdorp (lower part) formations of the Karoo basin of South Africa (e.g. Groenewald & Kitching 1995; Damiani et al. 2003; Smith & Botha 2005; Botha & Smith 2006, 2007). This assemblage zone has a type locality designated by Groenewald & Kitching (1995) around Lootsberg Pass. Lootsbergian time begins with the FAD of Lystrosaurus, which is the end of the Late Permian Platbergian LVF of Lucas (2005, 2006b). The end of the Lootsbergian is equivalent to the beginning of the Nonesian LVF, which is defined by the FAD of Cynognathus. Broom (1906) introduced two successive zones – Lystrosaurus and Procolophon – that Kitching (1970, 1977) later combined into a single, Lystrosaurus Zone. Keyser (1979) referred to this same zone as the Lystrosaurus-Thrinaxodon Assemblage Zone. The original name Lystrosaurus Zone (or Assemblage Zone) continues to be used (e.g. Groenewald & Kitching 1995; Lucas 1998a; Damiani et al. 2001; Botha & Smith 2006, 2007; Smith & Botha 2005). Characteristic tetrapod fossil assemblage. The characteristic tetrapod fossil assemblage of the Lootsbergian LVF is the Lystrosaurus Assemblage Zone of the Karoo basin, South Africa. It consists of amphibians, parareptiles, prolacertiforms, archosaurs, dicynodonts, therocephalians and cynodonts. Kitching (1977) reviewed the Lystrosaurus Assemblage Zone localities, Groenewald & Kitching (1995) provided a synopsis of the stratigraphic ranges of the genera, and Botha & Smith (2006, fig. 7) have presented the most recent data. The Lystrosaurus Assemblage Zone has long provided a standard for correlation of the oldest Triassic tetrapod assemblages, so it logically serves as the basis for the oldest Triassic LVF (though it encompasses the Permo-Triassic boundary and includes some uppermost Permian strata, see below). Index fossils. The following tetrapod genera are restricted to Lootsbergian time and are widespread and/or common enough to be useful as index fossils (Fig. 4): the amphibians Wetlugasaurus, Tupilakosaurus, Luzocephalus, and Lydekkerina; the parareptile Procolophon; the prolacertiform Prolacerta; the archosaur Proterosuchus (¼ Chasmatosaurus); the dicynodont Lystrosaurus; and the cynodonts Scaloposaurus and Thrinaxodon. Principal correlatives. Recognition of and correlation within the Lootsbergian is one of the most stable parts of the Triassic tetrapod timescale. Thus, the terms Lystrosaurus zone, beds or fauna

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taxa

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Nonesian

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Berdyankian

amphibians: Eocyclotosaurus Eryosuchus Luzocephalus Lydekkerina Mastodonsaurus Odenwaldia Paracyclotosaurus Parotosuchus Trematosaurus Trematosuchus Tupilakosaurus Wetlugasaurus parareptile: Procolophon prolacertiform: Prolacerta archosaurs: Arizonasaurus Erythrosuchus dicynodonts: Kannemeyeria Lystrosaurus Parakannemeyeria Shansiodon Sinokannemeyeria Stahleckeria cynodonts: Cynognathus Diademodon Massetognathus Scalenodon Scaloposaurus Thrinaxodon Trirachodon Fig. 4. Temporal ranges of selected genera of Early and Middle Triassic tetrapods.

have long been applied to a wide geographical range of strata/fossils of Lootsbergian age. Most significant correlatives are the vertebrate fossil assemblages of the: Wordy Creek Formation, eastern Greenland; Vokhmian, Sludkian and Ustmylian horizons of the Vetluga Series, Russian Urals; upper Guodikeng and lower Jiucaiyuan

formations, Junggur basin, China; Heshanggou Formation, Ordos basin, China; Panchet Formation, India; Sanga do Cabral Formation, Parana´ basin, Brazil; Rewan Formation, SE Galilee basin, Australia; Arcadia Formation, SW Bowen basin, Australia; and lower part of Fremouw Formation, Antarctica. Note that the alleged Lystrosaurus

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record from Laos (Repelin 1923; Piveteau 1938) has been re-identified as the Late Permian dicynodont Dicynodon (Battail et al. 1995; Battail 1997). The Wordy Creek Formation in eastern Greenland yields the amphibians Luzocephalus, Wetlugasaurus and Tupilakosaurus (Sa¨ve-So¨derbergh 1935; Nielsen 1954) and thus is of Lootsbergian age. These strata also yield Induan ammonites, and are key to correlation of the Lootsbergian to part of the Induan (see below). In the Russian Urals, the Lootsbergian interval is equivalent to Zone V of Efremov (1937, 1952), which has most recently been called the Vokhmian, Sludkian and Ustmylian horizons of the Vetlugan Series (Superhorizon) (Ivakhnenko et al. 1997; Shishkin et al. 2000b). Tetrapod taxa include anthracosaurs, the temnospondyls Luzocephalus, Benthosuchus, Wetlugasaurus and Tupilakosaurus, procolophonids, a prolacertiform, the proterosuchid Chasmatosuchus and other (mostly fragmentary) archosaurs and the dicynodont Lystrosaurus (Shishkin et al. 1995b; Ivakhnenko et al. 1997; Battail & Surkov 2000; Gower & Sennikov 2000; Novikov et al. 2000; Shishkin et al. 2000a, b; Spencer & Benton 2000). In northwestern China, land-vertebrates of Lootsbergian age come from the upper part of the Guodikeng Formation and the lowermost Jiucaiyuan Formation (both in the Cangfanggou Group) near Jimsar NE of Urumqi in western Xinjiang (e.g. Cheng 1981; Metcalfe et al. 2009). These vertebrates are the ‘Lystrosaurus fauna’ of northwestern China of some workers (e.g. Sun 1972), and they provided the basis for the Jimsarian LVF of Lucas (1993a). Taxa present are a prolacertid, a ?procolophonid, the proterosuchian Proterosuchus (¼ Chasmatosaurus), a regisaurid therocephalian and the dicynodont Lystrosaurus, of which seven species have been named, most of which are invalid (Colbert 1974; Colbert & Kitching 1977; Lucas 2001). In the Ordos basin of north–central China, near Fugu, Shanxi, the upper part of the Heshanggou Formation yields a vertebrate fauna that was the basis of the Fuguan LVF of Lucas (1993a). Taxa present are indeterminate capitosauroids, procolophonids, an erythrosuchid and an ordosiid therocephalian; based primarily on the procolophonids, these are of likely Lootsbergian age. In India, the Panchet Formation along the Damodar River northwest of Calcutta has produced a Lootsbergian vertebrate assemblage that includes a lydekkerinid, ?benthosuchid, ?capitosaurids, an indobrachyopid, trematosaurids, a procolophonid, the proterosuchian Proterosuchus and Lystrosaurus (Lydekker 1882; Sahni & Huene 1958; Tripathi 1961, 1969; Tripathi & Satsangi 1963; Hughes 1963; Ray 2005).

In southern Brazil, the Sanga do Cabral Formation in the Parana´ basin yields a rhytidosteid amphibian, indeterminate temnospondyls, Procolophon, ?thrinaxodontids and ?Lystrosaurus (e.g. Barbarena et al. 1985; Lucas 2002; Abdala et al. 2002; Cisneros 2008a, b; Cisneros & Schultz 2002; Dias-da-Silva et al. 2005, 2006a, b; Dias-da-Silva & Marsicano 2006; Dias-da-Silva & Schultz 2008). A putative Permian tetrapod record from the Buena Vista Formation of Uruguay (Pin˜eiro et al. 2003, 2004, 2007) is more likely correlative to the Lootsbergian Sanga do Cabral assemblage (Dias-da-Silva et al. 2006b). In eastern Australia, the Arcadia Formation (SW Bowen basin) and the Rewan Formation (SE Galilee basin) yield small assemblages of tetrapods of Lootsbergian age. The Arcadia Formation assemblage encompasses a diversity of mostly endemic amphibians, including fragmentary lydekkerinids, a primitive procolophonid, a possible Prolacerta, an archosaur similar to Proterosuchus and ?Lystrosaurus (e.g. Bartholomai 1979; King 1983; Thulborn 1983; Warren 1991; Damiani 2001; Warren et al. 2006). In the SE Galilee basin, the occurrence of Lydekkerina in the Rewan Formation supports a Lootsbergian age assignment (Warren et al. 2006). Southwest of the Transantarctic Mountains in southern Antarctica, the lower part of the Fremouw Formation yields a vertebrate fossil assemblage of Lootsbergian age that includes temnospondyls, a rhytidosteid, the procolophonid Procolophon, the prolacertiform Prolacerta, a proterosuchid or erythrosuchid, a rauisuchian, the dicynodonts Myosaurus and Lystrosaurus, the cynodont Thrinaxodon and scaloposaurs (e.g. Colbert 1972, 1991; Hammer 1990; Collinson et al. 2006). This Lootsbergian assemblage has been referred to as the lower Fremouw fauna or lower tetrapod fauna of the Fremouw Formation (Colbert 1972, 1991). Comments. Most Lootsbergian vertebrate fossil assemblages are readily recognized by the presence of Lystrosaurus. Procolophon and Proterosuchus are also important to the correlation of Lootsbergian tetrapod assemblages. However, temnospondyldominated assemblages occur that lack Lystrosaurus and thus are more difficult to correlate. I have used the temporal overlap of Lystrosaurus and the amphibians Tupilakosaurus and Luzocephalus in Russian strata as the primary basis for equating Lootsbergian dicynodont-dominated assemblages with temnospondyl-dominated assemblages. Cosgriff (1984) assigned several temnospondyldominated assemblages to his A1 ‘horizon’ (¼ Lootsbergian), even though these lack any index taxa of the Lootsbergian: the Knocklofty Sandstone/Shale in SE Tasmania (Cosgriff 1974),

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Perovkan

Nonesian

Lootsbergian B

C

A

B

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Lystrosaurus with Dicynodon

Lystrosaurus without Dicynodon and Procolophon

Lystrosaurus with Procolophon

Cynognathus without Kannemeyeria

Cynognathus with Kannemeyeria

Angonisaurus and Shansiodon

Definition. The term Nonesian LVF refers to the time between the FAD of the cynodont Cynognathus and the FAD of the amphibian Eocyclotosaurus. The characteristic tetrapod assemblage is found in the lower two-thirds of the Cynognathus Assemblage Zone, which is from the upper two-thirds of the Burgersdorp Formation in the Karoo basin of South Africa (e.g. Kitching 1995; Hancox et al. 1995; Hancox 2000). The type section of the Cynognathus Assemblage Zone encompasses Nonesi’s Nek, from which the name Nonesian is derived (Kitching 1995). Nonesian time begins with the FAD of

A

Nonesian LVF

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P

the Sticky Keep Formation in Svalbard (Wiman 1910, 1915; Nilsson 1942, 1943; Cox & Smith 1973), the upper Andavakoera Formation (Middle Sakamena Group or Formation) in NW Madagascar (Lehman 1961, 1966; Steyer 2002; Maganuco & Pasini 2009) and the Arcadia Formation of southern Queensland (Warren 1991). Except for the Arcadia Formation, I assign these assemblages a Nonesian age (see below). Lootsbergian time encompasses both the ‘Lystrosaurus zone’ and ‘Procolophon zone’ of classic usage (e.g. Broom 1906). Thus, two distinct tetrapod assemblages (at least in the Karoo basin) can be recognized within the Lootsbergian, simply based on the stratigraphic distribution of Procolophon. According to Botha & Smith (2007), all records of Lystrosaurus maccaigi in the Karoo basin are Permian (they co-occur with the Permian dicynodont Dicynodon), whereas L. curvatus straddles the Permo-Triassic boundary, and records of L. murrayi and L. declivus are Triassic. This provides a basis for a threefold subdivision of the Lootsbergian (Fig. 5): (1) Lootsbergian A is the time of overlap of Dicynodon and Lystrosaurus; (2) Lootsbergian B is the succeeding interval with Lystrosaurus without Procolophon; and (3) Lootsbergian C is the temporal overlap of Lystrosaurus and Procolophon. These subdivisions have some value outside of the Karoo basin. For example, in the Guodikeng Formation in the Junggur basin of northwestern China, there is a stratigraphic overlap of Lystrosaurus and Dicynodon (Lootsbergian A) followed by an interval of Lystrosaurus without Procolophon (Lootsbergian B) (Cheng 1981; Metcalfe et al. 2009). Indeed, in northwestern China, the co-occurrence of Lystrosaurus and Dicynodon at Dalongkou was first assigned to the upper Changhsingian Falsisca postera conchostracan zone and uppermost part of the F. eotriassica conchostracan zone by Kozur (1998a, b) (see also Kozur & Weems 2010). Therefore, a formal subdivision of the Lootsbergian into sub-LVFs has merit and should provide more precise correlation within the Lootsbergian interval.

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Fig. 5. Subdivisions of the Lootsbergian and Nonesian LVFs (based primarily on Hancox 2000). Restoration of Lystrosaurus by Matt Celeskey.

Cynognathus, which is the end of the Lootsbergian LVF. The end of the Nonesian is the beginning of the Perovkan LVF, which is defined by the FAD of Eocyclotosaurus. Broom (1906, 1907) coined the name Cynognathus ‘beds’, which was later transmuted to ‘zone’ by other workers (Watson 1914a, b; Kitching 1970,

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1977). Keyser & Smith (1978) renamed it the Kannemeyeria Assemblage Zone, and Keyser (1979) termed it the Kannemeyeria– Diademodon Assemblage Zone. Kitching (1984) called it the Cynognathus– Diademodon Assemblage Zone. The term Cynognathus Assemblage Zone has been used most recently (e.g. Kitching 1995; Rubidge et al. 1995; Lucas 1998a; Hancox 2000). Characteristic tetrapod assemblage. The characteristic assemblage of the Nonesian LVF occurs in subzones A and B of the Cynognathus Assemblage Zone of the Karoo basin (Hancox 2000) (Fig. 5). The tetrapod taxa present are amphibians, including Parotosuchus, Wellesaurus and Trematosuchus, captorhinids, a ?sphenodontid (or ?procolophonid), rhynchosaurs, the archosaurs Erythrosuchus and Euparkeria, the dicynodonts Kannemeyeria and Kombuisia, therocephalians and cynodonts, including Cynognathus, Diademodon and Trirachodon (e.g. Kitching 1977, 1995; Hancox & Rubidge 1994; Hancox et al. 1995; Shishkin et al. 1995a; Damiani 2001; Damiani & Rubidge 2003; Abdala et al. 2005). Index fossils. The following tetrapod genera are restricted to Nonesian time and are widespread and/or common enough to be considered index fossils (Fig. 4): the amphibians Parotosuchus, Odenwaldia, Wellesaurus, Trematosaurus and Trematosuchus and the cynodont Trirachodon. The LOs of the the archosaur Erythrosuchus, the cynodonts Cynognathus and Diademodon and of the dicynodont Kannemeyeria are in the Nonesian. The species K. simocephalus is restricted to Nonesian time, but the species K. cristarhynchus is younger, of Perovkan age. Principal correlatives. Principal correlatives of the type Cynognathus Assemblage Zone are: Wupatki and Torrey formations of the Moenkopi Group/ Formation, Utah/Arizona, USA; Sticky Keep Formation of Svalbard, Arctic Norway; Middle Buntsandstein (upper Volpriehausen, Hardegsen and Solling formations), Germany; Petropavlovsk Formation (Yarenskiy horizon) in the Russian Urals; lower part of Ermaying Formation, Ordos basin, China; Puesto Viejo and Rio Mendoza formations, Argentina; base of the Lower Sandstone of the Zarzaitine Series in Algeria; lower N’tawere Formation, Zambia; K7 horizon of the Kingori Sandstone, Tanzania; and upper Fremouw Formation, Antarctica. The Torrey Formation of the Moenkopi Group in Utah, USA, has yielded a skull of Parotosuchus (Lucas & Schoch 2002). Specimens of Wellesaurus as well as an Odenwaldia-like form are from the Wupatki Member of the Moenkopi Formation in

Arizona (Damiani 2001; Lucas & Schoch 2002; Heckert et al. 2005a; Nesbitt 2005). These records of Nonesian index taxa are of late Olenekian age (see below). In the Germanic basin, the Middle Buntsandstein (upper Volpriehausen, Hardegsen and Solling formations) yields fossils of Parotosuchus, Oldenwaldia and Trematosaurus, indicative of a Nonesian age (e.g. Schroeder 1913; Werneburg 1993; Lucas 1999; Schoch & Werneburg 1999; Lucas & Schoch 2002; Schoch 2008). Specifically, Odenwaldia occurs only in the Solling Formation, and Trematosaurus is common in the Hardegsen Formation and present in the Solling Formation. Parotosuchus is known from the Hardegsen and the Solling formations. One specimen of Parotosuchus (the holotype of P. helgolandicus) is known from the uppermost Volpriehausen Formation, from the upper Gerviellia beds assigned by Kozur & Bachmann (2008), based on conchostracans, to the Spathian. Temnospondyls of the Sticky Keep Formaton in Svalbard co-occur with early Olenekian (Smithian) ammonites (Buchanen et al. 1965; Tozer 1967). The temnospondyls are: Sasenisaurus, Peltostega, Aphanerama (¼ Lonchorhynchus), Lyrocephaliscus, Teretrema and Boreaosaurus (Wiman 1910, 1915, 1916; Nilsson 1942, 1943; Cox & Smith 1973). Such an acme in trematosaur diversity may characterize the Nonesian. I assign a Nonesian age to the Sticky Keep tetrapods based mostly on the marine evidence that they are Olenekian and that the Nonesian is equivalent to at least part of the Olenekian (see below). The Petropavlovsk svita in the Russian Urals (Yarenskiy horizon) yields anthracosaurs, temnospondyls (including Parotosuchus), procolophonids, a prolacertid, and various archosaurs, including erythrosuchids and rauisuchids (Shishkin et al. 1995b, 2000a, b; Ivakhnenko et al. 1997; Battail & Surkov 2000; Gower & Sennikov 2000; Novikov et al. 2000; Spencer & Benton 2000). The Parotosuchus record is the primary basis for a Nonesian age assignment. In China, the lower Ermaying Formation in the Ordos basin produces a vertebrate fauna upon which Lucas (1993a) based the Ordosian LVF. Taxa present are a procolophonid, a proterosuchian, euparkeriids, a therocephalian and the dicynodonts Parakannemeyeria and Kannemeyeria (¼ Shaanbeikannemeyeria) (Lucas 2001). The Kannemeyeria record as well as the overall composition of the assemblage suggest a Nonesian age. In Argentina, the upper part of the Puesto Viejo Formation produces the dicynodont Kannemeyeria, a traversodontid and Cynognathus (Bonaparte 1970, 1978, 1982). The co-occurrence of Cynognathus and Kannemeyeria supports a Nonesian age

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assignment. The correlative fauna from the middle part of the Rio Mendoza Formation (but see Zavattieri & Arcucci 2007 for a different correlation) includes Kannemeyeria (Vinceria andina Bonaparte is not Shansiodon, as Lucas [1993e] suggested, but instead is Kannemeyeria), traversodontids and a galeosaurid. Bonaparte (1981) described dicynodonts and proterosuchian postcrania from the lower part of the Puesto Viejo Formation. He referred to them as the Agua de los Burros local fauna. He assigned the dicynodonts to ‘Vinceria’ (¼ Kannemeyeria) and claimed correlation to the Lystrosaurus Assemblage Zone based on a mean value of K/Ar ages of 232 + 4 Ma from basalts and tuffs that bracket the fossils (Valencio et al. 1975, fig. 2). Given that we now know that the Induan is approximately 251 – 252 Ma (Bachmann & Kozur 2004; Kozur & Weems 2010; Mundil et al. 2010), the Argentinian dates (which are Carnian by current Triassic timescale calibration) do not support Bonaparte’s correlation, nor do the fossils, which instead suggest a Nonesian age. The base of the Lower Sandstone of the Zarzaitine Series in southeastern Algeria yields the amphibians Odenwaldia and ‘Wellesaurus’ (an indeterminate heylerosaurid according to Damiani 2001) as well as a ?brachyopid, trematosaurid and the prolacertiform Jesairosaurus (Lehman 1957, 1971; Welles 1993; Jalil 1990, 1993, 1994, 1997, 1999). The record of Odenwaldia supports a Nonesian age assignment. The lower part of the N’tawere Formation in Zambia produces Diademodon and Kannemeyeria (Crozier 1970). In the Ruhuhu Valley of Tanzania, the K7 horizon of the Kingori Sandstone Formation of Stockley (1932) yields Kannemeyeria (Cruickshank 1986). These are likely (though not definitively) Nonesian records. In Antarctica, the upper part of the Fremouw Formation yields capitosaurid temnospondyls including Parotosuchus, Cynognathus, a diademodontid and a kannemeyeriid (Colbert 1991; Hammer 1988, 1990, 1995; Damiani 2001; Collinson et al. 2006). This has long been regarded as an assemblage of the ‘Cynognathus zone’, and is of Nonesian age. Comments. Most Nonesian vertebrate assemblages have long been recognized by the presence of Cynognathus and/or Diademodon, but these taxa have temporal ranges that extend into the Perovkan. Parotosuchus is a key temnospondyl taxon to correlate many Nonesian assemblages (Damiani 2001). The temporal succession of Kannemeyeria species is important, with K. simocephalus restricted to the Nonesian and K. cristarhynchus a Perovkan taxon. Kitching (1977) reviewed the Cynognathus Assemblage Zone localities, and Kitching (1995)

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and Hancox (2000) provided a synopsis of the stratigraphic ranges of the genera. Watson (1942) and Kitching (1977) subdivided the Cynognathus Assemblage Zone into two subzones. Hancox & Rubidge (1994), Hancox et al. (1995), Shishkin et al. (1995a), Hancox (2000; Hancox et al. 1995, 2000) and Abdala et al. (2005) divided the Cynognathus Assemblage Zone into three stratigraphically discrete assemblages (Fig. 5). These assemblages have been called subzones A, B and C by Hancox et al. (1995), and the upper one is now assigned a Perovkan age (Hancox 2000; Abdala et al. 2005; Lucas et al. 2007e). This means that the South African Nonesian (which encompasses subzones A and B) is divisible into two biochronological units (Hancox 2000). A more important point is that recognizing subzone C as Perovkan means that not all of the classicallyrecognized ‘Cynognathus zone’ is Nonesian. Following Hancox (2000), the Nonesian can be subdivided into older (Nonesian A) and younger (Nonesian B) sub-LVFs (Fig. 5). Nonesian A begins with the FAD of Cynognathus, and Nonesian B begins with the FAD of Kannemeyeria. The FAD of Eocyclotosaurus (beginning of the Perovkan LVF) is the end of Nonesian B. In the Karoo basin, where Eocyclotosaurus is so far unknown, the LO of shansiodont dicynodonts approximates the beginning of Perovkan time (Fig. 5). In Nonesian A time in the Karoo basin, the amphibian Kestrosaurus is common and associated with Trematosuchus as well as theriodonts, Cynognathus, Diademodon, Trirachodon and Bauria. During Nonesian B time, characteristic taxa are Parotosuchus, Kannemeyeria, Cynognathus, Diademodon, Trirachodon, Bauria, Erythrosuchus and Euparkeria. Most of the Nonesian correlative tetrapod assemblages (see above) include Kannemeyeria, so they are of Nonesian B age.

Perovkan LVF Definition. The term Perovkan LVF refers to the time interval between the FADs of the amphibians Eocyclotosaurus and Mastodonsaurus giganteus (Fig. 1). The characteristic tetrapod assemblage is the vertebrate fossil assemblage of the Donguz svita (Eryosuchus fauna) in the Russian Urals (Shishkin et al. 1995b, 2000b; Ivahknenko et al. 1997). Lucas (1998a) termed this the Shansiodon Assemblage Zone, after the distinctive dicynodont Shansiodon (¼ Rhinodicynodon). These fossils are from an approximately 175-m-thick section exposed in the Donguz River drainage near the city of Perovka, from which the name of the LVF is taken (Lucas 1998a, fig. 8). The beginning of the Perovkan is defined by the FAD of the amphibian Eocyclotosaurus. The end of the Perovkan LVF is the beginning of the Berdyankian LVF,

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which is defined by the FAD of the amphibian Mastodonsaurus giganteus. Lucas (1998a) originally defined the beginning of Perovkan time as the FAD of the dicynodont Shansiodon. However, Shishkin (2000) has argued that the type assemblage of the Perovkan LVF is late Anisian, so it is younger than the Eocyclotosaurus assemblage that typically represents the Perovkan in western Europe and North America and is of unambiguous early Anisian age (Lucas & Schoch 2002). A more circumspect reading of the same data (e.g. Ivakhenko et al. 1997) simply regards the Donguz assemblage as Anisian, with no more precise correlation to the SGCS. Lucas (1993d) argued that the LO of the dicynodont Shansiodon is Anisian, and this is why Lucas (1998a) used it to define the beginning of the Perovkan. However, if the LO of Shansiodon is actually younger than the LO of Eocyclotosaurus, then records of Eocyclotosaurus (Upper Buntsandstein in Germany and France, upper Moenkopi Group in USA) are of Nonesian age. The temporal succession of Eocyclotosaurus and Shansiodon is not easily resolved, but Lucas et al. (2007e) noted that the LO of Kannemeyeria in China predates the LO of Shansiodon, and in South Africa the LO of Kannemeyria predates the LO of shansiodonts (Fig. 5), and there is no conclusive evidence that the youngest Nonesian assemblage in South Africa (subzone B of Hancox et al. 1995) is equivalent to the Eocyclotosaurus zone. Lucas et al. (2007e) therefore recognized problems in establishing the temporal succession of Perovkan assemblages, but believe all are broadly Anisian, and some (part of American Moenkopi Group, German Ro¨t Formation) are clearly early Anisian. The easiest way to remove ambiguity here is to redefine the beginning of the Perovkan as the FAD of Eocyclotosaurus, as did Lucas et al. (2007e). Characteristic tetrapod assemblage. Three principal sites in the Donguz svita produce the following taxa: various amphibians, including Eryosuchus, Bukobaja, Plagiosternum and Plagioscutum, a procolophonid, a prolacertid, a proterosuchid, the erythrosuchid Erythrosuchus, rauisuchids, a euparkeriid, the dicynodonts Kannemeyeria (¼ Rhadiodromus, Rabidosaurus, Edaxosaurus, Calleonassus and Rhinocerocephalus) and Shansiodon (¼ Rhinodicynodon), therocephalians, the cynodonts Scalenodon, Antecosuchus and a traversodontid (Shishkin et al. 1995b, 2000a, b; Ivakhnenko et al. 1997; Surkov 1999; Battail & Surkov 2000; Gower & Sennikov 2000; Spencer & Benton 2000; Tverdokhlebov et al. 2002). Index fossils. The following tetrapod genera are common and/or widespread enough to be useful

index taxa of the Perovkan (Fig. 4): the amphibians Eryosuchus, Eocyclotosaurus and Paracyclotosaurus, the archosaur Arizonasaurus, the cynodont Scalenodon and the dicynodonts Shansiodon, Parakannemeyeria and Sinokannemeyeria. Kannemeyria christarhynchus is a Perovkan index fossil, and the HOs of Kannemeyeria, Cynognathus and Diademodon are Perovkan. Principal correlatives. Principal correlatives of the type Perovkan assemblage are from the Holbrook and Anton Chico members of the Moenkopi Formation, Arizona–New Mexico USA; lower part of Wolfville Formation at Lower Economy, Fundy basin, Nova Scotia, Canada; Otter Sandstone of the United Kingdom; Upper Buntsandstein (Ro¨t Formation), Germany-France; lower Kelamayi Formation, Junggur basin, Xinjiang, China; upper Ermaying Formation, Ordos Basin, China; Yerrapalli Formation, India; Lower Zarzaitine Formation, Algeria; upper part of the Burgersdorp Formation in the Karoo basin of South Africa; Omingonde Formation, Namibia; and lower Manda Formation, Tanzania. The Holbrook and Anton Chico members of the Moenkopi Formation, in Arizona–New Mexico, USA, yield the characteristic Perovkan capitosauroid amphibian Eocyclotosaurus, very similar to E. lehmanni from the Upper Buntsandstein (Ro¨t Formation), as well as other capitosaurs, brachyopids, and the ctenosauriscid Arizonasaurus (Lucas & Morales 1985; Lucas & Hunt 1987; Morales 1987; Schoch 2000b; Boy et al. 2001; Lucas & Schoch 2002; Heckert et al. 2005a; Nesbitt 2005). A Shansisuchus-like erythrosuchian from the Anton Chico Member in New Mexico (Lucas et al. 1998b; Nesbitt et al. 2006) is consistent with a Perovkan age assignment. In the Fundy basin of Nova Scotia, Canada, the lower part of the Wolfville Formation (also referred to as the ‘Lower Economy Beds’) yields a small tetrapod assemblage that was the basis of the Economian LVF of Huber et al. (1993b). The presence of a trematosaur (cf. Cosgriffius) and the lepidosaur cf. Tanystropheus suggests a possible Perovkan age (Lucas & Huber 2003). The Otter Sandstone in Devon, United Kingdom, yields the temnospondyl Eocyclotosaurus, the rhynchosaur Fodonyx, the prolacertiform Tanystropheus, a procolophonid, a rauisuchian and a ?ctenosauriscid archosaur (Benton et al. 1994; Hone & Benton 2008). As Milner et al. (1990) stressed, ‘Mastodonsaurus’ lavisi Seeley from the Otter Sandstone is a nomen dubium upon which it is risky to draw stratigraphic conclusions, so I do not consider it a Perovkan record of Mastodonsaurus. Indeed, Damiani (2001) considered the type material of ‘M.’ lavisi to be indeterminate.

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In Germany and France, the Upper Buntsandstein (Ro¨t Formation) yields Eocyclotosaurus (Heyler 1969, 1976; Ortlam 1970; Kamphausen & Morales 1981; Lucas & Schoch 2002) and is of Perovkan age. In the Junggur basin of Xinjiang, China, the lower part of the Kelamayi (¼ Karamay) Formation produces a vertebrate fauna that consists of indeterminate labyrinthodonts (including the holotype of the nomen dubium ‘Parotosaurus’ [¼ Parotosuchus] turfanensis Young: Lucas & Hunt 1993b), a euparkeriid, an erythrosuchid and the dicynodonts Parakannemeyeria and Xiyukannemeyeria (Liu & Li 2003; Liu 2004). The bauriid therapsid Traversodontoides from Jiyuan, Henan may also be of Perovkan age (Cheng 1981; Sun 1989). The upper part of the Ermaying Formation in the Ordos basin in northern China produces what has been called the Perovkan-age ‘Sinokannemeyeria fauna’ or ‘kannemeyeriid fauna’ of China (e.g. Sun 1972; Cheng 1981; Lucas 2001). Lucas (1993a) based the Ningwuan LVF on this assemblage. The vertebrate fossil assemblage includes indeterminate labyrinthodonts, a procolophonid, erythrosuchids, an ?ornithosuchid, a ?euparkeriid, a cynodont, and the dicynodonts Shansiodon, Sinokannemeyeria and Parakannemeyeria (Lucas 2001). In the Pranhita–Godavari Valley of India, the Yerrapalli Formation yields an assemblage of Perovkan age. It includes the amphibian Eryosuchus, the rhynchosaur Mesodapedon, a prolacertid, the archosaur Erythrosuchus, a raisuchid, the dicynodonts Wadiasaurus and Kannmeyeria (¼ Rechnisaurus), and a trirachodontid, (e.g. Roychowdhury 1970a, b; Chatterjee 1980b; Damiani 2001; Sen 2003, 2005; Bandyopadhyay & Sengupta 2006). The Omingonde Formation in Namibia produced a Perovkan-age assemblage that includes an eryopoid temnospondyl, the dicynodonts Kannemeyeria cristarhynchus, Dolichuranus, and Rhopalorhinus, a bauriamorph, and cynodonts, including ?Cynognathus, Diademodon and Trirachodon (Keyser 1973a, b, 1978; Pickford 1995; Smith & Swartt 2002). In the Karoo basin of South Africa, the upper part of the Burgersdorp Formation yields the upper part of the Cynognathus Assemblage Zone (subzone C of Hancox 2000, see discussion above and Fig. 5). Characteristic taxa are the amphibian Paracylotosaurus, the dicynodonts Cynognathus, Diademodon and Cricodon, and the dicynodonts Angonisaurus and Kannemeyeria, which support a Perovkan age assignment (e.g. Hancox & Rubidge 1994, 1996; Damiani 2001; Damiani & Hancox 2003; Abdala et al. 2005). Paracylotosaurus is also known from the Denwa Formation in the Satpura basin, India and the Wianamatta Group of the Sydney basin, Australia (Damiani & Hancox 2003), so these may also be Perovkan records.

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The lower Manda Formation in Tanzania produces the amphibian Eryosuchus, the rhynchosaur Stenaulorhynchus, the archosaur ‘Mandasuchus’, the dicynodonts Shansiodon (¼ Tetragonius) and Angonisaurus and the cynodont Scalenodon (Huene 1938a, b; Crompton 1955; Cruickshank 1965, 1967; Cox & Li 1983; Damiani 2001). This is a Perovkan assemblage. Comments. Lucas (1998a) defined the Perovkan LVF as the time between the FAD of the dicynodont Shansiodon and the FAD of the temnospondyl Mastodonsaurus. Its characteristic assemblage is the tetrapod fauna from the Russian Donguz svita, so the land-vertebrate biochronology shifts here from superposed South African assemblages (the characteristic assemblages of the Lootsbergian and Nonesian LVFs) to the superposed Russian assemblages (the characteristic assemblages of the Perovkan and Berdyankian LVFs). This geographical shift poses problems for the biochronology, particularly in demonstrating the temporal succession (and not overlap) of Nonesian and Perovkan assemblages. Indeed, the reassignment of the upper ‘Cynognathus zone’ to the Perovkan LVF discussed above directly reflects such problems (Hancox 2000; Abdala et al. 2005; Lucas et al. 2007e). The easiest way to reduce ambiguity here was to redefine the beginning of the Perovkan as the FAD of Eocyclotosaurus (Lucas et al. 2007e). Perovkan tetrapod assemblages are best known in Russia and China where they contain numerous dicynodonts. Correlatives are either dicynodont dominated (Manda Formation, upper Burgersdorp Formation) or amphibian dominated (upper Moenkopi, upper Buntsandstein).

Berdyankian LVF Definition. The term Berdyankian LVF is the time interval between the FAD of the amphibian Mastodonsaurus giganteus and the FAD of the phytosaur Parasuchus (¼ Paleorhinus) (Fig. 1). The characteristic tetrapod assemblage is the vertebrate fossil assemblage of the Bukobay svita in the Russian Urals (e.g. Ivakhnenko et al. 1997; Shishkin et al. 2000b). Relevant vertebrate-fossil localities are near the Berdyank River, from which the LVF takes its name. The characteristic Berdyankian tetrapod assemblage is directly superposed on the characteristic Perovkan assemblage. The beginning of the Berdyankian is defined by the FAD of Mastodonsaurus giganteus, whereas the end of the Berdyankian is the beginning of the Otischalkian, which is defined by the FAD of Parasuchus. Characteristic tetrapod assemblage. The assemblage from the Bukobay Formation includes an

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anthracosaur, the amphibians Mastodonsaurus, Bukobaja, ?Cyclotosaurus, Plagioscutum and Plagiosternum, an erythrosuchid, a rauisuchid, and the dicynodonts ‘Elephantosaurus jachimovitschi’ Vyushkov (a Stahleckeria-like form) and a generically indeterminate kannemeyeriid (Shishkin et al. 1995b, 2000a, b; Ivakhnenko et al. 1997; Battail & Surkov 2000; Gower & Sennikov 2000). Index fossils. The following tetrapod genera are common and/or widespread enough to be index fossils of the Berdyankian (Fig. 4): the cynodont Massetognathus and the dicynodonts Dinodontosaurus and Stahleckeria. The LO of the amphibian Mastodonsaurus giganteus is Berdyankian. An acme in plagiosaur diversity and abundance characterizes Berdyankian time. No procolophonids are known from Berdyankian strata (Cisneros 2008a), but this must be due to a lack of discovery, not a real absence, as both pre- and post-Berdyankian procolophonids are known. Principal correlatives. The Lettenkohle (Lettenkeuper, Lower Keuper, Erfurt Formation) in Germany and the Chanarian LVF localities in Argentina and Brazil are the principal correlatives of the Berdyankian type assemblage. The Lettenkohle record is important because it establishes the Ladinian age of at least part of the Berdyankian (see below). The Lettenkohle fossils are from the Grenze bonebed, the laterally equivalent/overlying Vitriolschiefer and the Kupferzell locality, so they are above the unconformity that separates the Keuper from the underlying Muschelkalk. Lettenkohle tetrapods include a chroniosuchian, the amphibians Mastodonsaurus giganteus, Callistomordax, Plagiosternum, Plagiosuchus and Kupferzella, the rauisuchian Batrachotomus, the prolacertiform Tanystropheus and small cynodonts (e.g. Wild 1978, 1980; Schoch 1997, 2000a; Lucas 1999; Schoch & Werneburg 1999; Witzmann et al. 2008; Damiani et al. 2009; Gower & Schoch 2009). A Dinodontosaurus-like humerus from the Vitriolschiefer (Lucas & Wild 1995) may link the Lettenkohle to the South American Chanarian. However, a Dinodontosaurus-like radius is also known from the upper Anisian interval of the Muschelkalk in Germany, so this may indicate that the Berdyankian also encompasses part of late Anisian time (Lucas 2007b). The Chan˜ares local fauna from the Ischichuca (formerly Chan˜ares) Formation of the Ischigualasto– Villa Unio´n basin of northwestern Argentina includes various archosaurs such as Tarjadia, Lagerpeton, Marasuchus and Chanaresuchus, the dicynodont Dinodontosaurus, the traversodontids Massetognathus and Megagomphodon, the chiniquodontid Probelesodon and the probainognathid

Probainognathus (Bonaparte 1970; Romer 1973; Sereno & Arcucci 1993, 1994; Lucas & Harris 1996; Bonaparte 1997; Arcucci & Marsicano 1998; Hsiou et al. 2002). Bonaparte (1966, 1967, 1982) based the Chanarian ‘provincial age’ on this assemblage. The lower part of the Santa Maria Formation in the Parana´ basin of Rio Grande do Sul, Brazil yields vertebrate fossil assemblages from Candelaria and Chiniqua´ considered by Barberena (1977) and Barberena et al. (1985) to be two different local faunas of different ages. Lucas (2002) regarded them as a single biostratigraphic assemblage that includes a procolophonid, archosaurs, the dicynodonts Dinodontosaurus and Stahleckeria, chiniquodontids, and the traversodontids Massetognathus, Belesodon, Traversodon, Exaeretodon, Santacruzodo, Protuberum and Probelesodon (e.g. Abdala & Ribeiro 2003; Cisneros et al. 2004; Langer et al. 2007; Reichel et al. 2009). This assemblage and the Chanarian type assemblage in Argentina are assigned a Berdyankian age based largely on their dicynodonts and traversodontids and their stratigraphic position, which places them between the Nonesian and the Adamanian. Comments. Previously, I used the FAD of the genus Mastodonsaurus to define the beginning of the Berdyankian. This was based on a taxonomy in which Mastodonsaurus (typified by the species M. giganteus) was distinguished from the older (Perovkan) Heptasaurus (e.g. Schoch 1999; Schoch & Milner 2000). However, taxonomists who study these amphibians have suggested that Mastodonsaurus and Heptasaurus be combined into a single genus, Mastodonsaurus (Rayfield et al. 2009). Thus, I now use the FAD of the species M. giganteus to define the beginning of the Berdyankian so as not to be subject to the shifting opinions of taxonomists revising the genus-level taxonomy of stereospondyl amphibians. This preserves the original intent of the Berdyankian, as no temnospondyl worker has advocated the synonymy of Heptasaurus cappelensis and Mastodonsaurus giganteus at the species-level. As noted by Lucas (1998a), global correlations within the Berdyankian interval are confounded by the near endemism of South American tetrapod assemblages that are apparently of this age (the Dinodontosaurus faunas of Argentina and Brazil, classically assigned to the Chanarian LVA of Bonaparte 1966, 1967). Recognition of Berdyankianage assemblages in Russia and Germany is rendered easy by the presence of the key taxon Mastodonsaurus giganteus (Lucas 1999). The Berdyankian is difficult to correlate globally, largely because of a paucity of tetrapod assemblages of this age. Two clusters of localities (European and South American) are equated,

TRIASSIC TETRAPODS

largely on the basis of the Lettenkohle dicynodont and the conclusion that ‘Elephantosaurus’ is a ‘stahleckeriid’, possibly a synonym of Stahleckeria (Lucas & Wild 1995). The South American Chanarian LVF thus is the provincial secondary standard correlative to the Berdyankian. The Berdyankian may be relatively long, at least correlative to the latest Anisian and Ladinian (see below). Nevertheless, Berdyankian tetrapod fossil assemblages probably only represent the earlier part of this time interval. Indeed, the paucity of tetrapod assemblages of Berdyankian age represents one of the most substantial deficits in the global record of Triassic tetrapods. This is an important deficit because many characteristic Late Triassic tetrapod taxa, such as metoposaurs, phytosaurs, aetosaurs and dinosaurs, so far lack evolutionary antecedents that should occur in Berdyankian-age strata.

Otischalkian LVF Definition. The Otischalkian LVF is the time interval between the FADs of the phytosaurs Parasuchus (¼ Paleorhinus) and Rutiodon (Fig. 1). Lucas & Hunt (1993a) proposed the Otischalkian LVF based on the vertebrate fossil assemblage of the Colorado City Formation of the Chinle Group near the defunct town of Otis Chalk, Howard County, Texas, USA (Lucas & Anderson 1993a, b, 1994, 1995; Lucas et al. 1993, 1994, 1997a). The beginning of the Otischalkian is the FAD of Parasuchus. The end of the Otischalkian is the beginning of the Adamanian, which is defined by the FAD of the phytosaur Rutiodon. Characteristic tetrapod assemblage. The characteristic tetrapod assemblage of the Otischalkian is the assemblage of vertebrate fossils from just north of the defunct town of Otis Chalk in Howard County, Texas. Lucas et al. (1993) reviewed the fauna, which is from the Colorado City Formation of the Chinle Group. The following taxa are present: the amphibians Latiscopus, Buettneria and Apachesaurus, a procolophonid, the rhynchosaur Otischalkia, the archosaurs Doswellia, Trilophosaurus (¼ Malerisaurus) and Poposaurus, the aetosaurs Longosuchus (¼ Lucasuchus) and Coahomasuchus, and the phytosaurs Parasuchus and Angistorhinus (Lucas et al. 1993; Long & Murry 1995; Heckert & Lucas 1999; Spielmann et al. 2006c). Index fossils. The following tetrapod genera are restricted to Otischalkian time and are widespread and/or common enough to be useful as index fossils (Fig. 6): the aetosaur Longosuchus, and the archosaur Doswellia. Parasuchus and Angistorhinus are mostly of Otischalkian age, but also have early Adamanian records. The dicynodont Placerias

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has both Otischalkian and earliest Adamanian records. The LOs of the widespread temnospondyl Metoposaurus and of the rhynchosaur Hyperodapedon are Otischalkian, and these taxa are also known in Adamanian strata (Lucas et al. 2002a, 2007e). Principal correlatives. Besides Chinle Group correlatives, principal Otischalkian vertebrate assemblages are from the Sanfordian interval of the Newark Supergroup basins of eastern North America; Schilfsandstein (Stuttgart Formation) of the German Keuper; the Irohalene Member (T4) of the Timesgadiouine Formation, Argana Group, Morocco; and the basal part of the Maleri Formation, Pranhita –Godavari Valley, India. Otischalkian principal correlatives and the characteristic tetrapod assemblage encompass a broad geographical range of Chinle Group outcrops in Wyoming, New Mexico and Texas. They occur in units of the lower part of the Chinle Group that have been correlated with each other on a lithostratigraphic basis (Lucas 1993b). The most well-known principal correlative of the type Otischalkian fauna in the Chinle Group is the vertebrate-fossil assemblage from the Popo Agie Formation of Wyoming, principally Fremont County (Branson & Mehl 1928; Mehl 1928; Colbert 1957; Lucas 1994; Lucas et al. 2002a) that includes the amphibian Buettneria, the phytosaurs Parasuchus and Angistorhinus, the aetosaur Desmatosuchus, the archosaurs Poposaurus and Heptasuchus, the rhynchosaur Hyperodapedon, and the dicynodont Placerias. A less well-known principal correlative is the small assemblage from the Salitral Formation in Rio Arriba County, New Mexico that consists of a metoposaur, Longosuchus, a phytosaur, and an indeterminate dinosaur (Lucas & Hunt 1992). Heckert (2004; Heckert & Lucas 2006) provided some microvertebrate basis for recognition of the Otischalkian in Chinle Group strata, such as the LO of the ‘dinosaur’ Protecovasaurus and the archosaur Trilophosaurus buettneri (also see Spielmann et al. 2008). In the Newark Supergroup of eastern North America, the stratigraphically lower formations of the Deep River, Gettysburg, Newark and Fundy basins contain two distinct vertebrate fossil assemblages. The older of these was used by Huber et al. (1993b) as the basis of the Sanfordian LVF, after the characteristic assemblage from the middle Pekin Formation in the Sanford sub-basin of the Deep River basin complex. An age-equivalent assemblage from the middle Wolfville Formation (Fundy basin) is also assigned to this LVF. The collective Newark tetrapod fauna of this Sanfordian LVF includes the amphibian Metoposaurus, procolophonids, the traversodontids Arctotraversodon and Plinthogomphodon, the dicynodont Placerias, the rhynchosaur Hyperodapedon, the archosaur

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S. G. LUCAS

taxa

Otischalkian

Adamanian

Revueltian

Apachean

amphibians: Apachesaurus Buettneria Metoposaurus phytosaurs: Angistorhinus Mystriosuchus Nicrosaurus Parasuchus Pseudopalatus Redondasaurus Rutiodon aetosaurs: Aetosaurus Desmatosuchus Longosuchus Paratypothorax Redondasuchus Rioarribasuchus Stagonolepis Typothorax others: Doswellia Eudimorphodon Hyperodapedon Placerias Revueltosaurus Riojasaurus Fig. 6. Temporal ranges of selected genera of Late Triassic tetrapods.

Doswellia, the aetosaurs Desmatosuchus and Longosuchus, indeterminate rauisuchians (‘Zamotus’), the rauisuchian Postosuchus, the ‘sphenosuchian’ Dromicosuchus, indeterminate phytosaur fragments and fragmentary dinosaur remains (e.g. Cope 1871; Olsen et al. 1989; Hunt & Lucas 1990; Huber et al. 1993a; Hunt 1993; Sues et al. 1994, 1999, 2003; Langer et al. 2000b; Lucas et al. 2002a; Peyer et al. 2008; Dilkes & Sues 2009). The Sanfordian correlates with the Chinle Group Otischalkian LVF based on the shared presence of Buettneria, Hyperodapedon, Desmatosuchus, Longosuchus, Doswellia, and Placerias. In Germany, the Schilfsandstein produces Metoposaurus and Parasuchus but lacks Stagonolepis, so it can be assigned an Otischalkian age (Hunt &

Lucas 1991; Lucas 1999; Schoch & Werneburg 1999; Hungerbu¨hler 2001b). The 500-m-thick Irohalene Member of the Timesgadiouine Formation (interval T-5 of Dutuit 1966; Tixeront 1971) has produced most of the Late Triassic vertebrate fauna from Morocco. It contains the majority of vertebrate fossil localities described by Dutuit (1972, 1976, 1977, 1988, 1989a, b). Most of these occur in the lower part of the member and have produced a moderately diverse fauna that includes the amphibians Almasaurus and Dutuitosaurus, the phytosaur Parasuchus, the aetosaur Longosuchus, the dicynodont Placerias (¼ Moghreberia, ¼ Azarifeneria: Cox 1991; Lucas & Wild 1995), the dinosauriform Azendohsaurus (Gauffre 1993; Lucas 1998b; Jalil 1999) and at least one

TRIASSIC TETRAPODS

ornithischian dinosaur. Several of Dutuit’s (1976) localities occur in the upper part of the Irohalene Member, which is a distinct faunal horizon that includes the amphibian Arganasaurus, the phytosaur Angistorhinus, and the dicynodont Placerias. The presence of Parasuchus, Angistorhinus, Longosuchus and Placerias supports assigning the Irohalene Member tetrapod assemblage(s) an Otischalkian age. In the Pranhita–Godavari Valley of India, the basal Maleri Formation produces a tetrapod assemblage that includes the amphibian Metoposaurus, the rhynchosaur Paradapedon, the phytosaur Parasuchus, the archosaur ‘Malerisaurus’, an aetosaur, the theropod dinosaur Alwalkeria, a prosauropod (‘cf. Massospondylus’ of Kutty & Sengupta 1989), a large dicynodont, and the cynodont Exeraetodon (e.g. Huene 1940; Jain et al. 1964; Roychowdhury 1965; Chatterjee 1967, 1974, 1978, 1980a, 1982, 1987; Chatterjee & Roychowdhury 1974; Jain & Roychowdhury 1987; Bandyopadhyay & Sengupta 2006; Spielmann et al. 2006c). This is the only welldescribed Upper Triassic tetrapod assemblage from the Pranhita–Godavari Valley. It includes Parasuchus and Metoposaurus, taxa indicative of a likely Otischalkian age. Comments. The Otischalkian LVF was originally defined as the time between the FADs of the phytosaurs Parasuchus (¼ Paleorhinus) and Rutiodon (Lucas & Hunt 1993a; Lucas et al. 1997a; Lucas 1998a). It is important to note that a little advertised petition to the International Commission on Zoological Nomenclature by Chatterjee (2001) resulted in establishing a diagnostic lectotype for Parasuchus (long a nomen dubium: Hunt & Lucas 1991), so that this name should be regarded as the senior synonym of Paleorhinus (Lucas et al. 2007c). Furthermore, even though Hunt & Lucas (1991) provided a careful taxonomic revision of Parasuchus, and provided a clear diagnosis of the genus that has never been contested, some taxonomists have relegated all primitive phytosaurs to a metataxon (grade) and then claimed these phytosaurs (long and widely known as Paleorhinus/Parasuchus) are of no value to biostratigraphy (e.g. Fara & Hungerbu¨hler 2000; Rayfield et al. 2005, 2009). I reject such an approach to primitive phytosaur taxonomy and recognize Parasuchus as a diagnosable genus (Lucas et al. 2007c). I have long regarded Parasuchus as a robust index taxon of the Otischalkian (Hunt & Lucas 1991; Lucas et al. 2007c, d). However, recently developed Upper Triassic conchostracan biostratigraphy (Kozur & Weems 2005, 2007) and European records of the characteristic Adamanian aetosaur Stagonolepis suggest that some Parasuchus records should be considered early Adamanian in age (Kozur & Weems 2005). Thus, if all Stagonolepis records

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are regarded as Adamanian (as they were by Lucas 1998a), and the conchostracan-based correlations of the Adamanian are accepted, then records of Parasuchus from the German Kieselsandstein and Blasensandstein and the Polish Krasiejo´w locality are Adamanian. This is also consistent with the Chinle Group record of Parasuchus at the Placerias/Downs quarries in the Bluewater Creek Formation of the Chinle Group in Arizona, in what I have regarded as oldest Adamanian strata (Lucas et al. 1997a). Thus, recognizing that Parasuchus records are not strictly Otischalkian (some are early Adamanian: Fig. 6), and that Stagonolepis records are strictly Adamanian, clarifies correlation in the Otischalkian–Adamanian interval. The Otischalkian index taxa Longosuchus (¼ Lucasuchus) and Doswellia still stand. Metoposaurus also has only Otischalkian and early Adamanian records, though Milner & Schoch (2004) recently claimed its presence in the Revueltian Stubensandstein of Germany, a claim that met a detailed refutation from Lucas et al. (2007e). The last Otischalkian index fossil listed by Lucas (1998a) is the phytosaur Angistorhinus. Its records are Otischalkian (Long & Murry 1995) except one, near Lamy, New Mexico, USA, where it co-occurs with Rutiodon in the earliest Adamanian (Hunt et al. 1993, 2005) (Fig. 7).

Adamanian LVF Definition. The Adamanian is the time interval between the FAD of the phytosaur Rutiodon and the FAD of the aetosaur Typothorax coccinarum (Fig. 1). Lucas & Hunt (1993a) based the Adamanian LVF on the vertebrate fauna of the Blue Mesa Member of the Petrified Forest Formation in the Petrified Forest National Park, Arizona, USA (Lucas 1993b; Lucas & Hunt 1993a; Lucas et al. 1997a). Lucas (1998a) termed this the Rutiodon Assemblage Zone. The beginning of the Adamanian is defined as the FAD of the phytosaur Rutiodon. The end of the Adamanian is the beginning of the Revueltian, which is defined by the FAD of the aetosaur T. coccinarum. Characteristic tetrapod assemblage. The characteristic tetrapod assemblage of the Adamanian is the assemblage of vertebrate fossils found in the Blue Mesa Member of the Petrified Forest Formation in the Petrified Forest National Park, near the defunct railroad siding of Adamana, Arizona. Recent faunal lists have been provided by Murry & Long (1989), Long & Murry (1995), Heckert et al. (2005a) and Parker et al. (2006). The fauna includes the following tetrapods: the amphibians Apachesaurus and Buettneria, the aetosaurs Desmatosuchus

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S. G. LUCAS

Shinarump Formation

Chinle Group Petrified Forest Formation Blue Mesa Member

Sonsela Member Rainbow Forest Bed

Jim Camp Wash Bed

Agate Bridge Bed

Painted Desert Member

Los Esteros Member

Tres Lagunas Member

Garita Creek Formation

Trujillo Formation

Bull Canyon Formation

Santa Rosa Formation Tecolotito Member

east–central NM

Chinle Group

Petrified Forest National Park

Bluewater Creek Formation

phytosaurs

Rutiodon Paleorhinus Angistorhinus

Pseudopalatus Stagonolepis

aetosaurs

Desmatosuchus haplocerus

Rioarribasuchus chamaensis

Paratypothorax Placerias

Typothorax antiquum

others

Poposaurus Revuletosaurus hunti

Typothorax coccinarum R. callenderi

Trilophosaurus buettneri Trilophosaurus jacobsi REVUELTIAN

ADAMANIAN St. Johnsian

Lamyan

Barrancan

LVF subLVF

Fig. 7. Lithostratigraphy and tetrapod biostratigraphy of the Adamanian and Revueltian intervals in east–central New Mexico and in the Petrified Forest National Park, Arizona. The Lamyan interval is shaded (after Hunt et al. 2005).

(¼ Acaenasuchus), Stagonolepis, Adamanasuchus and Paratypothorax, Rutiodon-grade phytosaurs (including Leptosuchus and Smilosuchus), the rauisuchian Postosuchus, the archosaurs Hesperosuchus, Acallosuchus, Parrishea and Vancleavea,

and the dicynodont Placerias, as well as many microvertebrate taxa. Index fossils. The following tetrapod genera are restricted to Adamanian time and are widespread

TRIASSIC TETRAPODS

and/or common enough to be useful as index fossils (Fig. 6): Rutiodon-grade phytosaurs, including Leptosuchus and Smilosuchus, the trilophosaurid archosaur Spinosuchus and the aetosaur Stagonolepis. The HO of dicynodonts was long thought to be Adamanian. However, there is a putative Cretaceous record from Australia (Thulborn & Turner 2003), and Dzik et al. (2008) recently reported a Triassic dicynodont from Poland in strata they deemed Rhaetian based on palaeobotany. The HO of the widespread rhynchosaur Hyperodapedon is Adamanian (Lucas & Heckert 2001; Lucas et al. 2002a) (Fig. 8). Within the Chinle Group, various microvertebrate taxa, including Colognathus, Tecovasurus, and Crosbysaurus, are index taxa of the Adamanian (Heckert 2004; Heckert & Lucas 2006). Principal Correlatives. Besides the Chinle Group correlatives, major Adamanian faunas are those of the Conewagian interval of the Newark Supergroup basins of eastern North America; Lossiemouth Sandstone Formation, Scotland; Lehrberg Schichten interval of the German Keuper; the Krasiejo´w locality in Poland; Ischigualasto Formation, Argentina; and upper Santa Maria Formation, Brazil. In the Chinle Group, Adamanian vertebrates are widespread and include the vertebrate fossil assemblages of the Placerias and Downs’ quarries, Bluewater Creek Formation, Arizona (Camp & Welles 1956; Kaye & Padian 1994; Long & Murry 1995; Lucas et al. 1997a; Heckert 2004; Heckert et al. 2005a); the Bluewater Creek Formation and Blue Mesa Member of the Petrified Foreset Formation in the Blue Hills, Arizona; the Bluewater Creek Formation and Blue Mesa Member of the Petrified Forest Formation, McKinley and Cibola counties, New Mexico (Heckert 1997); the Los Esteros and Tres Lagunas members, Santa Rosa Formation, vicinity of Lamy, Santa Fe County, New Mexico (Hunt et al. 2005) (Fig. 7); Garita Creek Formation, Santa Rosa and vicinity, Guadalupe County, New Mexico (Hunt & Lucas 1993a); and Tecovas Formation, West Texas (Murry 1986, 1989; Long & Murry 1995). The fauna at the Placerias and Downs’ quarries has most recently been discussed by Kaye & Padian (1994), Long & Murry (1995), Lucas et al. (1997a) and Heckert (2004). It includes the amphibians Buettneria and Apachesaurus, the prolacertiform Tanytrachelos, the phytosaurs Parasuchus and Rutiodon/Leptosuchus, the aetosaurs Stagonolepis and Desmatosuchus (¼ Acaenasuchus), the rauisuchid Postosuchus, the archosaurs Trilophosaurus, Acallosaurus, Poposaurus, Chatterjeea, Hesperosuchus, Tecovasaurus and cf. Uatchitodon, an indeterminate ceratosaur and the dicynodont Placerias.

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The following tetrapod taxa are known from the Los Esteros Member, Santa Rosa Formation, near Lamy, New Mexico: the amphibian Apachesaurus, the phytosaurs Rutiodon and Angistorhinus, the aetosaurs Desmatosuchus, Tecovasuchus and Stagonolepis and the dicynodont cf. Ischigualastia (Hunt & Lucas 1993a, 1994; Hunt et al. 2005; Heckert et al. 2007b). The overlying Garita Creek Formation contains the following taxa: the amphibian Buettneria, phytosaurs, rauisuchians, and the aetosaurs Desmatosuchus, Stagonolepis and Paratypothorax (Hunt et al. 2005). The Tecovas Formation of West Texas yields the following tetrapod taxa: the amphibians Buettneria and Apachesaurus, the probable tetrapod Colognathus, the archosauromorphss Trilophosaurus, Parrishea, Tecovasaurus, and Crosbysaurus, the phytosaurs Rutiodon, Leptosuchus and Smilosuchus, the aetosaurs Desmatosuchus and Stagonolepis, the rauisuchian Postosuchus, and the oldest known mammal, Adelobasileus (Lucas & Luo 1993; Lucas et al. 1994; Long & Murry 1995; Spielmann et al. 2008). In the Deep River basin of North Carolina, an assemblage of the Conewagian LVF from the Cumnock Formation is superposed on the characteristic Sanfordian assemblage. Conewagian assemblages are characterized by the tetrapod assemblage in the basal Gettysburg Formation (Kozur & Weems 2010) along Little Conewago Creek in south-central Pennsylvania (Gettysburg basin: Huber et al. 1993b; Sullivan et al. 1995; Lucas & Sullivan 1997) and also are known from the Cow Branch Formation (Dan River basin), and upper Stockton and Lockatong formations (Newark basin). The most widespread and characteristic Conewagian tetrapod is the phytosaur Rutiodon, which co-occurs with the amphibian Buettneria, archosaurs of uncertain affinity, an aetosaur (Desmatosuchus), one or more ‘ornithischian dinosaurs’ (e.g. Pekinosaurus and Galtonia), the archosaur Tanytrachelos (¼ ?Gwyneddosaurus) and the lepidosauromorph Icarosaurus (e.g. Emmons 1856; Olsen 1980, 1988; Olsen et al. 1989; Sues 1992; Huber et al. 1993a; Hunt 1993; Hunt & Lucas 1994; Doyle & Sues 1995; Lucas & Huber 2003). Conewagian assemblages correlate with the Adamanan LVF of the Chinle Group, based on the shared presence of Buettneria, Rutiodon and other Rutiodon-grade phytosaurs (Smilosuchus of Long & Murry 1995), Desmatosuchus and broadly similar ‘ornithischian dinosaurs’ (e.g. Murry & Long 1989; Lucas et al. 1992, 1997a; Huber et al. 1993b; Hunt 1993; Hunt & Lucas 1994; Heckert 2004). The tetrapod assemblage of the Lossiemouth Sandstone Formation of Grampian (Elgin) Scotland comes from small quarries and the coastal section at Lossiemouth. Benton & Spencer (1995) provided a

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S. G. LUCAS

Stage

Carnian Adamanian

Wyoming

Chinle Group

lvf

Otischalkian

Popo Agie Formation

Nova Scotia

Wolfville Formation

Scotland

Lossiemouth Sandstone

Maleri Formation India

lower assemblage

upper assemblage

Zimbabwe

Pebbly Arkose Formation

Madagascar

Isalo II Beds

Caturrita Formation

Hyperodapedon assemblage zone

Brazil

Santa Maria Formation (Alemoa Member)

Argentina

Hyperodapedon record

Ischigualasto Formation

Hyperodapedon biochron Fig. 8. Global correlation of Hyperodapedon localities, which identify a Hyperodapedon biochron of Otischalkian– Adamanian age.

TRIASSIC TETRAPODS

detailed summary and indicate that all sites come from a narrow stratigraphic range, so I treat the vertebrates as a single biostratigraphic assemblage. It includes the procolophonid Leptopleuron, the sphenodontid Brachyrhinodon, the rhynchosaur Hyperodapedon, the aetosaur Stagonolepis, the ornithosuchid Ornithosuchus, the crocodylomorph Erpetosuchus, the probable ornithodiran Scleromochlus and the ‘dinosaur’ Saltopus. The presence of Hyperodapedon and Stagonolepis supports correlation of this assemblage to the Chinle Group Adamanian. In Germany, the stratigraphic interval between the Schilfsandstein and the Stubensandstein (Lehrberg Schichten, Blasensandstein and Kieselsandandstein) produces Stagonolepis, Parasuchus and Metoposaurus (e.g. Lucas 1999), and is assigned an Adamanian age (Kozur & Weems 2005). In Poland, the Krasiejo´w tetrapod assemblage includes the amphibians Cyclotosaurus and Metoposaurus, the phytosaur Parasuchus, the aetosaur Stagonolepis, the rauisuchian Teratosaurus and the dinosauriform Silesaurus (Dzik 2001, 2003; Sulej 2002, 2005, 2007; Sulej & Majer 2005; Dzik & Sulej 2007; Lucas et al. 2007d ). This assemblage is from strata c. 80 m above the Reed Sandstone (a Schilfsandstein equivalent) that are homotaxial to the German Lehrberg Schichten and is of Adamanian age. In Argentina, the Ischigualasto Formation is 500 –900 m thick and consists of drab mudstones, tuffs and sandstones that produce an extensive tetrapod assemblage including: the amphibian Promastodonosaurus, the archosaurs Saurosuchus, Sillosuchus, and Proterochampsa, the aetosaur Stagonolepis (¼ Aetosauroides), the rhynchosaur Hyperodapedon, the dinosaurs Herrerasaurus (¼ ?Ischisaurus ¼ Frenguellisaurus), Eoraptor and Pisanosaurus, the chiniquodontid cynodont Chiniquodon, the gomphodont cynodonts Exeraetodon, Proexaraetodon, and Ischignathus and the dicynodont Ischigualastia (e.g. Cabrera 1944; Reig 1959, 1961, 1963; Casamiquela 1960, 1962; Cox 1965; Bonaparte 1976; Rogers et al. 1993; Sereno et al. 1993; Bonaparte 1997; Alcober & Parrish 1997; Heckert & Lucas 2002c). The assemblage slightly overlaps and mostly overlies the Herr Toba bentonite that yielded a 40Ar/39Ar age of 227.8 + 0.3 Ma (Rogers et al. 1993), which was ‘recalculated’ to 231.4 Ma by Irmis and Mundil (2008). In Brazil, the principal Upper Triassic vertebrate assemblage from the Santa Maria Formation is from the vicinity of Santa Maria City. This is the Rhynchocephalia assemblage zone of Barberena (1977) or the Scaphonyx assemblage of Barberena et al. (1985), from the upper part of the Santa Maria Formation. The assemblage consists of abundant fossils of the rhynchosaur Hyperodapedon and

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the aetosaur Stagonolepis (¼ Aetosauroides); traversodontids, proterochampsids; the archetypal rauisuchian Rauisuchus and the primitive dinosaur Staurikosaurus (Barberena et al. 1985; Lucas 2002; Lucas & Heckert 2001; Langer et al. 2007). Clearly, the presence of Scaphonyx and Stagonolepis (‘Aetosauroides’) supports correlation with the vertebrates of the Ischigualasto Formation in Argentina, and therefore an Adamanian (¼ Ischigualastian) age (Lucas & Heckert 2001; Heckert & Lucas 2002c; Lucas 2002). The tetrapod assemblage of the Caturrita Formation, which overlies the Santa Maria Formation, includes a mastodonsauroid amphibian, the sphenodont Clevosaurus, the rhynchosaur Hyperodapedon, the proterochampsid Proterochampsa, the dinosaurs Guabisaurus and Saccasaurus, a phytosaur, the cynodonts Exaeretodon and Riograndia, the dicynodont Ischigualastia (¼ Jachaleria) and cynodonts (Arau´jo & Gonzaga 1980; Barbarena et al. 1985; Dornelles 1990; Bonaparte et al. 1999, 2001; Kischlat & Lucas 2003; Ferigolo & Langer 2006; Bonaparte & Sues 2006; Bonaparte et al. 2007; Langer et al. 2007; Dias-da-Silva et al. 2009). Most South American workers (e.g. Bonaparte 1982; Barberena et al. 1985; Langer 2005a; Rubert & Schultz 2004; Dias-da-Silva et al. 2007; Langer et al. 2007) advocate dividing the Brazilian Upper Triassic tetrapod succession into two biostratigraphically distinct assemblages largely based on their judgment that the dicynodonts Jachaleria and Ischigualastia are not the same taxon. They, therefore, correlate the Brazilian Caturrita Formation to the Argentinian Los Colorados Formation. Langer (2005b) also claimed that the Ischigualastian ¼ Otischalkian þ Adamanian, largely based on not recognizing the temporal range of Hyperodapedon as longer than the temporal range of the Ischigualastian. I do not accept either evaluation of the Brazilian Upper Triassic tetrapod biostratigraphy (Lucas 2002). In the Pranhita –Godavari Valley of India, the upper vertebrate fossil assemblage from the Maleri Formation is stratigraphically above the lower assemblage, but its stratigraphic range is not clear. This upper assemblage includes an aetosaur, prosauropods and a large dicynodont. Chigutisaurid amphibians (Compsocerops and Kuttycephalus: Sengupta 1995) and a ‘Rutiodon-like’ phytosaur are also present (Bandyopadhyay & Sengupta 2006). Therefore, this assemblage may be Adamanian, but needs further documentation. In western Madagascar, the Isalo group (‘Groupe d l’Isalo’ of Besarie 1930; also see Besarie & Collignon 1960, 1971) has long been divided into Isalo I, Isalo II and Isalo III based on perceived geological age. The Isalo II strata yield Late Triassic tetrapods, including metoposaurs, sphenodontids,

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phytosaurs, the rhynchosaur Hyperodapedon, the aetosaur Desmatosuchus, the archosaur Azendohsaurus, cynodonts and dicynodonts (Guth 1963; Westphal 1970; Dutuit 1978; Buffetaut 1983; Flynn et al. 1999, 2000, 2008; Langer et al. 2000a; Lucas et al. 2002a; Burmeister et al. 2006). The stratigraphic range of the Isalo II tetrapods is c. 1200 m, but the rhynchosaur Hyperodapedon is one of the stratigraphically lowest taxa in the assemblage. This means the Isalo assemblage is no older than Otischalkian and, based on the Desmatosuchus record, likely to be Adamanian. Comments. Lucas (1998a) listed as Adamanian index fossils the rhynchosaur Scaphonyx, the aetosaur Stagonolepis and Rutiodon-grade phytosaurs (including Leptosuchus and Smilosuchus). The dicynodont Ischigualastia (¼ Jachaleria) was also considered an Adamanian index taxon. Taxonomic revisions and range extensions have necessitated a reconsideration of some of these index taxa. Stagonolepis now co-occurs with Parasuchus at Krasiejo´w in southern Poland (Dzik 2001; Lucas et al. 2007d). This lends support to Heckert & Lucas’ (2000) conclusion that Ebrachosaurus singularis Kuhn 1936, from the Adamanian German Blasensandstein (type destroyed in World War II), was based on specimens of Stagonolepis. These European Adamanian records of Stagonolepis are consistent with regarding its stratigraphically lowest records in North America, such as at the Placerias/ Downs quarries in Arizona, as early Adamanian (Lucas et al. 1997a). An extensive revision of Late Triassic rhynchosaurs (Langer & Schultz 2000; Langer et al. 2000a, b) indicates that specimens previously assigned to Scaphonyx are mostly of Hyperodapedon. Lucas et al. (2002a) reviewed these records in detail and demonstrated that a Hyperodapedon biochron is of Otischalkian and Adamanian age (Fig. 8). Thus, the rhynchosaur Hyperodapedon cannot be used to discriminate the Otischalkian and Adamanian. Largely based on this, Langer (2005a, b; also see Schultz 2005) claimed that the Otischalkian and Adamanian cannot be distinguished and they should be abandoned and replaced by a single LVF, the Ischigualastian. To do so, Langer (2005b) dismissed phytosaur-based distinctions of the Otischalkian and Adamanian, basing his rejection largely on the cladotaxonomy of primitive phytosaurs advocated in published abstracts by Hungerbu¨hler (2001a; Hungerbu¨hler & Chatterjee 2002). Langer (2005b) also rejected aetosaur-based correlations based on the taxonomy of South American aetosaurs published by Heckert & Lucas (2000) and Lucas & Heckert (2001). Thus, Langer (2005b, p. 228) states that ‘Stagonolepis wellesi lacks a unique ornamentation

pattern of its dorsal paramedian osteoderms’, contrary to the published work of Lucas & Heckert, as well as those of Long & Ballew (1985), Parrish (1994), Long & Murry (1995) and Parker (2007), among others. Langer (2005b) also used the conclusions of Sulej (2002) regarding the taxonomy of Metoposaurus and Buettneria to question using amphibians to distinguish the Otischalkian and Adamanian. However, a review of the metoposaur specimens described by Sulej (2002) does not support some of his basic anatomical observations or his taxonomy (Lucas et al. 2007d). Rayfield et al. (2005, 2009) also argued for merging of the Otischalkian and Adamanian based largely on the same arguments as Langer (2005b), but Lucas et al. (2007e) have presented a detailed refutation of their arguments. What these workers have failed to recognize is that: (1) Otischalkian and Adamanian tetrapod assemblages are stratigraphically superposed and readily distinguished in the Chinle Group of the American Southwest; (2) there is no evidence that the ‘Ischigualastian’ of South America is Otischalkian and much more evidence that it is Adamanian, so Ischigualastian should not be redefined to encompass both Otischalkian and Adamanian time; and (3) recognition of distinct Otischalkian and/or Adamanian assemblages has been achieved in North America, South America, Europe, India and North Africa (e.g. Fig. 8). The fact that Langer (2005b) and Rayfield et al. (2005, 2009) do not accept a well-documented alpha taxonomy of Otischalkian and Adamanian index fossils is not a valid reason to merge the Otischalkian and Adamanian LVFs. Recent work in the Chinle Group of the western USA has refined the stratigraphic ranges of known tetrapod taxa and has produced new records in strata of Adamanian age. These new data are principally from the Petrified Forest National Park in Arizona (Heckert & Lucas 2002a; Hunt et al. 2002; Woody 2003, 2006; Heckert 2004; Woody & Parker 2004; Heckert et al. 2005a) and the extensive exposures of the Chinle Group in east–central New Mexico (Lucas et al. 2001, 2002b), though there are also other new records from the Tecovas and Trujillo formations in Texas (Heckert 2004; Heckert et al. 2006; Martz & Small 2006). Clearly, there is a ‘transitional’ fauna between the Adamanian and Revueltian LVFs (Woody & Parker 2004), and this prompted Hunt et al. (2005) to subdivide the Adamanian into two subfaunachrons, St. Johnsian (older) and Lamyan (younger), of regional biochronological significance (Fig. 7). The aetosaur Tecovasuchus is a St. Johnsian index taxon (Heckert et al. 2007b), whereas the aetosaur Typothorax antiquum is a Lamyan index taxon (Hunt et al. 2005).

TRIASSIC TETRAPODS

Heckert & Lucas (2006) built upon the microvertebrate collections documented by Heckert (2001, 2004) to demonstrate that there are multiple microvertebrate index taxa of Adamanian (St. Johnsian) time, including the xenacanth ‘Xenacanthus’ moorei, the enigmatic tetrapod Colognathus obscurus and the archosaurs (possibly ornithischian dinosaurs) Tecovasaurus murryi, Crosbysaurus harrisae, and Krzyzanowskisaurus hunti. So far, these taxa are presently known only from the Chinle Group of the American Southwest, so they may not be of broad biostratigraphic utility.

Revueltian Definition. The Revueltian is the time interval between the FAD of the aetosaur Typothorax coccinarum and the FAD of the phytosaur Redondasaurus (Fig. 1). Lucas & Hunt (1993a) introduced the term Revueltian LVF to refer to the time equivalent to the vertebrate fossil assemblage of the Bull Canyon Formation in east –central, New Mexico, USA (Lucas et al. 1985; Hunt 1994, 2001; Hunt & Lucas 1997). Lucas (1998a) termed this the Pseudopalatus Assemblage Zone. The name of the LVF is for Revuelto Creek, one of the key collecting areas in eastern New Mexico. Revueltian time begins with the FAD of the aetosaur T. coccinarum. The end of the Revueltian is the beginning of the Apachean, which is defined by the FAD of the phytosaur Redondasaurus. Characteristic tetrapod assemblage. The characteristic tetrapod assemblage of the Revueltian is that of the Bull Canyon Formation in east –central New Mexico (Quay and Guadalupe counties), and the following taxa are present: the amphibian Apachesaurus, the turtle Chinlechelys, the phytosaur Pseudopalatus and other Pseudopalatus-grade phytosaurs, the aetosaurs Rioarribasuchus, Paratypothorax, Typothorax coccinarum, and Aetosaurus, the suchian Revueltosaurus, the ‘dinosaur’ Lucianosaurus, the rauisuchian Postosuchus, the chatterjeeids Shuvosaurus (¼ Effigia) and Chatterjeea, the sphenosuchian Hesperosuchus; and the cynodont Pseudotriconodon (e.g. Hunt 1994, 2001; Lucas et al. 2001; Joyce et al. 2009). Index fossils. The following tetrapod taxa are restricted to Revueltian time and are widespread and/or common enough to be useful as index fossils: the crurotarsan Revueltosaurus, the aetosaurs Aetosaurus, Rioarribasuchus and Typothorax coccinarum, and Pseudopalatus-grade phytosaurs. The pterosaur Eudimorphodon is present in Revueltian assemblages in Italy and Greenland (e.g. Jenkins et al. 2001; Dalla Vecchia 2003, 2006) and can also be considered a Revueltian index taxon (but

469

see Dalla Vecchia 2009). The stratigraphic cooccurrence of dinosaurs and dinosauromorphs (Sullivan & Lucas 1999; Ezcurra 2006; Irmis et al. 2007; Nesbitt et al. 2007, 2009; Spielmann et al. 2007b; Nesbitt & Chatterjee 2008) also aids in recognition of Revueltian time. Principal correlatives. Besides Chinle Group assemblages, which are primarily from Texas, New Mexico and Arizona (e.g. Zeigler et al. 2003; Heckert et al. 2005a, b; Parker et al. 2006; Spielmann et al. 2007a, b; Nesbitt & Stocker 2008), the principal Revueltian tetrapod assemblages are those of the Newark Supergroup of eastern North America of Neshanician and Cliftonian (part) age; Ørsted Dal Member of the Fleming Fjord Formation, Greenland; Stubensandstein (Lo¨wenstein Formation) of the German Keuper; Calcare di Zorzino (Zorzino Limestone) and Dolomia di Forni (Forni Dolomite), northern Italy; and lower part of Dharmaran Formation, India. In West Texas-eastern New Mexico, the Bull Canyon Formation of the Chinle Group yields extensive assemblages of Revueltian tetrapods, including the characteristic tetrapod assemblage (e.g. Hunt 2001; Lehman & Chatterjee 2005). In the Chama basin of north–central New Mexico, the Petrified Forest Formation of the Chinle Group also yields Revueltian tetrapods, especially from the Snyder and Canjilon phytosaur-dominated bonebeds (Zeigler et al. 2003; Heckert et al. 2005b; Nesbitt & Stocker 2008). In northern Arizona, two Chinle Group units, the Painted Desert Member of the Petrified Forest Formation and the overlying Owl Rock Formation, have produced numerous Revueltian fossils, especially from the Petrified Forest National Park and from localities on Ward’s Terrace north of Flagstaff (e.g. Kirby 1989, 1991, 1993; Heckert et al. 2005a; Spielmann et al. 2007a). In eastern North America, the provincial Neshanician LVF is based on a limited fossil assemblage typified by the aetosaur Aetosaurus arcuatus (Lucas et al.1998a; Lucas & Huber 2003). This taxon is present in ‘Lithofacies Association II’ of the Chatham Group (Durham sub-basin of the Deep River basin), the Newark Basin (range zone: Warford through Neshanic Members of the lower Passaic Formation), and the middle New Haven Arkose of central Connecticut. Other vertebrates from the Neshanician LVF include indeterminate metoposaurid and phytosaur teeth, skull and scute fragments (e.g. ‘Belodon validus’), a rauisuchian, crocodylomorph, traversodontid and a sphenodontid (lower New Haven Arkose) as well as a dominance of the primitive neopterygian Semionotus sp. over other fish taxa, a trend also apparent in ageequivalent strata of the Chinle Group and German Keuper (Huber et al. 1993c; Lucas & Huber 2003).

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The Cliftonian LVF is based on a low-diversity assemblage defined by the distribution of the procolophonid Hypsognathus fenneri. This taxon is common in the type area, from the middle (?Mettlars Member) to the upper (?Member TT) Passaic Formation of the northern Newark basin (e.g. Baird 1986). It is also known from the upper New Haven Arkose of the Hartford basin, central Connecticut, and the basal Blomidon Formation in the Fundy basin, Nova Scotia (Sues et al. 1997). The Fundy basin specimen of Hypsognathus was obtained from pebble conglomerate at the base of the Blomidon Formation, which unconformably overlies the Wolfville Formation. The only other vertebrates that occur in the interval of Cliftonian age are indeterminate phytosaur remains (including the holotype of ‘Clepsysaurus pennsylvanicus’ Lea 1851) from the Ukrainian Member of the Passaic Formation in the Newark basin, moderately diverse tetrapod footprint assemblages at many horizons in the Passaic Formation (e.g. Szajna & Silvestri 1996; Lucas & Sullivan 2006), and an indeterminate sphenodontid from the upper New Haven Arkose (Olsen 1980; Sues & Baird 1993; Lucas & Huber 2003). The Malmos Klint and overlying Ørsted Dal Members of the Fleming Fjord Formation in eastern Greenland yield fossil tetrapods of Revueltian age (Jenkins et al. 1994, 1997, 2008). The Malmos Klint Member has produced fragmentary fossils of plagiosaurid amphibians, the amphibian Cyclotosaurus, possible phytosaur fragments and the prosauropod dinosaur Plateosaurus. The Ørsted Dal Member assemblage is much more diverse: the amphibians Gerrothorax and Cyclotosaurus, the turtle cf. Proganochelys, unidentified sphenodontians, the aetosaurs Aetosaurus and Paratypothorax, the pterosaur Eudimorphodon, the prosauropod dinosaur ‘Plateosaurus’, a theropod dinosaur, theropod dinosaur footprints (Grallator), and the mammals Kuehneotherium, cf. ?Brachyzostrodon and Haramiyavia. As Jenkins et al. (1994) argued, this assemblage shares many taxa with the German Stubensandstein. More specifically, other than Plateosaurus, most taxa from the Ørsted Dal Member are known in the Lower Stubensandstein, to which I correlate the Greenland assemblage. In Germany, the best known and most diverse Keuper tetrapod assemblage is that of the Lower Stubensandstein (Lo¨wenstein Formation). This assemblage includes the amphibians Cyclotosaurus and Gerrothorax, the earliest European turtles (Proganochelys and Proterochersis), Pseudopalatusgrade phytosaurs (Nicrosaurus), the aetosaurs Aetosaurus and Paratypothorax, rauisuchians (Teratosaurus), theropod dinosaurs, and the prosauropod dinosaurs Sellosaurus and Thecodontosaurus (e.g. Benton 1993; Hungerbu¨hler 1998; Lucas 1999; Schoch & Werneburg 1999; Schoch 2007). The

phytosaurs, aetosaurs, and rauisuchians provide a strong basis for assigning a Revueltian age to the Lower Stubensandstein (Lucas & Hunt 1993a; Hunt 1994; Lucas 1999). The younger, Middle and Upper Stubensandstein, produce a similar, but less diverse assemblage, so I also assign them a Revueltian age. Whether or not the lowest occurrence of Mystriosuchus in the Middle Stubensandstein is of biochronologic significance is not clear. The assemblages of the Upper Stubensandstein and Knollenmergel (Tro¨ssingen Formation) are almost entirely dinosaurian – 95% or more of the fossils are of dinosaurs (Benton 1986, 1991). This contrasts sharply with the Lower and Middle Stubensandstein assemblages, in which dinosaurs are a much smaller percentage of the fossils collected. However, I regard this change to dinosaur domination as largely a local facies/taphonomic effect, not a biochronologically significant event (Hunt 1991). It seems likely, but not certain, that the Knollenmergel assemblage is of Apachean age (see below). In the Lombardian Alps of northern Italy, after the regional progradation of platform carbonates during the early-middle Norian (Dolomia Principale), extensional tectonism produced intraplatform depressions occupied by patch reefs, turbiditic debris flows and lagoonal to freshwater facies (Jadoul 1985; Jadoul et al. 1994). Tetrapods from these intraplatform strata, the Zorzino Limestone at the Cene and Endenna quarries in Lombardy, are the diapsids Endennasaurus and Vallesaurus, the prolacertiform Longobardisaurus, the rhynchocephalian Diphydontosaurus, the drepanosaurids Drepanosaurus and Megalancosaurus, the phytosaur Mystriosuchus, the aetosaur Aetosaurus, the pterosaurs Eudimorphodon and Peteinosaurus and the placodont Psephoderma (e.g. Wild 1989; Pinna 1993; Renesto 2006). In Germany, Mystriosuchus is well known from the Middle Stubensandstein and Aetosaurus from the Lower-Middle Stubensandstein, so a Revueltian age of the Zorzino Limestone is certain. The Calcare di Zorzino also crops out in Austria, where it yields specimens of Langobardisaurus and the pterosaur Austriadactylus, a likely synonym of Preondactylus (Dalla Vecchia 2009; S. Renesto, written commun. 2009). Also, in Austria, unpublished specimens of Mystriosuchus are known from Totes Gebirge (possibly Dachstein) (S. Renesto, written commun. 2009). The other Italian Late Triassic tetrapod sites are in the Forni Dolomite (Dolomia di Forni) in the Veneto Prealps of northeastern Italy. They yield the drepanosaurids Drepanosaurus and Megalancosaurus, and the pterosaurs Eudimorphodon and Preondactylus (Dalla Vecchia 1995) and a specimen of Langobardisaurus under study by S. Renesto (written commun. 2009). The presence of Eudimorphodon supports a Revueltian age assignment.

TRIASSIC TETRAPODS

Upper Triassic tetrapod assemblages from the Indian Subcontinent come from the Pranhita– Godavari Valley of south–central India. Several summaries (Jain et al. 1964; Kutty 1969; Kutty & Roychowdhury 1970; Sengupta 1970; Jain & Roychowdhury 1987; Yadagiri & Rao 1987; Kutty et al. 1988; Kutty & Sengupta 1989; Bandyopadhyay & Roychowdhury 1996; Bandyopadhyay & Sengupta 2006) have been published, but other than the lower Maleri assemblage (see above), relatively few of the fossils have been adequately documented in print, forcing me to rely largely on unsubstantiated genus-level identifications to evaluate the ages of the tetrapod assemblages. A case in point is the Dharmaram Formation, which yields two stratigraphically discrete vertebrate fossil assemblages (lower and upper). The stratigraphic range of the lower assemblage has not been published, and it includes a phytosaur that Kutty & Sengupta (1989, table 2) list as Nicrosaurus, aetosaurs, including a so-called ‘Paratypothorax-like’ form, and prosauropod dinosaurs. Based primarily on the supposed Nicrosaurus record, I consider the lower assemblage of the Dharmaram Formation a possible Revueltian correlative. Comments. Hunt & Lucas (1993c) suggested that, perhaps along the lines of the CliftonianNeshanician subdivision used in the Newark Supergroup, the Revueltian merits subdivision, and Hunt (1994, 2001) subdivided it into three sub-LVFs of regional utility. Two of these, the Barrancan (early Revueltian) and Lucianoan (later Revueltian) are readily correlated in the western USA using various index fossils (e.g. Heckert & Lucas 2006). Some of the discussion of the Revueltian has focused on whether or not it is readily distinguished from the next younger Apachean LVF (Long & Murry 1995; Rayfield et al. 2005, 2009). These discussions are rooted in taxonomic arguments, as the type assemblages of the Revueltian and Apachean are stratigraphically superposed in east –central New Mexico, USA and thus are obviously time successive. Typothorax, Aetosaurus and Pseudopalatusgrade phytosaurs were listed as Revueltian index fossils by Lucas (1998a). However, recognition of an older, Adamanian species of Typothorax, T. antiquum, by Lucas et al. (2002b) has modified this; it is the species T. coccinarum that is a Revueltian index fossil, and this is part of what prompted Hunt et al. (2005) to redefine the beginning of the Revueltian as the FAD of T. coccinarum, a decision followed by Lucas et al. (2007e) and also used here. Typothorax coccinarum stands as a robust index fossil of the Revueltian across the Chinle Group. Indeed, its likely descent from T. antiquum as part of an anagenetic evolutionary lineage (Lucas et al. 2002b) is significant to the Triassic tetrapod

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biochronology in that the beginning of a LVF can be defined by a true species-level evolutionary event, not the appearance of a genus-level taxon. Aetosaurus is one of the most robust index fossils of the Triassic tetrapod timescale (Fig. 9). Lucas et al. (1998b) presented a detailed taxonomic revision based on study of all North American and European specimens. Aetosaurus has a marine record in the middle Norian of northern Italy (Wild 1989), and all of its nonmarine records are Revueltian. Criticism of the use of Aetosaurus, typified by Rayfield et al. (2005, 2009), is based on the claim that because Aetosaurus has been portrayed as the plesiomorphic sister taxon of other aetosaurs in cladistic analyses (e.g. Heckert & Lucas 2000) it ‘must’ have a long ghost lineage that therefore renders it useless in biostratigraphy. I regard this as specious cladotaxonomic reasoning (Lucas et al. 1999a, 2007c, e). Thus, the position of a taxon on a cladogram has nothing to do with its biostratigraphic utility unless all the assumptions of the cladogram – and the existence of a ghost lineage is nothing more than an assumption – are brought into the biostratigraphic analysis. Indeed, an alternative interpretation of the Heckert & Lucas (2000) cladogram of aetosaurs, one that views Aetosaurus as a highly derived, dwarfed and simplified form, would produce a very different ‘ghost lineage’. Aetosaurus thus is a taxonomically stable and robust Revueltian index fossil (e.g. Fraas 1877; Huene 1921; Walker 1961; Wild 1989; Parrish 1994; Heckert et al. 1996, 2007a; Heckert & Lucas 1998; Small 1998; Lucas et al. 1998b, 1999a; Heckert & Lucas 2000; Parker 2007). Pseudopalatus-grade phytosaurs include Pseudopalatus, Nicrosaurus and Mystriosuchus, all taxa restricted to Revueltian time. Like the use of Rutiodon-grade phytosaurs to identify the Adamanian, this is a convenient and concise way to refer to a group of broadly contemporaneous phytosaur taxa whose stratigraphic ranges are well established, but whose genus- and species-level nomenclature remain in flux (compare, e.g. the differing phytosaur taxonomies of Ballew 1989; Hunt 1994; Long & Murry 1995; and Hungerbu¨hler 2002). Heckert & Lucas (1997) suggested that Revueltosaurus might serve as an index taxon of Revueltian time. At that time Revueltosaurus, which was known solely from teeth, was considered to be an ornithischian dinosaur. Parker et al. (2005) documented associated skulls and postcrania of Revueltosaurus callenderi, demonstrating that that taxon is actually a crurotarsan archosaur. However, they noted that, following Hunt (1989), Padian (1990) and others, the teeth are indeed diagnostic, and the taxon is valid. Heckert & Lucas (2006) then showed that in the Chinle Group Revuletosaurus is restricted to strata of Revueltian age.

Newark Supergroup, eastern USA Newark Basin (NJ)

Deep River Basin (NC)

Juvavites magnus

New Haven Arkose

K I Graters Member

Malayites dawsoni

ET

Sitkinoceras kerri

Warford Member

Aralalta Group

Lithofacies Association III

Bull Canyon Formation (part)

Lithofacies Association II

Trujillo Formation C

Biochron

Zorzino Limestone

middle

Passaic Formation (part)

Aetosaurus arcuatus (holotype)

Middle (Alaunian)

Perkasie Member

Italy (Lombardian Alps)

Germany

Neshanic Member

LM

Early (Lacian)

NORIAN

Drepanites rutherfordi

Eastern Greenland

Rock Point Formation

Q Himavatites columbianus

New MexicoColorado USA

Orsted / Dal Member

Steinmergel Keuper

Hartford Basin (CT)

Stubensandstein

ammonoid zones

Fleming Fjord Formation

substage

Chinle Group (part)

STAGE

S. G. LUCAS

Dolomia Principale

Aetosaurus biochron

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lower

Malmros Klint Member

Aetosaurus occurrences Fig. 9. Global correlation of Aetosaurus localities, which identify an Aetosaurus biochron of Revueltian age.

This demonstrates the irrelevance of the assumed position of a taxon in a phylogeny to biostratigraphy. The changing phylogenetic position of Revueltosaurus alters neither its biostratigraphic significance nor its biochronological utility. What is biostratigraphically important about Revueltosaurus is that it is distinctive (easily identified), relatively common and/or widespread, and known from a restricted stratigraphic interval. Whether it is an ornithischian (as previously supposed) or a crurotarsan (the current phylogenetic hypothesis) is irrelevant to its biostratigraphic and biochronological utility.

Apachean Definition. The Apachean LVF is the time interval between the FAD of the phytosaur Redondasaurus and the FAD of the crocodylomorph Protosuchus (Fig. 1). Lucas & Hunt (1993a) introduced the term Apachean LVF to refer to the time equivalent to the vertebrate fossil assemblage of the Redonda Formation (Chinle Group) in east–central New Mexico, USA (Lucas et al. 1985; Hunt 1994; Hunt & Lucas 1997; Lucas 1998a; Lucas et al. 2001; Spielmann et al. 2006a, b). Apachean time begins

with the FAD of the phytosaur Redondasaurus. The end of Apachean time is the beginning of the Wassonian LVF, which is the FAD of the crocodylomorph Protosuchus (Lucas & Huber 2003; Lucas & Tanner 2007a, b). Characteristic tetrapod assemblage. The characteristic tetrapod assemblage of the Apachean LVF is from the Redonda Formation of the Chinle Group in Guadalupe and Quay Counties, New Mexico, USA. The following taxa are present: the amphibian Apachesaurus, a sphenodontid, a procolophonid, the phytosaur Redondasaurus, the aetosaur Redondasuchus, the rauisuchian Redondavenator, the sphenosuchian Vancleavea, a rauisuchian, theropod dinosaurs and a ?cynodont (e.g. Hunt 1994; Hunt & Lucas 1993b, 1997; Heckert et al. 2001; Hunt et al. 2005; Spielmann et al. 2006a, b). Index fossils. The following tetrapod genera are restricted to Apachean time and are widespread and/or common enough to be useful as index fossils: the phytosaur Redondasaurus, the aetosaur Redondasuchus and the dinosaur Riojasaurus. Principal correlatives. Principal correlatives of the type Apachean assemblage are the Whitaker

TRIASSIC TETRAPODS

quarry in the Rock Point Formation of the Chinle Group at Ghost Ranch, New Mexico, the Cliftonian LVF assemblages (in part) of the Newark Supergroup, the Knollenmergel (Tro¨ssingen Formation), time-equivalent upper Arnstadt Formation and the ‘Rhaetian Bonebed’ of the Germanic Basin, the Coloradan LVF of Argentina and the tetrapod assemblage of the Lower Elliot Formation in South Africa. Some of the fissure-fill assemblages in the uppermost Mercia Mudstone Group and/or lowermost Penarth Group of the United Kingdom (Fraser 1994; Benton & Spencer 1995; Whiteside & Marshall 2008) may be Apachean correlatives. Some of the so-called Rhaetian vertebrate sites in France, such as Saint-Nicolas-de-Port, may be Apachean correlatives as well (Lucas & Huber 2003). At Ghost Ranch in New Mexico, the Whitaker quarry bone bed is dominated by skeletons of the theropod dinosaur Coelophysis bauri (Colbert 1989). Nevertheless, it also includes the sphenodont Whitakersaurus, at least one drepanosaur, a rauisuchian skeleton (cf. Postosuchus), the sphenosuchians Hesperosuchus and Vancleavea, the chatterjeeid Shuvosaurus (¼ Effigia) and the phytosaur Redondasaurus (e.g. Hunt & Lucas 1993b; Clark et al. 2000; Harris & Downs 2002; Hungerbu¨hler 2002; Hunt et al. 2002; Lucas et al. 2003; Rinehart et al. 2004; Nesbitt 2007; Lucas et al. 2005, 2007e; Heckert et al. 2008; Renesto et al. 2009). In Argentina, the Los Colorados Formation consists of siliciclastic red beds approximately 800 m thick. Near its base, a single tetrapod fossil – a dicynodont skull, the holotype of ‘Jachaleria’ colorata Bonaparte 1970 – was collected. The remainder of the tetrapod fossils from the Los Colorados Formation are from its middle and upper parts but have not been stratigraphically organized. The assemblage includes the turtle Palaeochersis, the ornithosuchid Riojasuchus, the aetosaur Neoaetosauroides, the rauisuchid Fasolasuchus, the crocodylomorphs Hemiprotosuchus and Pseudhesperosuchus, the prosauropod dinosaurs Riojasaurus and Coloradisaurus, the theropod dinosaur Zupaysaurus and the tritheledontid cynodont Chaliminia (e.g. Bonaparte 1970, 1971, 1978, 1980, 1997; Lucas & Hunt 1994; Rougier et al. 1995; Arcucci et al. 2004). The correlative Quebrada del Barro and El Tranquilo formations also produce prosauropods (e.g. Riojasaurus, ‘Mussaurus’) (Casamiquela 1980; Bonaparte & Vince 1979; Bonaparte & Pumares 1995). The Los Colorados assemblage clearly is of Late Triassic age (Arcucci et al. 2004) and must be post-Ischigualastian. However, its endemism makes it difficult to correlate precisely. I tentatively consider it an Apachean correlative based primarily on its abundant prosauropods. The age of the tetrapod assemblage from the Lower Elliott Formation in South Africa has long

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been considered Late Triassic. Lucas & Hancox (2001) reviewed the age of this assemblage, which is dominated by sauropodomorph dinosaurs, but also has rare amphibians (a large chigutisaurid), a possible rauisuchian (Basutodon), the ornithischian dinosaur Eocursor, a traversodontid (Scalenodontoides) and the characteristic Late Triassic footprint ichnogenus Brachychirotherium (Kitching & Raath 1984; Lucas & Hancox 2001; Butler et al. 2007). This is the ‘Euskelosaurus range zone’ of Kitching & Raath (1984), the youngest Triassic tetrapod assemblage in the Karoo basin. Yates (2003) re-evaluated the prosauropods of the Lower Elliott Formation and concluded that most are indeterminate sauropodomorphs or basal sauropods. He noted similarities of indeterminate prosauropods from the Lower Elliott Formation to Riojasaurus from the Los Colorados Formation of Argentina, and similarities between the basal sauropod Antetonitrus from South Africa and Lessemsaurus from Argentina (Yates & Kitching 2003). These conclusions suggest a Lower Elliott –Los Colorados correlation, and thus a tentative Apachean age assignment. In the United Kingdom, fissure fills such as Durdham Down in Clifton yield fossils that include phytosaurs, aetosaurs, dinosauriforms and dinosaurs (e.g. Fraser 1994; Fraser et al. 2002; Galton 2005, 2007a, b; Whiteside & Marshall 2008). Unfortunately, other than a tentative record of Aetosaurus based on a single osteoderm (Lucas et al. 1999b), the fissure fill tetrapods are mostly endemic taxa of no biochronological significance or cosmopolitan taxa with long age ranges, such as the sphenodontian Clevosaurus. Recently, Whiteside & Marshall (2008), based primarily on the palynoflora, assigned the Tytherington fissure fill a Rhaetian age, and extrapolated this age to the other fissures. If this Rhaetian age is correct, then the fissure fill tetrapods are of Apachean age. However, as Lucas & Hunt (1994, p. 340) noted, ‘a single age should not necessarily be assigned to the fossils from one fissure and . . . . individual fossils from the fissures may range in age from middle Carnian to Sinemurian’. Therefore, I continue to regard as problematic the precise age of the Triassic tetrapod assemblages from the British fissure fills. Comments. The Apachean is the most difficult Triassic LVF to correlate globally. This almost certainly reflects a provincialization of the global tetrapod fauna near the end of the Triassic. Some of the apparent endemism of Apachean land-vertebrate assemblages may also be due to facies, sampling and taphonomic biases. Thus, rather than recognize a global Apachean LVF, it may ultimately be necessary to recognize two or more provincial LVFs during this time interval.

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There is no evidence that any part of the Apachean is of Jurassic age. The FAD of the crocodylomorph Protosuchus, which defines the beginning of the next LVF, the Wassonian, appears to correspond closely to the beginning of the Jurassic (Lucas & Tanner 2007a, b). Thus, Protosuchus occurs in units assigned an Early Jurassic based on diverse evidence: the McCoy Brook Formation (Newark Supergroup), the upper Stormberg Group of South Africa and the upper part of the Dinosaur Canyon Member of the Moenave Formation in Utah-Arizona (Colbert & Mook 1951; Sues et al. 1996; Lucas et al. 2005; Lucas & Tanner 2007a, b). The Moenave record of Protosuchus is stratigraphically superposed above Apachean body fossil assemblages of the uppermost Chinle Group (Lucas et al. 1997b, 2005; Lucas & Tanner 2007a, b). Furthermore, it is correlative to the Lower Jurassic conchostracan assemblages from the Whitmore Point Member of the Moenave Formation (Lucas & Tanner 2007a; Kozur & Weems 2010). Relatively recent recognition that Apachean-age strata extend above the Chinle Group into part of the Moenave–Wingate (lower Glen Canyon Group) lithosome has been based, in part, on the occurrence of a Redondasaurus skull in the lower part of the Wingate Sandstone in southeastern Utah (Lucas et al. 1997b; Lucas & Tanner 2007a, b). (Note that Spielmann et al. 2007a, fig. 8A–B illustrated a cast of this skull and mistakenly attributed it to the Revueltian Owl Rock Formation). Lucas (1998a) listed three Apachean index fossils: the aetosaur Redondasuchus, the phytosaur Redondasaurus and the dinosaur Riojasaurus. Riojasaurus is known from Argentina and may be present in the Lower Elliott Formation in South Africa. The Apachean is readily distinguished in North America by its primary index fossils, Redondasaurus and Redondasuchus. However, some workers (Long & Murry 1995; Martz 2002) have questioned the validity of Redondasaurus and Redondasuchus, proclaiming the former a synonym of Pseudopalatus and the latter a synonym of Typothorax. Nevertheless, Heckert et al. (2001) and Spielmann et al. (2006a, b) reaffirmed the distinctiveness and validity of Redondasuchus and Redondasaurus, and Hungerbu¨hler (2002) also recognized Redondasaurus as distinct from Pseudopalatus.

Correlation of the LVFs to the Marine SGCS Introduction Records of nonmarine Triassic tetrapods in marine strata (Lucas & Heckert 2000), palynostratigraphy, magnetostratigraphy and radioisotopic ages provide

some basis for correlation of the LVFs to the standard global chronostratigraphic scale (Fig. 10). Nevertheless, reliable data for this correlation remain relatively sparse, so the correlation of the LVFs to the SGCS is still imprecise in many time intervals.

Lootsbergian The base of the Triassic (¼ Permo-Triassic boundary [PTB], ¼ base of Induan Stage) has been formally defined by the FAD of the conodont Hindeodus parvus at a global stratotype section and point (GSSP) located at Meishan in southern China (Yin et al. 2001). This means it is possible to attempt to correlate a potential Triassic base in the nonmarine section to a fixed, agreed-upon point in the marine timescale. It is important to ask how the Lootsbergian correlates to the marine PTB in order to establish the synchrony or diachroneity of marine and nonmarine events across the PTB. However, such correlation is not simple, because no sections are known where strata bearing terrestrial tetrapods can be directly correlated (say by interfingering lithostratigraphy) to the marine record across the PTB. Thus, magnetostratigraphy, isotope stratigraphy, conchostracan biostratigraphy and palynostratigraphy have been used to correlate the nonmarine and marine records across the PTB. Lucas (2009) provides a detailed discussion of this correlation, which is briefly reviewed here. There is a well documented negative d13C excursion at the PTB in marine sections that closely coincides with the major extinction that precedes the formally-defined PTB (e.g. Payne et al. 2004; Yin et al. 2005, 2007; Richoz 2006). Diverse analyses indicate that the marine PTB is within the lower third of a long normal-polarity chron (e.g. Ogg 2004; Steiner 2006; Hounslow & Muttoni 2010). Palynostratigraphy has also been used by some to correlate marine to nonmarine sections at the PTB (e.g. Morante 1996; Looy et al. 1999, 2001; Twitchett et al. 2001; Collinson et al. 2006), particularly the fungal abundance spike documented in marine and nonmarine sections that some have considered to correspond to the PTB marine mass extinction (e.g. Eshet et al. 1995; Visscher et al. 1996; Steiner et al. 2003). In the conchostracan biostratigraphy, which is well correlated with the marine scale, the PTB coincides with the boundary between the Falsisca postera Zone and the Falsisca verchojanica Zone (Kozur 1998a, b, 1999; Bachmann & Kozur 2004; Kozur & Weems 2010). As in the marine section, this conchostracan zonal boundary lies in the lower third of a long normal magnetostratigaphic zone that straddles the PTB, and it is characterized by a minimum in d13C in continental beds (Bachmann &

TRIASSIC TETRAPODS

J

Hettangian

201

Rhaetian

Norian

210

475

Wassonian Apachean

Revueltian

Aetosaurus, Drepanosaurus, Edennasaurus, Eudimorphodon, Macrocnemus, Megalancosaurus, Mystriosuchus,Peteinosaurus, Preonodactylus, Sikkanisuchus

Late

220 Adamanian

230

TRIASSIC

Carnian

Otischalkian

Berdyankian Ladinian

Metoposaurus, Parasuchus

cf. Cyclotosaurus, aff. Dinodontosaurus, Mastodonsaurus

250 252

Early

Middle

240

Anisian

Perovkan

Olenekian

Nonesian

Induan

Lootsbergian Platbergian

Permian Changxingian

Eocyclotosaurus, Lotosaurus Aphaneramma, Lyrocephaliscus, Parotosuchus, Peltostega, Sasenisaurus, Tertrema Luzocephalus, Stoschiosaurus, Tupilakosaurus, Wetlugasaurus

Fig. 10. Marine records of Triassic nonmarine tetrapod correlated to the marine SGCS and the Triassic land– vertebrate faunachrons. See Lucas & Heckert (2000) for details. Restoration of Eocyclotosaurus by Matt Celeskey.

Kozur 2004; Korte & Kozur 2005b). This minimum in d13C occurs in continental lakes without facies changes, and the conchostracan boundary occurs in the Dalongkou section in northwestern China close to the HO of Dicynodon (Kozur 1998a, b; Metcalfe et al. 2009; Kozur & Weems 2010).

At the Meishan section in southern China, a sharp drop in d13C values coincides with the maximum amount of marine extinction, and this mass extinction and carbon isotope excursion are older than the PTB defined by the lowest occurrence of the conodont Hindeodus parvus (Fig. 11).

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Fig. 11. Magnetostratigraphic correlation of marine PTB section at Meishan, China (based on Yin et al. 2005) to PTB tetrapod extinction interval in Karoo basin of South Africa (based on Ward et al. 2005). The sections are correlated at the base of the normal polarity magnetozone that encompasses the PTB. However, they are not scaled to each other based on time intervals or section thickness, so the only certain point of correlation indicated is the base of the normal polarity magnetozone that encompasses the PTB. From Lucas (2009). Restoration of Lystrosaurus by Matt Celeskey.

However, in sections without a weathered boundary clay (e.g. Shahreza, Iran, and Gerenava´r, Bu¨kk Mts, Hungary) or without a boundary clay, the d13C minimum lies around the FAD of H. parvus, at the PTB (Korte & Kozur 2005a). In the Karoo basin of South Africa, d13C isotope data through the PTB have been used to correlate to the marine d13C excursion (MacLeod et al. 2000; Ward et al. 2005). However, these isotope data do not convincingly support the conclusion that the highest occurrence of Dicynodon in the Karoo basin is equivalent to the PTB. Indeed, Tabor et al. (2007) recently published an analysis of d13C across the PTB in the Karoo basin and argued that changes in that record are driven by local facies changes and are not a reflection of atmospheric carbon values. Therefore, the d13C record in the Karoo basin cannot be reliably correlated to the d13C record in marine strata across the PTB. However, at Dalongkou in northwestern China the HO of Dicynodon is close to the PTB defined and correlated to the marine scale by conchostracans. Therefore, the minimum in d13C in the Karoo basin may be a primary signal, and the HO of Dicynodon is close to the PTB.

The PTB marine extinction took place during a relatively long interval of normal magnetic polarity that straddles the PTB, well documented in a variety of marine sections (Ogg 2004; Steiner 2006) as well as in continental sections (Szurlies 2004; Bachmann & Kozur 2004). In the Karoo basin there is an interval of normal polarity that encompasses the highest occurrence of Dicynodon and is part of the stratigraphically thick (c. 60 m) interval of low d13C values (Schwindt et al. 2003; De Kock & Kirschvink 2004; Ward et al. 2005; Steiner 2006) (Fig. 11). These magnetostratigraphic data indicate that the lowest occurrence of Lystrosaurus (in an interval of reversed polarity) is older than the PTB (as already suggested by King & Jenkins 1997; Kozur 1998a, b; and Botha & Smith 2007, among others), and that the highest occurrence of Dicynodon is closer to the PTB (Fig. 11). I favour the magnetostratigraphic correlation of the Meishan and Karoo sections, and it is consistent with all other available correlation data. The correlation indicates that in the Karoo basin the base of the Lootsbergian (LO of Lystrosaurus) pre-dates the main marine extinction event. The LO of

TRIASSIC TETRAPODS

Lystrosaurus cannot be used to place the PTB in nonmarine sections; the highest occurrence of Dicynodon is a much better proxy for the PTB. Thus, the earliest Lootsbergian is of latest Permian (Changshingian) age. Correlation of the rest of the Lootsbergian to at least part of the marine Induan Stage is clear (Lucas 1998a; Lucas et al. 2007e). However, whether the Lootsbergian equates to part, all or more than Induan time is not possible to determine with the available data. The Wordy Creek Formation in eastern Greenland has a record of Lootsbergian amphibians interbedded with marine late Griesbachian –early Dienerian (middle Induan) age strata. Thus, the stratigraphically lowest record of Luzocephalus here is in the Ophiceras commune ammonite zone, and the genus extends up through the ‘Proptychites rosenkrantzi Zone’. Most of the Wordy Creek Formation amphibians come from the younger ‘Anodontophora fassarensis beds’, which are the youngest Lower Triassic strata in this section (Nielsen 1935; Sa¨ve-So¨derbergh 1935). This indicates a range of Luzocephalus from late Griesbachian through early Dienerian (middle Induan), but the other temnospondyl taxa are of middle or late Dienerian (late Induan) age (Tru¨mpy 1961; Silberling & Tozer 1968; Tozer 1994). Luzocephalus, Tupilakosaurus, and Wetlugasaurus occur in the Vokhmian Horizon of the Vetlugian Series of the Russian Urals. This fauna includes Lystrosaurus, an index taxon of the Lootsbergian land–vertebrate faunachron, so the amphibian records from Greenland establish a middle– late Induan age for at least part of Lootsbergian time. In northwestern Madagascar the upper part of the marine Andavakoera Formation (Dienerian) yields a diverse assemblage of temnospondyls: ?Benthosuchus, ?Wetlugasaurus, Mahavisaurus, Aphaneramma, Ifasaurus, Tertrema, Tertremoides, Trematosaurus, Wantzosaurus and Deltacephalus (Swinton 1956; Lehman 1961, 1966, 1979). The Benthosuchus and Wetlugasaurus identifications are not reliable (Cosgriff 1984; Damiani 2001), though the amphibians from the Andavakoera Formation may be of Lootsbergian age. This may indicate correlation of part of the Lootsbergian and the Dienerian. Shishkin (2000, p. 65) asserted that the Lootsbergian includes assemblages younger than Induan, but no credible data support his claim. For example, he stated (p. 65) that ‘the Hesshanggou assemblage of China [which Lucas 1998a assigned a Lootsbergian age] . . . is actually latest Spathian or Spathian– Anisian in age’. However, there is no direct way to correlate Hesshanggou Formation red beds in Shanxi (long correlated by Chinese workers to the ‘Procolophon zone’ of the Karoo: Cheng 1981) to the SGCS (Lucas 1993a, 1998a, 2001; Lucas et al. 2007e). In another case, Damiani et al. (2000)

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reported a generically-indeterminate trematosaurid jaw from the South African Lootsbergian strata and claimed it extends Lootsbergian time up to the late Olenekian, largely because of its resemblance to Olenekian Trematosaurus. An equally likely possibility is that Damiani et al. (2000) simply documented an Induan-age trematosaurid. Thus, the possibility exists that Lootsbergian time is as young as early Olenekian, but no reliable data are known to support a Lootsbergian –Olenekian correlation.

Nonesian Cross correlation of the Nonesian to at least part of the Olenekian is clear because of the occurrence of the Nonesian index temnospondyl Parotosuchus in marine upper Olenekian (Spathian) strata in the Mangyshlak Peninsula of western Kazakstan. Thus, from Mangyshlak, Lozovsky & Shishkin (1974) documented Parotosuchus sequester from marine upper Olenekian (Spathian) strata that yield Tirolites and other ammonites. Parotosuchus is an index taxon of Nonesian time, and the Kazak record thus provides a direct Nonesian-late Olenekian correlation. Furthermore, a Spathian conchostracan fauna of the Germanic Basin in the Hardegsen Formation (with Parotosuchus) is well correlated with marine beds in Hungary (with Spathian ammonoids) and northern Siberia (Kozur & Weems 2010). In the western USA, the Nonesian Torrey Formation of the Moenkopi Group overlies the early Olenekian (Smithian) ammonite-bearing Sinbad Formation, whereas the Nonesian Wupatki Member of the Moenkopi Formation is clearly younger than the late Olenekian (Spathian) Virgin Limestone (e.g. McKee 1954; Blakey 1974; Morales 1987; Steiner et al. 1993; Lucas & Schoch 2002; Goodspeed & Lucas 2007; Lucas et al. 2007a). This suggests a Smithian–Spathian (Olenekian) age for the Moenkopi Nonesian tetrapods, and supports a broad Nonesian –Olenekian correlation. The Sticky Keep Formation in Svalbard yields amphibians that co-occur here with early Olenekian (Smithian) ammonites (Buchanen et al. 1965; Tozer 1967). The amphibians are: Sasenisaurus, Peltostega, Aphaneramma, Lyrocephaliscus, Teretrema and Boreaosaurus (Wiman 1910, 1915; Nilsson 1942, 1943; Cox & Smith 1973). These trematosaurs are believed to have been euryhaline amphibians that may have actually lived in marine environments. They also reflect a high diversity and abundance of trematosaurs characteristic of the Nonesian. However, the Svalbard trematosaur taxa are mostly endemic and thus only provide stage-of-evolution evidence for an Olenekian– Nonesian cross-correlation.

478

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Perovkan A fairly direct correlation can be made of some Perovkan tetrapod assemblages to the SGCS (Lucas & Schoch 2002). Thus, strata of the Ro¨t Formation (Upper Buntsandstein) in southwestern Germany–eastern France are lower Anisian marginal marine to interbedded nonmarine/marine facies of well-established age because of their close physical relationship to the Lower Muschelkalk. Indeed, marine facies of the lower Ro¨t contain early Anisian conodonts, the early Anisian (Aegean) ammonoid Beneckeia tenuis and agediagnostic holothurian sclerites (e.g. Kozur 1993), and magnetostratigraphic correlation of the Ro¨t Formation to marine magnetostratigraphy indicates an early Anisian age (Szurlies 2007; Hounslow et al. 2008). Furthermore, conchostracans of the Ro¨t correlate with Aegean and lower Bithynian marine intercalations (Kozur & Weems 2010). The common amphibian from the Ro¨t Formation, Eocyclotosaurus, is an index taxon of the Perovkan found in both Europe and the western United States (e.g. Ortlam 1970; Morales 1987; Lucas & Schoch 2002). The Ro¨t records of Eocyclotosaurus thus provide a Perovkan –early Anisian correlation. The Gogolin Formation (lowermost Muschelkalk) in Polish Silesia yields fragmentary temnospondyl and archosaur fossils that include the types of Mentosaurus waltheri, Eurycervix posthumus, and ‘Xestorrhytias perrini’, all of which are indeterminate mastodonsaurids, and the rauisuchian Zanclodon silesiacus Jaekel, based on a single tooth. Ammonite biostratigraphy places the Gogolin Formation in the lower Anisian (e.g. Kaim & Niedz´wiedzki 1999). The tetrapod material, however, is too fragmentary to be of much biochronological utility. Nevertheless, the available material closely resembles some of the tetrapods from the Upper Buntsandstein (Ro¨t Formation) of southwestern Germany–eastern France, and thus supports a Perovkan–early Anisian correlation. Magentostratigraphic correlation of the Perovkan Otter Sandstone in Great Britain indicates it is of late Anisian age (Hounslow & McIntosh 2003). Abdala et al. (2005) assigned the Perovkan Cynognathus zone C in the Karoo basin to the late Anisian based largely on the palynological content of its probable correlatives, such as the Wianamatta Group in the Sydney basin of Australia. Thus, there is good evidence that the Perovkan is equivalent to most of Anisian time.

Berdyankian The German section provides the best data for a Berdyankian– Ladinian correlation. Thus, the Berdyankian taxon Mastodonsaurus giganteus ranges

from the Upper Muschelkalk through the Lettenkeuper, strata of late Ladinian (Longobardian) age (Schoch 1999). The lower Ladinian Partnach Formation of western Austria yielded a temnospondyl jaw fragment that Sander & Meyer (1991) identified as cf. Cyclotosaurus sp. However, this specimen could just as well belong to Mastodonsaurus (cf. Schoch 1999), so it is of limited biochronological significance. Stur (1873) reported ?Mastodonsaurus giganteus from the Lunz Sandstone in the Austrian Alps. This is an early Carnian (Julian) record, broadly correlative to the German Schilfsandstein. However, I have examined the material Stur described, and it is not diagnostic of M. giganteus; it could just as well belong to Cyclotosaurus. Therefore, this record also is of limited biochronologic significance. The Brazilian and Argentinian Dinodontosaurus assemblages are unambiguously correlated to each other, and have generally been considered Ladinian based on flimsy palynostratigraphic evidence (Lucas & Harris 1996; Lucas 2002). Tetrapod evidence to correlate the Dinodontosaurus assemblages to the European Berdyankian is also not robust; it consists of fragmentary remains of Dinodontosaurus-grade and Stahleckeria-grade dicynodonts from the German Muschelkalk and Russian Bukobay Formation, respectively, not on shared alpha taxa (Lucas & Wild 1995; Lucas 1998a; Lucas et al. 2007b). At present, this South American –European correlation remains weakly supported and merits further study. This is one area where magnetostratigraphy (in South America) will be of assistance. Thus, all available robust data for correlating the Berdyankian to the SGCS indicate that it is equivalent to the late Ladinian. This may indicate that there is a global gap equivalent to the early Ladinian in the Triassic tetrapod record.

Otischalkian There are two records of Otischalkian tetrapod index taxa in marine strata in Austria that support an Otischalkian– Carnian correlation: 1. Raibler Schichten, Austria: Koken (1913) described Metoposaurus santaecrucis from a conglomeratic sandstone in the upper part of the Raibler Schichten. This is an early Carnian (Julian) record, and thus correlates part of the Otischalkian (index taxon ¼ Metoposaurus) to the early Carnian. 2. Opponitzer Schichten, Austria: Huene (1939) described a skull fragment of the phytosaur Parasuchus (¼ Francosuchus) from the lower part of the Opponitzer Schichten (Kalk) near Lunz, Austria. The occurrence is of late Carnian (Tuvalian) age (Janoscheck & Matura 1980), but it cannot be tied

TRIASSIC TETRAPODS

precisely to a particular ammonite zone (Hunt & Lucas 1991). In Germany, Otischalkian tetrapods from the Schilfsandstein are as old as early Carnian (late Julian). Palynostratigraphy assigns a late Carnian age to the lower Chinle Group, including the strata of Otischalkian age, and an early Carnian age to the oldest Sanfordian strata of the Newark Supergroup (Litwin et al. 1991, 1993; Cornet 1993). Sequence stratigraphy of the Chinle Group advocated by Lucas (1991, 1993b), Lucas & Marzolf (1993) and Lucas & Huber (1994) assigns lower Chinle Group strata to a single sequence, the Shinarump–Blue Mesa sequence. This sequence can be correlated to a late Carnian marine sequence in Nevada (Lupe & Silberling 1985; Lucas & Huber 1994), and recent studies of detrital zircons are consistent with these correlations (Dickinson & Gehrels 2008; Dickinson et al. 2009). Magnetostratigraphy correlates lower Chinle Group strata to the late Carnian portion of the Newark Supergroup magnetostratigraphy (Kent et al. 1995; Molina-Garza et al. 1996; Muttoni et al. 2004). Therefore, the Otischalkian clearly is Carnian, equivalent to the early Carnian and part of the late Carnian.

Adamanian I have long considered the Adamanian to be of late Carnian age based on palynostratigraphy, sequence stratigraphy and magnetostratigraphy (see references cited above under marine cross-correlation of the Otischalkian). In West Texas, Otischalkian and Adamanian tetrapod assemblages are stratigraphically superposed (Lucas 1993b; Lucas & Anderson 1993a, b, 1994, 1995; Lucas et al. 1993, 1994). Therefore, Adamanian time is younger than the Otischalkian. Revueltian vertebrates are stratigraphically above Adamanian vertebrates in Arizona, New Mexico and Texas. Therefore, Adamanian vertebrates are either the youngest Carnian vertebrates known or the oldest Norian vertebrates known (or both). Kozur & Weems (2007, 2010) discussed at length the biostratigraphic evidence to support a late Carnian (Tuvalian) correlation of the Adamanian. This is the concordance of three biostratigraphies – palynostratigraphy, conchostracan biostratigraphy and vertebrate biostratigraphy – that all indicate that the Adamanian is Tuvalian. Particularly significant is the record in the Newark Supergroup of eastern North America, where, for decades, palynostratigraphy placed the Carnian –Norian boundary at or just above the base of the Passaic Formation (at the Warford Member), a placement supported by conchostracan and tetrapod biostratigraphy (and by megafossil plant biostratigraphy: Ash 1980, 1987) (see summary by Huber et al.

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1993a). Correlations to the Chinle Group based on palynomorphs, conchostracans and tetrapods indicate that the Adamanian LVF is older than the base of the Passaic Formation. Based on counting cycles in the Newark, the estimated age of the Passaic Formation base (and the base of the Norian) is about 217 Ma (Kent & Olsen 1999), but in this counting a complete Rhaetian was assumed. However, according to Kozur & Weems (2005, 2007), most of the Rhaetian is missing in the Passaic Formation, where only the uppermost precession cycle of c. 20 000 years yielded uppermost Rhaetian conchostracans, whereas below these beds late Norian conchostacans are present. Recent correlations of Newark magnetostratigraphy, however, have been used to argue for a much older Norian base in the Newark section (Muttoni et al. 2004), one that would be close to the base of the Lockatong Formation, with an estimated age of c. 228 Ma based on Newark cycle counting. Furthermore, in an abstract, Irmis & Mundil (2008) reported a 206Pb/238U age of 219.2 + 0.7 Ma for an Adamanian horizon of the Chinle Group in west –central New Mexico. On face value, the Chinle date and the interpretation of Newark magnetostratigraphy of Muttoni et al. (2004) indicate that the Adamanian is Norian. Nevertheless, the correlations Muttoni et al. (2004) propose between the Newark and the marine Late Triassic magnetostratigraphy from Pizzo Mondello are fraught with problems, mostly because the marine section contains far fewer magnetochrons than does the presumed age-equivalent interval of the Newark. Furthermore, the correlation has abandoned the only well-documented biostratigraphic datum in the Newark that allows a correlation to marine strata: the Carnian –Norian boundary at the approximate base of the Passaic Formation (see above). Thus, the proposed Pizzo Mondello-Newark magnetostratigraphic correlation lacks an independent biostratigraphic datum by which to correlate. Furthermore, the Pizza Mondelo marine section is thin (c. 430 m of limestone-dominated section represent late Carnian and much of Norian time) in comparison to the more than 4-km-thick Newark section. Therefore, it is not surprising that the Pizzo Mondello section yields a magnetostratigraphic record that does not directly correspond, in both reversal frequency and pattern, to the Newark section. I thus believe there is real reason to question the reliability of the magnetostratigraphic correlations advocated by Muttoni et al. (2004). I maintain a late Carnian (Tuvalian) age for the Adamanian, choosing biostratigraphic data over what I judge to be less reliable correlations based on magnetostratigraphy. As for the date reported in an abstract by Irmis & Mundil (2008), without supporting data its reliability cannot be fully evaluated.

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However, if it is a reliable age, it dates part of the late Carnian to c. 219 Ma, which means the base of the Norian would be younger than 219 Ma, in agreement with the 217 Ma age suggested by Kent & Olsen (1999), and the Norian is not as long as concluded by Muttoni et al. (2004).

Revueltian Two Italian records are critical to correlation of the Revueltian to part of the Norian: 1. Zorzino Limestone, Lombardian Alps, Italy: The Zorzino Limestone (Calcare de Zorzino) has been correlated to the mid-Norian (uppermost Alaunian) Himavatites columbianus ammonite zone (Jadoul et al. 1994; Roghi et al. 1995). Nonmarine tetrapods from this unit at the Ceˆne and Endenna quarries in Lombardy include the Revueltian index taxa Mystriosuchus, Aetosaurus and Eudimorphodon (Wild 1989; Renesto 2006). 2. Forni Dolomite, Veneto Prealps, Italy: the Forni Dolomite (Dolomia di Forni) in northeastern Italy is the same age as the Zorzino Limestone, midNorian (Roghi et al. 1995). Its nonmarine tetrapods include the Revuletian index taxon Eudimorphodon (Dalla Vecchia 1995). The Italian records thus provide direct evidence that at least part of the Revueltian ¼ middle Norian (Alaunian). I consider the Revueltian to correlate approximately with the early-middle Norian, which is consistent with the Italian data (Lucas 1997a). Palynostratigraphy, magnetostratigraphy and sequence stratigraphy suggest the characteristic Revueltian tetrapod assemblage in the Chinle Group of New Mexico, USA, is of Norian age (Lucas 1997a, 1998a). Based on stratigraphic position (Huber et al. 1993b; Lucas & Huber 2003), magnetostratigraphy (Witte et al. 1991; Kent et al. 1995; Muttoni et al. 2004), and palynomorphs (Cornet 1977), the Neshanician LVF in the Newark Supergroup of eastern North America is of early to middle Norian or just of middle Norian age. Stratigraphic position (Huber et al. 1993b; Lucas & Huber 1993), magnetostratigraphy (Witte et al. 1991; Kent et al. 1995; Muttoni et al. 2004), and palynomorphs (Cornet 1977; Fowell & Olsen 1993; Lucas & Tanner 2007b) indicate the Cliftonian LVF is of late Norian –Rhaetian age. Thus, a Norian correlation of the Revueltian is certain, with well supported correlation to the early and middle Norian.

Apachean Apachean time is post-Revueltian (c. mid-Norian) and pre-Jurassic. Magnetostratigraphy of the uppermost Chinle Group in the Four Corners and in eastern New Mexico (Reeve & Helsley 1972; Molina-Garza et al. 1996, 2003), correlated to the Newark Supergroup magnetostratigraphy (Kent et al. 1995;

Muttoni et al. 2004; Hounslow & Muttoni 2010, this volume), also suggests the Apachean is latest Triassic (‘Norian– Rhaetian’). Earlier arguments that the Apachean is equivalent to the Rhaetian (Hunt 1993; Lucas 1993b, 1998a) cannot be sustained in the light of new data. These arguments were largely based on a stage-of-evolution assessment of the Apachean phytosaur Redondasaurus. This phytosaur is more derived than the Knollenmergel (late Norian) phytosaurs of the German Keuper, so Redondasaurus was therefore assigned a Rhaetian age. However, the Norian aetosaur Aetosaurus occurs in Rock Point strata in Colorado (Small 1998) and Rock Point strata in New Mexico, and the Rock Point palynomorphs suggest a Norian age (Litwin et al. 1991). Clearly, the Apachean is younger than the Revueltian (early –middle Norian), so I regard it as late Norian to Rhaetian in age (Lucas et al. 2005, 2007e; Lucas & Tanner 2007a, b). The stratigraphically highest Apachean assemblage from the American Southwest is in the Dinosaur Canyon Member of the Moenave Formation and laterally equivalent Wingate Sandstone (Lucas et al. 2005, 2006; Lucas & Tanner 2007a, b). There are several compelling reasons to assign a Late Triassic age to this assemblage: (1) the Apachean phytosaur Redondasaurus is present, and no phytosaur is known from Jurassic strata; (2) the footprint ichnogenus Brachychirotherium is present and not known anywhere from Jurassic strata; (3) the lower Dinosaur Canyon Member is laterally equivalent to strata of well established Late Triassic age (upper Rock Point Formation of the Chinle Group); (4) the Wingate Formation basal contact is gradational with underlying Upper Triassic strata of the Rock Point Formation; and (5) magnetostratigraphy of the Dinosaur Canyon interval is reasonably correlated to the magnetostratigraphy of uppermost Triassic strata of the Newark Supergroup in eastern North America (MolinaGarza et al. 2003). Although it is possible to assign the Dinosaur Canyon assemblage to the Late Triassic, its precise correlation to the marine timescale is uncertain. Probably it equates to part or all of Rhaetian time, simply because the Dinosaur Canyon interval is the youngest Triassic interval on the Colorado Plateau and is conformably overlain by strata that apparently correlate to the earliest part of the Early Jurassic (Hettangian) (Lucas & Tanner 2007a, b; Kozur & Weems 2010). This supports a correlation of the Apachean with the late Norian–Rhaetian.

Triassic Footprint Biostratigraphy In this volume, Klein & Lucas (2010) present a Triassic footprint biostratigraphy and biochronology that build on, revize and synthesize previous

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efforts, including those of Haubold (1969, 1971, 1984, 1986), Demathieu & Haubold (1972, 1974), Olsen (1980, 1983), Lockley & Hunt (1995), Hunt & Lucas (2007a, b), Lucas (2003, 2007a) and Klein & Haubold (2007). Triassic tetrapod footprints have a Pangaea-wide distribution; they are known from North and South America, Greenland, Europe, North Africa, China, Australia, Antarctica and South Africa. They often occur in nonmarine Triassic strata that lack well-preserved body fossils, so their biostratigraphic utility has been of some interest. In Triassic strata, several characteristic footprint assemblages and ichnotaxa have restricted stratigraphic ranges and thus represent distinct time intervals. Key Triassic footprint ichnotaxa are archosaur tracks: Rotodactylus, the chirotherian ichnotaxa Protochirotherium, Synaptichnium, Isochirotherium, Chirotherium, Brachychirotherium and grallatorids (theropod dinosaur tracks). Nevertheless, non-archosaur footprints are common, especially the ichnotaxa Rhynchosauroides, Procolophonichnium, Capitosauroides and several dicynodontrelated or mammal-like forms that dominate some footprint assemblages. From the temporal distribution pattern Klein & Lucas (2010) identified five distinct tetrapodfootprint-based biochrons: (1) dicynodont tracks (Lootsbergian); (2) Protochirotherium (Synaptichnium): also includes Rhynchosauroides and Procolophonichnium (Nonesian); (3) Chirotherium barthii, also includes C. sickleri, Isochirotherium, Synaptichnium, Rotodactylus, Rhynchosauroides, Procolophonichnium, dicynodont tracks and Capitosauroides (Nonesian –Perovkan); (4) Atreipus –Grallator (‘Coelurosaurichnus’), which also includes Synaptichnium, Isochirotherium, Sphingopus, Parachirotherium, Rhynchosauroides, Procolophonichnium (Perovkan–Berdyankian); and (5) Brachychirotherium, which also includes Atreipus –Grallator, Grallator, Eubrontes, Apatopus, Rhynchosauroides, dicynodont tracks (Otischalkian–Apachean). Tetrapod footprints are thus useful for Triassic biostratigraphy and biochronology, but, compared to the tetrapod body fossil record with eight biochrons, the five footprint-based biochrons provide less temporal resolution. Nevertheless, in nonmarine Triassic strata where body fossils are rare, footprints can be useful for biostratigraphy and biochronology.

Conclusion The global Triassic timescale based on tetrapod evolution developed in the 1990s has been criticized because of: (1) perceived problems with the alpha taxonomy of some of its index fossils; (2) possible temporal overlap of the Nonesian and Perovkan

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LVFs; (3) changes and additions to the stratigraphic ranges of some index taxa; and (4) perceived problems of correlation to the SGCS. Taxonomic disagreements lie at the heart of many arguments over biostratigraphy, but I believe that the extensive taxonomies developed for many of the Triassic index taxa, especially temnospondyls, phytosaurs, aetosaurs, dicynodonts and cynodonts, provide a sound basis for their use in biostratigraphy. Shifting opinions about taxonomy of these tetrapods will remain, and that will always affect correlations based on tetrapod fossils. Lucas et al. (2007e) resolved the problems of potential overlap or gaps around the Nonesian –Perovkan boundary by redefining the beginning of the Perovkan to obviate such problems. Stratigraphic range extensions and changes are the regular outgrowth of collecting and careful biostratigraphic study in the field. They always force adjustments to any biochronological scheme rooted in sound biostratigraphy. Problems with correlation of the Triassic LVFs to the SGCS persist largely because in much of the nonmarine Triassic section few reliable data are available for correlation to the marine timescale. Clearly, we need a nonmarine Triassic tetrapod biochronology with which to correctly sequence the history of tetrapod evolution on land. Advances in the scheme proposed in the 1990s have come from new fossil discoveries, more detailed biostratigraphy and additional alpha taxonomic studies based on sound evolutionary taxonomic principles. As the work reviewed here demonstrates, the global Triassic timescale based on tetrapod biochronology remains a robust tool for both global and regional age-assignment and correlation. The synthesis reported here is based on work funded by the National Geographic Society, DAAD and the Janet Stearns Memorial Trust. I am grateful for the collaboration of A. Heckert, P. Huber, A. Hunt, J. Spielmann and L. Tanner on many problems of Triassic tetrapod taxonomy and biostratigraphy. I thank A. Heckert, H. Kozur, J. Ogg, S. Renesto, J. Spielmann and R. Weems for their careful and helpful reviews of the manuscript.

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W ESTPHAL , F. 1970. Phytosaurier-Hautplatten aus der Trias von Madagaskar-ein Beitrag zur Gondwana– Pala¨ogeographie. Neues Jahrbuch fu¨r Geologie und Pala¨ontologie Monatshefte, 1970, 632– 638. W HITESIDE , D. I. & M ARSHALL , J. E. A. 2008. The age, fauna and palaeoenvironment of the Late Triassic fissure deposits of Tytherington, South Gloucestershire, UK. Geological Magazine, 145, 105– 147. W ILD , R. 1978. Die Saurier von Kupferzell-Bauersbach. Wu¨rttembergisch Franken Jahrbuch, 62, 1– 16. W ILD , R. 1980. The fossil deposits of Kupferzell, southwest Germany. Mesozoic Vertebrate Life, 1, 15–18. W ILD , R. 1989. Aetosaurus (Reptilia: Thecodontia) from the Upper Triassic (Norian) of Cene near Bergamo, Italy, with a revision of the genus. Rivista del Museo Civico di Scienze Naturali, 14, 1 –24. W IMAN , C. 1910. Ein paar Labyrinthodonten Reste aus der Trias Spitzbergens. Bulletin of the Geological Institute of the University of Uppsala, 9, 34– 40. ¨ ber die Stegocephalen saus der Trias W IMAN , C. 1915. U Spitzbergens. Bulletin of the Geological Institute of the University of Uppsala, 13, 1 –34. W IMAN , C. 1916. Neue Stegocephalenfunde aus Posidonmyaschiefer Spitzbergens. Bulletin of the Geological Institute of the University of Uppsala, 13, 209– 222. W ITTE , W. K., K ENT , D. V. & O LSEN , P. E. 1991. Magnetostratigraphy and paleomagnetic poles from Late Triassic–earliest Jurassic strata of the Newark basin. Geological Society of America Bulletin, 103, 1648–1662. W ITZMANN , F., S CHOCH , R. R. & M AISCH , M. W. 2008. A relict basal tetrapod from Germany: first evidence of a Triassic chronosuchian outside Russia. Naturwissenschaften, 95, 67–72. W OODY , D. T. 2003. Revised Geological Assessment of the Sonsela Member, Chinle Formation, Petrified Forest National Park, Arizona. MSc thesis, Northern Arizona University, Flagstaff. W OODY , D. T. 2006. Revised stratigraphy of the lower Chinle Formation (Upper Triassic) of Petrified Forest

National Park, Arizona. Museum of Northern Arizona Bulletin, 62, 17–45. W OODY , D. T. & P ARKER , W. G. 2004. Evidence for a transitional fauna within the Sonsela Member of the Chinle Formation, Petrified Forest National Park, Arizona. Journal of Vertebrate Paleontology, 24(3), 132A. Y ADAGIRI , P. & R AO , B. R. J. 1987. Contribution to the stratigraphy and vertebrate fauna of the Lower Jurassic Kota Formation, Pranhita–Godavari Valley, India. The Palaeobotanist, 36, 230–244. Y ATES , A. M. 2003. A definite prosauropod dinosaur from the Lower Eliot Formation (Norian: Upper Triassic) of South Africa. Palaeontologia Africana, 39, 63–68. Y ATES , A. M. & K ITCHING , J. W. 2003. The earliest known sauropod dinosaur and the first steps toward sauropod locomotion. Proceedings of the Royal Society of London, B270, 1753– 1758. Y IN , H., Z HANG , K., T ONG , J., Y ANG , Z. & W U , S. 2001. The global stratotype section and point (GSSP) of the Permian–Triassic boundary. Episodes, 24, 102–114. Y IN , H., T ONG , J. & Z HANG , K. 2005. A review on the global stratotype section and point of the Permian– Triassic boundary. Acta Geologica Sinica, 79, 715–728. Y IN , H., F ENG , Q., L AI , X. & T ONG , J. 2007. The PermoTriassic crisis is prolonged, and the PTB mass extinction is multiphase. New Mexico Museum of Natural History and Science Bulletin, 41, 414. Z AVATTIERI , A. M. & A RCUCCI , A. N. 2007. Edad y posicio´n estratigrafica de los tetra´podos del Cerro Bayo de Potrerillos (Tria´sico), Mendoza, Argentina. Ameghiniana, 44, 133 –142. Z EIGLER , K. E., H ECKERT , A. B. & L UCAS , S. G. (eds) 2003. Paleontology and geology of the Upper Triassic (Revueltian) Snyder quarry, New Mexico. New Mexico Museum of Natural History and Science Bulletin, 24, 1 –132.

Index Note: Page numbers in italic denote figures. Page numbers in bold denote tables. Acadiestheriella 390, 398 Acadiestheriella cameroni 320 –321, 351, 398, 399 Acallosaurus 465 Acallosuchus 464 Acanthotetrapaurinella 176 Acanthotriletes varius 293, 303 Accinctisporites 300 Acrochordiceras carolinae 241 Adamanasuchus 464 Adamanian LVF 383, 448, 449, 462, 463– 469 correlation with Carnian Stage 475, 479 –480 Adamanian –Revueltian boundary 355– 356 Adelobasileus 465 Aegean substage 30 Aegeiceras 242 Aegeiceras ugra 143, 146 Aetosaurus 338, 462, 469, 470, 471, 472, 473, 475, 480 Aetosaurus arcuatus 386, 469 Aghdarbandites ismidicus 240 Alaunian substage 30, 32 Albaillella 170 Alberti, Friedrich August von (1795–1878), Trias 18, 19–20 Alisporites 289, 300, 302 Alisporites warepanus 300 Almasaurus 462 Alpine marine Triassic cyclostratigraphy 124– 128 early work 20 Alwalkeria 463 ammonoids 6, 10–12, 221– 256 Alpine regions 230–231 Anisian 234, 238, 239, 240, 242 Anisian–Ladinian boundary 148, 149 Bagolino section 228 biochronology 254 biostratigraphy 33–34, 254 Canada 229– 230, 235, 236, 245, 247–249, 251 Carnian 239, 246, 247–250 Carnian–Norian boundary 153, 154 Chiangsingian– Induan boundary 141 Dienerian 232–234, 236 early work 20, 21, 22–24, 221–229 Germanic Basin 230 Griesbachian 232– 234, 235 Himalayas 231 Induan 232, 233

Induan– Olenekian boundary 66–68 Ladinian 234, 242, 244, 245, 254 Ladinian–Carnian boundary 148, 151, 254 Lower Triassic 232– 234, 235, 236, 237 zones 206, 234 Middle Triassic 234–239, 240, 241, 242, 243, 244–245 zones 242, 246 Norian 239, 246, 248, 250, 251, 252 Norian– Rhaetian boundary 156 North American 23– 24 Olenekian 232, 233, 234 Olenekian– Anisian boundary 143– 144, 144, 146 palaeoenvironment 255 Permian– Triassic boundary 64 Rhaetian 92, 239, 246, 251, 252 Siberia 231– 232 Smithian 232, 233, 234 South China 231 Spathian 232, 233, 234 taxonomic problems 252 –254 Tethys regions 230– 231, 238, 242, 240, 244, 249, 250, 252 Trans-Caucasia 231 Upper Triassic 239, 243, 246, 247– 252 zones 246, 252 USA 230, 237, 241, 243 zones 10– 12, 209, 234, 242, 246, 251, 255 Amuria 176, 185, 192 Anagymnotoceras 242 Anapiculatisporites 267, 274 Anasibirites 68, 234 Anasibirites kingianus 70, 72, 237 Anaticapitula 187 Anatropites spinosus 250 Angistorhinus 461, 462, 463, 464, 465 Angonisaurus 455, 459 Angulocircus 179 Angulopaurinella 175 Angustisulcites 266, 274 Angustisulcites grandis 273 Angustisulcites klausii 273 Anisian Stage 2, 11, 25, 29–30 ammonoids 234, 238, 239, 240, 242 correlation with Perovkan LVF 475, 478 magnetostratigraphy 68, 69, 71, 78

radio-isotope dating 46, 47, 48, 49, 52 radiolarians 170– 174, 183, 187, 189–190 stable-isotope record 107, 108, 109–110 see also Olenekian –Anisian boundary Anisian–Ladinian boundary ammonoids 148, 149 conodonts 148 magnetostratigraphy 73–76, 78 radio-isotope dating 42, 47, 47, 52 radiolarians 190 Anisicyrtis 169, 171 Annulispora 268 Annulispora folliculosa 300 Annulispora microannulata 300 Annulobulbocyrtium 175, 187 Annulohaeckelella 176 Annulopoulpus 180 Annulosaturnalis 178, 188 Annulotriassocampe 174, 183 anoxia, ocean 107, 108, 109, 110 Ansubuga 174 Antecosuchus 458 Antetonitrus 473 Anyuanestheria 333, 382, 386 Anyuanestheria wingatella 322, 352, 359 zone 381, 384 Apachean LVF 360, 448, 449, 462, 472–474 correlation with Rhaetian Stage 475, 480 Apachesaurus 461, 462, 463, 465, 469, 472 Aparimella 201, 202, 214 Apatopus 434–435, 437, 439, 440, 481 Aphaneramma 456, 475, 477 Apiculatisporis lentus 300 Apiculatisporites parvispinosus 278 Apiculatisporites plicatus 266, 273 40 Ar/39Ar isotope dating 3, 42, 43–44, 54 Late Triassic 49, 51 Middle Triassic 46–49, 47, 50 Permian–Triassic boundary 45 Triassic– Jurassic boundary 51 Aratrisporites 264, 266, 267, 273, 274, 278, 300 Aratrisporites astigmosus 267 Aratrisporites compositus 300 Aratrisporites flexibilis 300 Aratrisporites tenuispinosus 301 Araucariacites 299

502 Arcestes 111, 246 Archaeoacanthocircus 181, 188 Archaeocenosphaera 174, 192, 193 Archaeohagiastrum 187 Archaeosemantis 169, 171, 190 Archaeospongoprunum 170, 185 Arcicubulus 171 Arctoceras tuberculatum 237 Arctotraversodon 461 Arganasaurus 463 Argentina, Late Triassic palynology 300, 302 Arizonasaurus 453, 458 Asseretosporites 267 Asseretosporites gyrate 267 Astrocentrus 173 Atalantria 186 Atreipus-Grallator 427, 430, 431, 434, 435, 436, 438, 439, 440, 481 Aulisporites 286, 292 Aulisporites astigmosus 267, 274, 275, 278, 292 Auritulinasporites scanicus 300 Australia, Late Triassic palynology 298– 299 Austriadactylus 470 Austrisaturnalis 178, 188 Austrotrachyceras 246 Austrotrachyceras austriacum 250 Austrotrachyceras obesum 247 Ayrtonius 182, 184, 188 Azendohsaurus 462, 468 Baculatisporites comaumensis 302 Bagolino GSSP 47, 148, 227, 228, 277 Balatonites 239, 242, 253 Balatonites balatonicus 29, 240 Balatonites ottonis 73 Balatonites shoshonensis 46, 241 Baloghisphaera 175 Baratuna 172, 183 Basutodon 473 Batrachopus 435, 440 Batrachotomus 460 Baumgartneria 175 Bauria 457 Beatricea 187 Beijianglimnadia dalongkouensis 362 Beijianglimnadia qitaensis 362 Beijianglimnadia rotunda 363 Belesodon 460 Bellerophon Formation, PT boundary 64, 65 Beneckeia buchi 70, 73 Beneckeia tenuis 70, 73 bennettitaleans 9 Benthosuchus 70, 454, 477 Berdyankian LVF 448, 449, 453, 459– 461 correlation with Ladinian Stage 475, 478

INDEX Berlahmium 179 Betraccium 181, 184, 188 Beturiella 173 Bipemphigus 324, 362 Bipemphigus gennisi 362 Bipemphigus liaoningensis 362 biochronology ammonoids 254 tetrapod footprints 438– 440 tetrapods 447– 481 bioproductivity crises 107, 109, 110, 112, 113 biostratigraphy 5–13 ammonoids 221–256 ammonoids v. conodonts 33–34 conchostracans 316–404 importance of 319– 321 early work 20, 21, 22 radiolarians 163–194 tetrapods 447– 481 biotic events 34 stable isotope record 103, 104, 107– 108, 109, 110 Bipedis 182, 186, 188, 193 Bithynian substage 30 Bittner, Alexander (1850– 1902), Triassic timescale 30, 31 Bitubopyle 178, 184 bivalves 201–215 Lower Triassic 206, 207 –208, 208 Middle Triassic 208, 209, 210 non-marine 12– 13 palaeobiogeography 204– 206 palaeoecology 203– 204 Upper Triassic 209, 210– 211, 212, 213 Blechschmidtia 181 Blomidon Formation conchostracans 341, 351 cyclostratigraphy 128, 129–130 Blue Mesa Member 355, 464, 465 Bluewater Creek Formation conchostracans 352–353, 355 tetrapods 464, 465 Bogdanella 177 bolide impact, carbon isotope excursion 107, 108 Boraeosaurus 456, 477 Boreal marine fossils 34–35 Borinella nepalensis 143, 145 Bosniacyrtis 176 Brachychirotherium 422, 423, 425, 427, 428, 429, 430, 432, 433, 434, 438, 439, 440, 473 Brachychirotherium hassfurtense 430, 433, 436 Brachychirotherium parvum 433 Brachychirotherium thuringiacum 430, 433, 436 Brachyrhinodon 467 Brachyzostrodon 470

Braginella 182, 188 Brasilichnium 438 Briery Creek Basin, conchostracans 339, 341, 346 British Columbia, ammonoids 230 Brodispora striata 289, 305 Buch, Leopold von (1774–1853) Alpine marine Triassic 20 Keuper 18–19 Buchenstein Beds ammonoids 254 magnetostratigraphy 77–78 U– Pb dating 125, 126 Buckman, Sydney S. (1860– 1929) 22– 23 Budurovignathus diebeli 153 Budurovignathus gabriellae 148 Budurovignathus hungaricus 150 Budurovignathus mungoensis 152, 153 Budurovignathus praehungaricus 148, 150 Buettneria 461, 462, 463, 465, 468 Buettneria perfecta 344 Bukkenites nanus 235 Bukobaja 458 Bulbilimnadia 335, 404 Bulbilimnadia froelichi 343, 345 zone 381, 389 Bulbilimnadia killianorum 335, 343, 356–357, 400, 402, 404 zone 381, 389 Bulbilimnadia sheni 343, 389 zone 381, 389 Bulbocyrtium 179 bundling patterns 120, 121, 132 Bunter see Buntsandstein Buntsandstein 327, 328, 329 astronomical timescale 94 cyclostratigraphy 122, 123– 124 early work 18, 19–20 nonmarine magnetostratigraphy 70, 72 Cacheutasporites wielandii 300 Cadargasporites cuyanensis 300 Calamospora 266, 278, 300 Calamospora tener 276 Callialasporites 299 Callialasporites dampieri 288 Callistomordax 460 Calvo¨rde Formation 332, 367– 368 Camarazonosporites 292, 293 Camarazonosporites laevigatus 303 Camarazonosporites rudis 267, 274, 289, 300 Camerosporites 286, 293 Camerosporites pseudoverrucatus 278, 292 Camerosporites secatus 267, 274, 278, 288, 289, 292, 302, 305 phase 289, 291 zone 270, 274 –275

INDEX Camerosporites verrucosus 289, 300 Canada, ammonoids 229– 230, 238, 235, 236, 245, 247–249, 251 Canesium 179 Cannanoropollis 278 Canoptum 178, 185, 187, 188, 192, 193 Cantalum 182, 184 Caphtorocyrtium 181, 188 Capitosauroides 422, 425, 427, 481 Capnodoce 181, 188 Capnuchosphaera 179, 187 carbon isotopes 103– 104, 105 Lower Triassic 106–109 Middle Triassic 109 –110 Permian–Triassic boundary 65, 71, 106– 108 Triassic –Jurassic boundary 90, 92 Upper Triassic 110–113, 286 carbonate platforms, cyclicity 131– 132 Carinacyclia 179 Carinaheliosoma 179 Carnian Stage 2, 12, 25, 30, 31 ammonoids 239, 246, 247–250 conchostracans 355, 359 correlation with Adamanian LVF 475, 479 –480 correlation with Otischalkian LVF 475, 478–479 GPTS magnetostratigraphy 81–82 palynomorphs, northern hemisphere 288–292, 301, 302 radio-isotope dating 49, 51, 52 radiolarians 171, 173– 181, 187–188, 189– 190 stable-isotope record 107, 110–111 see also Ladinian–Carnian boundary Carnian-Norian boundary ammonoids 153, 154 carbon isotope record 110 conchostracans 355, 359 conodonts 153–154, 155, 158 magnetostratigraphy 79, 80– 82 Newark Supergroup 86–87 radiolarians 191 Carniepigondolella samueli 154 Carnisporites 276 Catenohalorites 246 Catoma 181 Cauletella 170 Celluronta 171 Cenosphaerocapsula 178 Central Atlantic Magmatic Province (CAMP) 51, 285, 286, 390 Rhaetian carbon isotope excursion 112– 113 Ceratites 239 Ceratites binodosus 29, 30, 225, 253 Ceratites reitzi 253–254

Ceratites trinodosus 29, 30, 225, 253 Cerebellocapsula 177 Cerebropollenites macroverrucosus 288, 292, 297, 303 Cerebropollenites thiergartii 89, 92, 268, 288, 297 zone 270, 276 –277 Chaliminia 473 Chanaresuchus 460 Charlottea 186 charophytes 9 Chasmatosporites 268, 292, 293, 294 Chasmatosuchus 454 Chatterjeea 465, 469 Changhsingian Stage carbon isotope excursion 107 conchostracans 325, 361– 367 magnetostratigraphy 64, 65, 66 radio-isotope dating 44–45, 53 sulphur isotopes 108 Changhsingian –Induan boundary ammonoids 141 conodonts 139– 140, 141, 142 China, northwest conchostracans 324–327 tetrapods 454 China, south ammonoids 231 cyclostratigraphy 122 –123 Chiniquodon 467 Chinle Group 352, 355 conchostracans 356 cyclostratigraphy 130 –131 tetrapods 461, 464, 465, 468, 469, 472 Chinlechelys 469 Chiosella 146, 147 Chiosella gondolelloides 144, 146, 147, 329 Chiosella timorensis 68, 69, 71, 144, 146, 147, 315, 329 Chirotherium 422 Chirotherium barthii 422, 425, 426, 427, 438, 439, 481 Chirotherium lulli 434, 437 Chirotherium moquiense 422 Chirotherium rex 422 Chirotherium sickleri 422, 425, 426, 438, 439, 481 Choristoceras 51 Choristoceras crikmayi 111, 251 Choristoceras marshi 252 chronostratigraphy 1– 3, 2, 17– 35 American 23–25, 35 biotic events 34 early work 17–32 Germanic Basin 8 nomenclature stability 32– 33 Old World v. New World 35 provinciality 34– 35 subdivision 25– 32 taxonomy 33 Cingulizonates rhaeticus 276, 292, 297

503 Circulina 267, 287, 306 Citriduma 182, 186, 188, 193 clam shrimp see conchostracans Claraia 201, 202, 203, 206, 207, 213 Claraia aurita 206, 207, 324 Claraia clarae 207, 208 Claraia concentrica 207 Claraia dieneri 206, 207 Claraia griesbachi 207 Claraia stachei 206, 207 Claraia wangi 206, 207 Clarkina praetaylorae 140 Classopollis 89, 92, 267, 274, 276, 287, 293, 298, 302, 306 Classopollis chateaunovi 300 Classopollis meyeriana 267, 268, 275, 276, 288, 294, 298, 304 Classopollis murphyae 288, 298, 304 Classopollis torosus 267, 268, 288, 294, 298, 299, 304 Classopollis zwolinskae 267, 275, 276 clathrates see methane hydrate dissociation ‘Clepsysaurus pennsylvanicus’ 470 Clevosaurus 467, 473 Cliftonian LVF 470 climate, as control on sedimentation 131 Coahomasuchus 461 coal gap 109 Cochloceras 246 Coelophysis bauri 473 ‘Coelurosaurichnus’ see Atreipus-Grallator Colognathus 465 Colognathus obscurus 469 Coloradisaurus 473 Colorado Plateau, cyclostratigraphy 130–131 Columbites parisianus 237 Compsocerops 467 Conbaculatisporites 292 Concavisporites 92, 268, 276 Concentricisporites 266, 273, 278 Concentricisporites bianulatus 278, 292 Conchostraca 317 conchostracans 7, 10– 12, 315–404 biostratigraphic importance 319–321 characteristics 316– 319 Dalongkou 324– 327 distribution 319– 321, 357– 361 early work 321– 324 Asia 322– 323 Gondwana 323 North America and Europe 321– 322 Tethys 323– 324 eggs 317, 320 Germanic Basin 321–322, 325, 327–335, 357, 358– 359, 360, 368, 369

504 conchostracans (Continued) Hettangian, zonation 380 –389, 381, 389– 390 Lower Triassic 357–358 zonation 367–376 Middle Triassic 358 zonation 376–380 Newark Supergroup 321–322, 338–351 palaeoecology 317 Permian–Triassic, zonation 325, 361–367 Southwestern USA 351–357, 354 taxonomy 317–319 Upper Triassic 358–361 zonation 380–389 zonation 10–12, 325, 361– 390 Congestheriella elliptoidea 384, 385 conifers 9, 266, 267, 297, 298 conodonts 5, 94 Anisian–Ladinian boundary 73, 148 biostratigraphy 33–34 Carnian– Norian boundary 79, 80, 86– 87, 153– 154, 155, 158 Changhsingian –Induan boundary 139–140, 141, 142 Induan–Olenekian boundary 66, 140, 143, 144, 145 Ladinian–Carnian boundary 148–149, 151, 152, 153 Norian Stage 82–84 Norian–Rhaetian boundary 156, 157, 158 Olenekian–Anisian boundary 68, 71, 143–148 Permian–Triassic boundary 64, 139–140, 141, 142 Rhaetian Stage 84–86, 87, 89, 92 stage boundary definition 25, 139–158, 140 zones 10– 12 Conospongocyrtis 174 Continental Flood Basalt (CFB) events 45, 51 Convolutispora microrugulata 297, 300 Cordaitina minor 274 Cordevolic substage 31 Cornia 319, 322 Cornia beijiangensis 362 Cornia germari 320, 321, 322, 324, 367, 368, 369, 371 –372 Cornia germari-M. subcircularis zone 370, 372 Cornutella 176, 185 Cornutisporites 267 Cornutisporites rugulatus 276 Cornutisporites seebergensis 276 Corollina 267, 287, 306 Coronatubopyle 177 Corum 177, 187 Costatoria costata 358, 378

INDEX Courtilloticeras stevensi 143 Craterisporites rotundus 299, 300 Cricodon 459 Crosbysaurus 465, 469 Crucella 180, 186, 192, 193 Cryptostephanidium 171, 183 Culpeper Basin 339, 340, 342–343 conchostracans 347–348 cycadeoids 9 Cycadopites 269, 289, 300 cycads 9 Cyclestheriidae 317 cyclostratigraphy 4, 119– 133 Alpine marine Triassic 124– 128 bundling patterns 120, 121, 132 Colorado Plateau 130– 131 Germanic Basin 122, 123 –124 Latemar Massif 54, 77– 78, 121, 125– 126, 132 Lockatong Formation 128–130 Newark Supergroup 128– 130, 132 orbital forcing 119–120 South China 122–123 Cyclotosaurus 383, 467, 470, 475, 478 Cycloverrutriletes presselensis 265, 277 Cynognathus 448, 450, 452, 453, 455– 457 Cyrtopleurites 246 Dachstein Limestone, cyclostratigraphy 126–128 Dagys, A.S. (1933– 2000), ammonoids 227 Dalongkou, conchostracans 324– 327 Dan River/Danville Basin, conchostracans 338– 341, 346 Danubea 186 Daonella 201, 202, 203, 210, 214 Daonella americana 209, 210 Daonella dubia 209, 210 Daonella elegans 208, 209, 210 Daonella elongata 209, 210 Daonella lommeli 209, 210 Daonella moussoni 208, 209, 210 Daonella nitinae 209 Daonella picheri 209, 210 Daonella sturi 208, 209, 210 Daonella subarctica 208, 209 Daxatina 153 Daxatina canadensis 75, 76, 148, 229, 246, 249, 254 Daxatina desatoyense 243 Decussatisporites 286 Deep River Basin, conchostracans 338 Deep Run Basin 339, 342 conchostracans 347 Deflandrecyrtium 176, 184, 187, 193 Deltoidospora 92, 276 demagnetization 62

Densoisporites 70, 71, 264, 276, 278, 293 Densoisporites nejburgii 65, 264, 265 zone 269, 270, 273, 277 Densoisporites playfordii 65, 264, 269, 273 Densoisporites psilatus 300 Desmatosuchus 461–462, 462, 463, 465, 468 Desmatosuchus haplocerus 464 Diademodon 453, 456–457, 459 diagenesis, stable isotopes 105 Diaplexa tigjanensis, zone 377, 378 Diaplexa xuanhanensis 378 Dicapnuchosphaera 181, 188 Dicroidium 9 Dictyonatella 318 Dictyonatella dictyonata 369 Dictyophyllidites 298, 299 Dictyophyllidites mortonii 300, 302 Dicyclosporis 268 Dicynodon 70, 71, 325– 326, 364, 365, 366– 367, 455, 476 Dicynodontipus 422, 424, 427 Diener, Carl (1862–1928), Triassic timescale 20, 22, 29, 224 Dienerian substage 25, 29 ammonoids 232– 234, 236 stable-isotope record 107, 108 Difalsisca 367 Dimorphites 246 Dinodontosaurus 449, 460, 475, 478 Diphydontosaurus 470 Discofulmen 176 Discretella 143 Disphaerocapsula 177 Distcyclosporis 268 Divatella 178 Dolomites Anisian– Ladinian boundary, magnetostratigraphy 73, 74 Latemar Massif, cyclostratigraphy 54, 77– 78, 121, 125– 126, 132 Doswellia 461– 462, 462, 463 Doubingerispora filamentosa 274 Draculacampe 176 Drepanites 246 Drepanites ruthefordi 248 Drepanosaurus 470, 475 Dreyericyrtium 175, 187 Droltus 186, 193 Dromicosuchus 462 Dumitricasphaera 177 Duplexisporites 286 Duplexisporites problematicus 299, 300 Duplicisporites continuus 289 Duplicisporites granulatus 278, 288, 289, 292, 293, 301 Duplicisporites verrucosus 278, 292 Durham Sub-basin, conchostracans 338, 339, 346

INDEX Dutuitosaurus 462 Dyupetalum 266 Dyupetalum vicentinense 266 Ebrachosaurus singularis 468 eccentricity cycles 120, 125, 133 Echinitosporites iliacoides 267, 274, 291 Echinolimnadia 362 Echinolimnadia mattoxi 362– 363 Elbistanium 177 Ellesmerella 202, 214 Ellesmerella aranea 206, 208, 208 Ellipsovelatisporites plicatus 289 Ellipsovelatisporites rugosus 293, 294 Ellisus 181 Empirea 180, 186 Endennasaurus 470 Endosporites 264 Endosporites pallidus 264 Endosporites papillatus 269 Entactinia 170, 190 Enteropleura 201, 202, 203, 209, 210, 214 Enteropleura jenksi 208, 209 Enzonalasporites 276, 286, 293 Enzonalasporites vigens 82, 267, 274, 278, 288, 289, 292, 293, 294, 299, 301, 305 Eocursor 473 Eocyclotosaurus 77, 80, 475, 478 biochronology 448, 453, 455, 457–459 Eogymnotoceras thompsoni 241 Eomonotis 202, 203, 209, 211, 214 Eomonotis pinensis 209, 211, 212 Eomonotis scutiformis 209, 211 Eonapora 173, 183 Eoprotrachyceras curionii 76, 148, 190, 229, 242, 244 Eoprotrachyceras lahontanum 148 Eoprotrachyceras matutinum 245 Eoprotrachyceras subasperum 148, 149, 243 Eoraptor 467 Eosauropus 431, 434, 438, 439 Eosolimnadiopsis gallegoi 318, 381, 382 Ephedripites 269 Ephedripites primus 305 Epigondolella abneptis 79, 80, 356 Epigondolella bidentata 85–86, 157 Epigondolella englandi 157 Epigondolella mosheri 157, 158 Epigondolella nodosa 79, 80 Epigondolella quadrata 154, 155, 191 Eptingium 171, 183, 184, 192 Equisetosporites chinleanus 289 Erpetosuchus 467 Eryosuchus 453, 458, 459 Erythrosuchus 453, 456, 457, 458, 459

‘Estheria’ Beds 333, 334 ‘Estheria’ pennsylvanica 344 Estheriella 319, 332, 370 Estheriella bachmanni 367, 368, 369–370, 371 zone 369–370, 370 Estheriella costata 320, 368, 371 Estheriella marginostriata 368, 369, 370–371 zone 370–371, 370 Estheriella nodosocostata 320, 368, 371 zone 370, 371 –372 ‘Estherites’ atsuensis 359 ‘Estherites’ nakazawai 359 Etalian–Kaihikuan stage boundary, 40 Ar/39Ar dating 42 Eubrontes 430, 432–433, 435, 438, 440, 481 Eudimorphodon 462, 469, 470, 475, 480 Euestheria albertii albertii 318, 352, 369 zone 377–378 Euestheria albertii mahlerselli 317–318, 369, 391– 392 zone 375 Euestheria albertii mahlerselli-P. alsatica alsatica zone 376–377 Euestheria brodieana 332, 335, 343, 351, 356–357, 360– 361, 388, 389, 400 zone 381, 388 –389 Euestheria bunopasi 386 Euestheria buravasi 338 zone 381, 386 Euestheria dactylis 379 Euestheria exsecta 352, 369, 375, 376 zone 375 Euestheria franconica 333, 379 zone 377, 379 Euestheria gutta 364, 365, 366, 367, 369, 371, 372 Euestheria gutta gutta 373 Euestheria hausmanni 383, 386 Euestheria hubeiensis 379 Euestheria jakutica 365 Euestheria lepida 379 Euestheria minuta 318, 320, 333, 345, 358, 380 zone 377, 379 –380 Euestheria oertlii 365, 366, 367, 369, 370 Euestheria shizibaoensis 379 Euestheria winterpockensis 318, 342, 380 Euflemingites 29, 66, 67, 68, 140 Eumorphotis 207 Euparkeria 456, 457 Eurycervix posthumus 478 Eurygnathodus costatus 143, 145 Eurygnathodus hamadai 143, 145

505 eustasy, Triassic cyclicity 121– 122, 130, 131 Evazoum 431, 434, 438, 439 Exeraetodon 460, 463, 467 Exter Formation 331, 334 extinction, mass PT boundary 44–45, 65, 66, 71, 474–476 carbon isotope excursion 65, 71, 107 –108, 109, 474– 476 radiolarians 189 Triassic– Jurassic boundary 51, 286 radiolarians 189, 192 Falcispongus 175 Falcisporites australis 299, 300 Falcisporites nuthallensis 300 Falsisca 318, 324, 326–327, 362–365 Falsisca beijiangensis 326, 362 Falsisca bolodekitensis 364 Falsisca eotriassica 326, 363–364, 366, 368 zone 325, 361–362, 363–364, 370 Falsisca jeskinoica 363 Falsisca kanandaensis 326, 362 Falsisca novozilov 366 Falsisca podrabineki 364, 365 Falsisca postera 364–365, 366, 368 zone 325, 361–362, 364–365, 370 Falsisca secreta 363 Falsisca turaica 326, 362, 363 Falsisca turaica-F. zavjalovi zone 363 Falsisca verchojanica 325, 364–367, 368 zone 365– 367, 370 Falsisca zavjalovi 326, 362, 363, 364 Falsisca zavjalovi-T. minutus zone 325, 362 Farmville Basin, conchostracans 339, 341–342, 346 Fasolasuchus 473 Fassanian substage 30– 31 Feixanguan Formation 64, 65, 66, 67 fern spike 298 Ferresium 182, 184, 188, 189 Fischer plots 120, 121 ‘flat clams’ see bivalves Flemingites 140, 234 Flemingites flemingianus 72 Flemingites verchojanica 72 Flexispongus 174 Follicucullus 170 Fontinella 182, 185, 193 footprints see tetrapods, footprints Foremanellina 175, 187 Frankites regoledanus 244 Frankites sutherlandi 149, 245 Frechites 242

506 Frechites occidentalis 148, 149 Fu¨chsel, George Christian (1722– 1773), Muschelkalk 18 Fueloepicyrtis 171, 183, 187 Fundy Basin 341, 345, 351 conchostracans 350 cyclostratigraphy 128, 129 Galtonia 465 Ganderian substage, conchostracans 319– 320 see also Dienerian substage Gastrulocapsula 178 geomagnetic polarity timescale (GPTS) 4, 61–94 bar-code patterns 63 correlation 62–63 Geopollis 293, 297 Germanic Basin ammonoids 230 chronostratigraphy 8, 328 conchostracans 321– 322, 327–335, 357, 358– 359, 360, 368, 369 cyclostratigraphy 122, 123– 124 Gerrothorax 470 Gettysburg Basin, conchostracans 339, 340, 343 –344, 348–349 ginkgophytes 9 Ginsburg model 132 glacio-eustasy 131 Gladigondolella carinata 145 Gladigondolella malayensis 148 Gladigondolella tethydis 79, 145, 147 Gliscopollis 293, 297 global warming, carbon isotope excursion 107, 108, 109, 286 Globolaxtorum 182, 185, 189, 192 Glomeropyle 170, 190 Gnomohalorites 246 Gnomohalorites cordilleranus 251 Goestlingella 173 Gomberellus 175 Gondolella trammeri 73 Gondwana, conchostracans 323 Gonionotites maurolicoi 80 Goomboorian diorite, K– Ar age 42 Gordonispora fossulata 278, 292, 302 Gorgansium 182, 186, 193 Grallator 430, 432–434, 435, 436, 438, 439, 440, 470, 481 Granuloperculatipollis 293, 297 Granuloperculatipollis rudis 82, 267, 274, 275, 276, 288, 289, 292, 294, 304 zone 270, 275 Gregoriusella 333, 353, 355, 383, 392– 393 Gregoriusella bocki 382 Gregoriusella fimbriata 333, 382 Gregoriusella polonica 356, 360, 393, 394 zone 381, 388, 389

INDEX Grenzbitumenzone, Monte San Giorgio ammonoids 226 radio-iosotope dating 42, 46, 48, 49 Griesbachian substage 25, 27, 29, 30 ammonoids 232–234, 235 carbon isotope excursion 107, 108– 109 sulphur isotopes 108 groundwater storage 131 Gruipeda 435, 437 GSSP definitions 1– 3, 25, 26, 27, 30, 31, 32, 34, 139, 190, 193– 194, 229, 315 Guabisaurus 467 Guembelites 246 Guembelites jandianus 250 Gu¨mbel, Karl Wilhelm von (1823–1898), Alpine marine Triassic 20, 32 Guodikeng Formation conchostracans 324–327, 325, 364– 366 tetrapods 454 Guthoerlisporites cancellosus 289, 300, 302 Gwyneddichnium 437, 440 Gymnites 242 Gymnotoceras 242 Gymnotoceras rotelliformis 148, 149 Gyronites 140, 242 Haeckelicyrtium 180, 186, 193 Hagiastrum 182, 186 Halobia 201, 202, 203, 209, 210– 211, 214 Halobia asperella 209, 211 Halobia beyrichi 209, 211, 212 Halobia cordillerana 209, 211, 212 Halobia darwini 209, 211 Halobia distincta 209, 211 Halobia halorica 209, 211 Halobia lenticularis 209, 212 Halobia mediterranea 209, 211 Halobia multistriata 359 Halobia obruchevi 209, 211 Halobia ornatissima 209, 211 Halobia popowi 209, 211 Halobia radiata 209, 211 Halobia rugosa 209, 211 Halobia styriaca 211, 212 Halobia subfallax 211 Halobia superba 209, 211 Halobia talajaensis 211 Halobia zhilnensis 209, 211 Halobia zitteli 209, 210–211, 212 Halorites 246 Halorites macer 252 Haramiyavia 470 Hartford Basin 341, 345 conchostracans 350

Hauer, Franz Ritter von (1822–1899), Alpine marine Triassic 20 Hedenstroemia 234 Hedenstroemia hedenstroemi 66, 67, 68 Hegleria 170, 190 Heliosaccus dimorphus 267, 274, 289 zone 270, 274 Heliosaturnalis 178 Heliosporites reissingeri 276, 288, 293, 297, 303 Hemicycloleaia 362 Hemicycloleaia mitchelli 362–363 Hemiprotosuchus 473 Heptacladus 172, 183 Heptasaurus 460 Herrerasaurus 467 Hesperosuchus 464, 465, 469, 473 Hetalum 181 Hettangian Stage 3 carbon isotope excursion 113 conchostracans, zonation 380– 389, 381, 389–390 radiolarians 185– 187, 191– 193 see also Rhaetian–Hettangian boundary Hexacatoma 177 Hexaporobrachia 177, 188 Hexapylomella 179 Hexaspongus 174 Hexatortilisphaera 171 Himalayas, ammonoids 231 Himavatites 246 Himavatites columbianus 248 Himavatites hogarti 250 Hindeodus latidentatus 139, 142 Hindeodus parvus 63–65, 66, 67, 106, 139– 140, 142 Hindeodus praeparvus 139, 142 Hindeosphaera 174, 183 Hinedorcus 173, 183 Hollandites 242 Hornestheria 395, 397 Hornestheria sollingensis 369, 395, 396, 397 zone 375– 376 Howellisaura 321, 390 Howellisaura berryi 338 Howellisaura ovata 384 zone 381, 385 Howellisaura princetonensis 352, 384 zone 381, 384 –385 Howellisaura winterpockensis 342 Hozmadia 169, 171, 190 Huglusphaera 179, 188 Hungarosaturnalis 176, 187, 188 Hyatt, Alpheus (1838–1902), ammonoids 22 Hyd Group, U– Pb age 43 Hydasp Stage 29–30 hypercapnia 107, 108

INDEX Hyperodapedon 461, 462, 465, 466, 467, 468 Hypsognathus fenneri 470 Icarosaurus 465 ichnotaxonomy 421 Icrioma 180, 184, 192 Illinites 266 Illinites chitonoides 73, 266, 273 Illinites kosankei 273 Illyrian substage 30 India palynology 299 tetrapods 454, 471 Induan Stage 2, 10, 25, 27 ammonoids 232, 233 radio-isotope dating 45, 53 radiolarians 169, 170 see also Changhsingian– Induan boundary Induan– Olenekian boundary 315 conodonts 140, 143, 144, 145 magnetostratigraphy 66– 68 stable-isotope record 107, 108–109 Indusiisporites parvisaccatus 300 Infernopollenites 286, 289, 293, 299 Institisporites 266, 274 Ipswich Microflora 286, 287, 298, 300 Isalo Group tetrapods 467– 468 Isarcicella 139 Isarcicella staeschei 142 Isaura midlothianensis 342 Ischignathus 467 Ischigualastia 465, 467, 468 Ischigualastian LVF 468 Ischyosporites variegatus 276 Ishigaum 170 Isochirotherium 422, 423, 425, 427, 428, 429, 438, 439, 481 isotope stratigraphy see radioisotope ages; stable-isotope record Jachaleria 467, 468 ‘Jachaleria’ colorata 473 Japonites 143, 234, 242 Japonocampe 181 Joannites jacobus 243 Judicarites 253 Jugasporites 65, 264, 269 Julian substage 30, 31 Justium 180 Juvavian Stage 30, 31 Juvavites magnus 250, 251 K– Ar dating 42– 43, 46, 54 Kahlerosphaera 180, 184, 187 Kannemeyeria 449, 453, 455, 456– 457, 458, 459 Kapes 77, 80 Karnospongella 177

Karoo Basin PT boundary extinctions 71, 476 tetrapods 451 –452, 459 Kashmirites 234 Kashmirites densistriatus 45, 66 Katorella 173 Kayenta Formation, conchostracans 357 Kellnerites 228 Keuper 18, 329, 334 cyclostratigraphy 122, 124 early work 18–20 Keuperisporites baculatus 267, 274 Khramov, A.N., magnetostratigraphy 62 Kin Kin Beds, K–Ar age 42 Klamathites macrolobatus 248 Klausipollenites 65, 269, 300 Knollenmergel 330 –331, 334 Koipato rhyolite porphyry, Pb-alpha dates 42 Kozuricyrtium 180 Kraeuselisporites 264 Kraeuselisporites reissingeri 82, 89, 92, 292, 297 Krasiestheria parvula 382 Krystyn, L., ammonoids and conodonts 227 Krzyzanowskisaurus hunti 469 Kuehneotherium 470 Kuglerina meieri 278 Kulacella 174 Kungalaria 182, 185, 192 Kupferzella 460 Kuttycephalus 467 Kyrtomisporis 289, 292, 293 Kyrtomisporis ervii 278, 292, 302 Lacian substage 30, 31– 32 Ladinian Stage 2, 11, 30–31, 190 ammonoids 234, 239, 242, 244, 245, 254 correlation with Berdyankian LVF 475, 478 correlation with Perovkan LVF 475, 478 radio-isotope dating 47, 47, 52 radiolarians 171–178, 187 stable-isotope record 107, 109– 110 see also Anisian– Ladinian boundary Ladinian–Carnian boundary ammonoids 148, 151, 254 carbon isotope record 110 conodonts 148– 149, 151, 152, 153 magnetostratigraphy 74– 75, 76, 78 palynology 291–292 Ladinocampe 175, 187 Laevicaudata 317 Lagenella martini 289, 301 Lagerpeton 460

507 Land–mammal ages (LMAs) 450 land–vertebrate faunachrons (LVFs) Triassic 9, 360, 421, 448, 449, 450–480 correlation to marine SGCS 474– 481 Lardaroceras 73 Laricoidites intragranulosus 300 Latemar Massif, cyclostratigraphy 54, 77–78, 121, 125–126, 132 Latentifistula 170 Latiscopus 461 Laxitextella dorsorecta 353, 383, 384 Laxitextella freybergi 318, 338, 344, 359 zone 381, 383, 385 Laxitextella freybergi-P. schwanbergensis zone 381, 383 –384 Laxitextella laxitexta 333, 381, 382 Laxitextella multireticulata 318, 333 zone 380– 381 Laxitextella seegisi 352, 358, 384 zone 381, 382–383 Laxtorum 182, 186, 188, 193 Lehmann, Johann Gottlob (1719– 1767), Bunter 18 Leiofusa 269 Leiophyllites 234 Leptolepidites 302 Leptopleuron 467 Leptosuchus 464, 465, 468 Lessemsaurus 473 Liassobetraccium 187 Ligulatubus 176 Limbosporites antiquus 300 Limbosporites lundbladii 267, 268, 276, 292, 293, 294, 295, 297 Limitisporites 264, 269 Lioestheria 358 Lioleaiina 319, 333, 379 Lioleaiina jakutica 372 Lioleaiina radzinskii 368, 369, 373–374 zone 370, 373–374 Lioleaiina triasiana 372 Livarella 182, 184, 188, 193 Lobactinocapsa 174 Lockatong Formation conchostracans 344 cyclostratigraphy 128–130 Lofer cycles 126–128 Loffa 181 Longobardian substage 30– 31 Longobardisaurus 470 Longobardites 242 Longosuchus 461– 463 Lootsbergian LVF 421, 448, 449, 452–455 correlation with PT boundary 474–477 Loupanus 183, 185 Lucianosaurus 469

508 Lueckisporites 70, 71, 264 ‘Lueckisporites’ cf. singhii 278, 288, 292, 302, 305 Lueckisporites virkkiae zone 269, 270 Lunatisporites 264, 269, 291, 294 Lunatisporites acutus 302 Lunatisporites noviaulensis 289 Lunatisporites rhaeticus 288, 292 Lundbladispora 70, 71, 264 Lundbladispora denmeadi 300 Lundbladispora obsolete-P. pantii zone 269, 270 Lundbladispora willmottii 269 Luzocephalus 71, 452, 453, 454, 475, 477 Lycopodiacidites 267, 274, 293 Lycopodiacidites rugulatus 303 lycopsids 9, 264–265 Lydekkerina 452, 453, 454 Lyrocephaliscus 456, 475, 477 Lysemelas 182, 188 Lystrosaurus 324–326, 327, 363, 364, 365, 366– 367, 452 biochronology 448, 450, 452, 453, 454– 455 Maclearnoceras 242 Maclearnoceras maclearni 245 magnetostratigraphy 4, 61– 94, 91 Anisian–Ladinian boundary 73– 76 Carnian Stage 81– 82 Carnian– Norian boundary 79, 80– 81, 86–87 correlation 62–63 early work 61– 62 Induan–Olenekian boundary 66– 68 Ladinian–Carnian boundary 74– 75, 76 Lower Triassic 66– 71 nonmarine 70, 71–73 Middle Triassic GPTS 76–78 nonmarine 77, 78, 80 Newark Supergroup 86–88, 89, 90, 93, 94 Norian Stage GPTS 82–84 Norian–Rhaetian boundary 84– 86 Newark Supergroup 87–88, 90 Olenekian Stage 66–71 Olenekian-Anisian boundary 68, 69, 71 Permian–Triassic boundary 63– 66, 476 Triassic– Jurassic boundary 89, 90, 92–93 Magniestheria 358 Magniestheria deverta 369 zone 374– 375 Magniestheria deverta bogdoensis 374 Magniestheria deverta deverta 374

INDEX Magniestheria endybalica 372 Magniestheria lerichi 367, 369, 371– 372 Magniestheria mangaliensis 317, 318, 320, 367, 369, 371– 372, 374 zone 374 Magniestheria rybinskensis 368, 369, 373– 374 Magniestheria rybinskensis-L. radzinskii zone 370, 373–374 Magniestheria subcircularis 317, 368, 372 zone 370, 372– 373 Magniestheria truempyi 72, 317, 323, 368, 372–373 zone 370, 373 Malayites dawsoni 248 ‘Malerisaurus’ 463 Malschenberg Sandstone 330– 331, 334 Manassas Formation 342 ‘Mandasuchus’ 459 Marasuchus 460 Massetognathus 453, 460 Mastodonsaurus giganteus, biochronology 448, 453, 457– 458, 459–460, 475, 478 Meekoceras gracilitatis 237 Megagomphodon 460 Megalancosaurus 470, 475 Megasitum 324, 326, 362 Megasitum vanum 362 Meginoceras 242 Meginoceras meginae 245 Meishan GSSP PT boundary 44, 45, 65, 106, 139 carbon isotope excursion 107, 108, 475, 476 magnetostratigraphy 64, 65 Mentosaurus waltheri 478 Menucoestheria 323 Mesodapedon 459 Mesohimavatites 246 Mesosaturnalis 182, 186, 188, 193 Metapolygnathus 154, 155 Metapolygnathus acuminatus 152 Metapolygnathus communisti 79, 80, 153, 154, 155 Metapolygnathus echinatus 79, 80, 153, 155, 158, 191 Metapolygnathus intermedius 152, 153 Metapolygnathus nodosus 49, 79, 80 Metapolygnathus parvus 79, 80, 154, 155 Metapolygnathus polygnathiformis 148– 149, 152, 153 Metapolygnathus primitius 153–154, 155, 191, 356 Metapolygnathus tadpole 152, 153 methane hydrate dissociation 104, 107, 108, 109, 112–113 Metoposaurus 383, 461– 463, 467, 468, 475, 478

Micrhystridium 269 Milankovitch cycles 4, 119– 120, 125, 129, 130, 132– 133 Minutosaccus 286, 299, 300 Minutosaccus crenulatus 289 miospores Norian– Rhaetian 292–298, 303, 304 Upper Triassic, southern hemisphere 298– 300, 302, 305 Misikella hernsteini 85– 86, 157, 158 Misikella posthernsteini 32, 84– 86, 157, 158, 315 Moenave Formation, conchostracans 356–357 Moenkopi Group conchostracans 351–352 nonmarine magnetostratigraphy 70, 72 Mojsisovics, Edmund von (1839–1907) ammonoids 222, 223, 224– 225, 253 Old World Triassic timescale 20, 21, 22, 29– 32, 35, 224, 225 Molinestheria 319 Molinestheria seideli 367 –368, 368 zone 367– 368, 370 Molinestheria seideli postera 369, 371 Molinestheria seideli seideli 367, 371 Molzaxis 172 Monicasterix 172, 183 Monocapnuchosphaera 181, 188 Monospongella 173, 183 Monostylosphaera 178 Monosulcites 300 Monotis 201, 202, 203, 204, 209, 211, 214 Monotis ochotica 209, 211 Monotis rhaetica 209, 211, 212 Monotis salinaria 209, 211 Monotis subcircularis 209, 211, 212 monsoon circulation 131, 264, 267 Monzoni igneous complex, K –Ar age 43 Mosherella newpassensis 152, 153 Mostlericyrtium 180, 188 Mud, Induan– Olenekian boundary GSSP 66, 67, 140, 144 Muellericyrtium 172, 183 Muelleritortis 176, 187 Multiarcusella 175 Multimonilis 177, 187 Muschelkalk 18 conchostracans 333 cyclostratigraphy 122, 124 early work 18, 19–20 ‘Mussaurus’ 473 Myophoria vulgaris 358, 378 Myosaurus 454 Mystriosuchus 462, 470, 471, 475, 480

INDEX Nabolella 174, 184, 192 Nammalian Stage, definition 29 Nandartia 172 Nathorstites 242 Natraglia 180, 184 Nazarovella 178 Neara Volcanics, K– Ar ages 42, 43 Neoaetosauroides 473 Neocanoptum 182, 185, 189 Neogondolella 140, 150 Neogondolella aequidentata 148, 150 Neogondolella alpina 148, 150 Neogondolella carinata 140, 142 Neogondolella constricta 148, 150 Neogondolella cornuta 148, 150 Neogondolella discreta 140, 142 Neogondolella krystyni 140, 142 Neogondolella liebermani 150 Neogondolella nassichuki 140, 142 Neogondolella postcornuta 148 Neogondolella pridaensis 148, 150 Neogondolella regalis 145, 147 Neogondolella szaboi 148 Neogondolella taylorae 140, 142 ‘Neogondolella’ trammeri 148, 150 Neogondolella zhejiangensis 140, 142 Neohimavatites 246 Neophyllites 92 Neopylentonema 173, 183 Neoraistrickia densata 300 Neoraistrickia taylorii 302 Neospathodus chii 143 Neospathodus concavus 143 Neospathodus cyclodontus 143 Neospathodus dieneri 66, 67, 68, 140, 145 Neospathodus homeri 69, 71 Neospathodus krystyni 66, 67 Neospathodus kummeli 66, 67 Neospathodus pakistanensis 66, 67, 140, 143, 145 Neospathodus posterolongatus 143, 145 Neospathodus spitiensis 143, 145 Neospathodus triangularis 69, 71, 146, 147 Neospathodus waageni 27, 66, 67, 140, 143, 145 Neospathodus waageni eowaageni 143 Neospathodus waageni waageni 143 Nevadisculites taylori 240 Nevadites 253– 254 Nevadites humboldtensis 148 Nevadites hyatti 148, 243 Nevadites secedensis 149, 254 Nevanellus 179 Nevesisporites 299 Nevesisporites limatulus 300 Nevesisporites vallatus 278, 292 New Haven Arkose 341, 345 New Zealand, Late Triassic palynology 299– 300

Newark Basin, conchostracans 341, 344–345, 349–350, 359, 360 Newark Supergroup 335–338, 336, 337 conchostracans 321–322, 338– 351 cyclostratigraphy 128 –130, 132 magnetostratigraphy 86– 88, 89, 90, 93, 94, 479 tetrapods 461 –462, 469–470, 472 Nicomedites 242 Nicomedites osmani 240 Nicoraella germanica 69, 71, 146, 147, 329 Nicoraella kockeli 69, 70, 71, 73, 329 Nicrosaurus 462, 470, 471 Nitrader 187 Nodocapnuchosphaera 181, 188 Nodotetrasphaera 177 Nodotrisphaera 177 Nofrema 172, 183 Nonesian LVF 448, 449, 453, 455–457 correlation with Olenekian Stage 475, 477 Norestheria 402, 403 Norestheria barnaschi 331, 402, 403–404 Norestheria barnaschi-S. mcdonaldi zone 386–387 Norian Stage 2– 3, 12, 25, 31–32, 80 ammonoids 239, 246, 248, 250, 251, 252 carbon isotope record 110–111 conchostracans 359–360, 385– 388 correlation with Revueltian LVF 475, 480 duration 355–356 GPTS magnetostratigraphy 82–84 radiolarians 171, 173 –182, 188 see also Carnian– Norian boundary Norian–Rhaetian boundary ammonoids 156 carbon isotope excursion 110 conodonts 156, 157, 158 magnetostratigraphy 84– 86 Newark Supergroup 87–88, 89, 90 nonmarine 86, 87 miospores, northern hemisphere 292– 298, 303, 304 Norigondolella navicula 79, 80, 153, 155 Norigondolella steinbergensis 157 Nuskoisporites 269 Nuskoisporites dulhuntyi 65, 264, 269 obliquity cycles 120, 125 Octosaturnalis 182, 188

509 Octostella 178, 184 Odenwaldia 453, 456, 457 Oertlispongus 175, 187 Olenekian Stage 2, 10, 25, 27 ammonoids 232, 233, 234 correlation with Nonesian LVF 475, 477 magnetostratigraphy 66–72 radio-isotope dating 46– 47, 53 radiolarians 169, 170–171 see also Induan–Olenekian boundary Olenekian –Anisian boundary 315 ammonoids 143 –144, 144, 146 conodonts 143– 148 magnetostratigraphy 68, 69, 71 Oncodella paucidentata 157 Onslow Microflora 286, 287, 298 Ophiceras 234 Ophiceras commune 27, 235 Oppel, Albert (1831–1865), ammonoids 222 Orbiculiformella 183, 185, 189, 192, 193 orbital forcing 119–120, 132, 133 Ornatisaturnalis 176, 187 Ornithosuchus 467 Oruatemanua 170 Osmundacidites 268, 300 ostracods 12 Otapiria 211 Otapiria alpina 211, 212 Otapiria dissimilis 211 Otapiria ussuriensis 211 Otischalkian LVF 448, 449, 461–463, 466 correlation with Carnian Stage 475, 478– 479 Otoceras 234 Otoceras boreale 27, 64, 65, 66, 235 Otoceras concavum 27, 29, 64, 65, 235 Otoceras woodwardi 27, 29 Otozoum 431, 434 Otynisporites eotriassicus 326 Ovalipollis 267, 286, 292, 293 Ovalipollis ovalis 289, 298, 303 Ovalipollis pseudoalatus 82, 87, 92, 268, 274, 276, 277, 278, 288, 289, 294, 302 Owl Rock Formation, cyclostratigraphy 130, 131 oxygen isotopes 104 –106 Middle Triassic 110 Pachus 181 Palaeestheria dorsorecta 360 palaeobiogeography, bivalves 204–206 Palaeochersis 473 palaeoclimate 131, 266, 268, 285 palaeoecology bivalves 203–204 conchostracans 317

510 palaeoenvironment ammonoids 255 palynoflora 266 –267, 268 palaeoflora 9, 12 see also palynoflora palaeogeography, Upper Triassic 285 Palaeolimnadia 333, 353, 359 Palaeolimnadia alsatica alsatica 369 E. albertii mahlerselli-P. alsatica alsatica zone 376–377 Palaeolimnadia alsatica detfurthensis 352, 369, 375, 376 Palaeolimnadia atsuensis 359 Palaeolimnadia baitianbaensis 389– 390 Palaeolimnadia cishycranica 364 Palaeolimnadia longmenshanensis 389 Palaeolimnadia machaolingensis 379 Palaeolimnadia mecsekensis 375 Palaeolimnadia nakazawai 359 Palaeolimnadia nodosa 352, 369, 375, 376 Palaeolimnadia schwanbergensis 338, 359 zone 381, 383–384, 385–386 Palaeolimnadia semicircularis 389 Palaeolimnadia triangularis 378 Palaeolimnadia wianamattensis 358 palaeomagnetic data see magnetostratigraphy Palaeosaturnalis 180, 185, 193 Palaeosemantis 172 Palaeospongisporis europaeus 278 Paleorhinus 87, 383, 463, 464 Palhindeolithus 170 palynoflora 264–268 Lower Triassic 265–266, 277 Middle Triassic 266– 267, 277 palaeoenvironment 266 –267, 268 Permian–Triassic boundary 264–265, 277 Triassic– Jurassic boundary 268, 286–287 Upper Triassic 267–268, 278 palynology Ladinian–Carnian boundary 291–292 Triassic 263–278 Triassic– Jurassic boundary 288–307 correlation 296–298 palynomorphs 6– 7, 10– 12 Carnian, northern hemisphere 288–292, 301, 302 Norian–Rhaetian, northern hemisphere 292–298, 303, 304 PT boundary extinction 70, 71 Upper Triassic, southern hemisphere 298–302 zones 10– 12, 268–278

INDEX palynostratigraphy Triassic, NW Europe 268 –278 correlation 277– 278 Triassic–Jurassic boundary 287– 288 Pangaea 204–206, 285 palaeoclimate 131 Pantanellium 181, 186, 188, 192, 193 Panthalassa 204– 206 Papiliocampe 180, 188 Paraceratites 242, 253 Paraceratites trinodosus 29, 240 Parachirotherium 427, 428, 429, 438, 440, 481 Parachirotherium postchirotheroides 427 Paracirculina quadruplicis 288, 291, 293, 301 Paracirculina scurrilis 288 Paracochloceras 246 Paracochloceras amoenum 251 Paracrochordiceras 30, 143, 234, 242 Paracrochordiceras asseretoi 240 Paracyclotosaurus 453, 458, 459 Paradanubites 234 Paradapedon 463 Parafrechites meeki 148, 149 Paragondolella excelsa 148 Paragondolella inclinata 152, 153 Paragondolella liebermani 148 Paraheptacladus 174 Parakannemeyria 453, 456, 458, 459 Paranevadites furlongi 148, 243 Paranevadites gabbi 148 Pararchaeospongoprunum 170, 190 Pararuesticyrtium 173, 183, 188 Parasepsagon 172, 183 Parasuchus, biochronology 448, 459, 461– 463, 465, 467, 468, 475, 478 Paratriassoastrum 180, 184, 192 Paratriassocampe 172, 183 Paratypothorax 462, 464, 465, 469, 470 Parentactinia 169, 171, 185, 190, 193 Parentactinosphaera 173, 183 Paroetlispongus 173 Paronaella 183, 185, 192, 193 Parotosuchus 70, 73, 453, 456, 457, 459, 475, 477 Parrishea 464, 465 Partitisporites 267, 274, 294 Partitisporites novimundanus 278, 288, 292, 305 Parvibrachiale 183, 185 Passaic Formation 341, 344, 479 cyclostratigraphy 128, 129 Patinasporites densus 278, 288, 289, 291, 292, 293, 298, 301, 305 Patinasporites toralis 294 Patrulius 172, 183, 187

Paurinella 170, 190 Pegoxystris 170, 190 Pekinosaurus 465 Pelsonian substage 30 Peltostega 456, 475, 477 Pentabelus 171 Pentactinocapsa 172, 183, 187 Pentactinocarpus 176, 184, 192 Pentactinorbis 174, 183 Pentaspongodiscus 173 Peribositria 201, 203, 206, 207– 208, 214 Peribositria backland 208 Peribositria borealis 206, 208 Peribositria mimer 206, 208, 208 Peribositria sibirica 206, 208 Perinopollenites elatoides 297, 300, 305 Perinosporites thuringiacus 267, 293 Permian-Triassic boundary carbon isotope excursion 106– 108 conchostracan zonation 325, 361– 367 conodonts 139– 140, 141, 142 correlation with Lootsbergian LVF 474–477 magnetostratigraphy 63–66, 476 mass extinction 474 radiolarians 189 palynoflora 264– 265, 277 radio-isotope age 44– 45, 53 sulphur isotopes 108 Peromonolites 267, 268 Peromonolites elatoides 276 Perotrilites minor 266, 273 Perovkan LVF 448, 449, 453, 455, 457–459 correlation with Anisian-Ladinian Stages 475, 478 Pessagnollum 171 Peteinosaurus 470, 475 Petrified Forest Formation 355– 356 tetrapods 463, 464, 465, 468, 469 Phormedites 246 Phylloceras 92 Pinacoceras metternichi 31, 32 Pinuspollenites minimus 288, 292, 297 Pisanosaurus 467 Placerias 461– 463, 464, 465 biochronology 449 Placites 246 Plafkerium 174, 184 Plagioscutum 458 Plagiosternum 458, 460 Plagiosuchus 460 Planispinocyrtis 171, 183 Plateosaurus 330, 331, 470 biochronology 449 Platycuccoceras Beds 329 Platysaccus 300 Platysaccus leschikii 273 Playfordiaspora velata 299

INDEX Pleuromeia 264 Plinthogomphodon 461 Ploechingerella 178, 188 Podosporites amicus 266, 267, 274, 275 polarity, geomagnetic see geomagnetic polarity timescale pollen 265, 266–268 see also palynoflora; palynology; palynomorphs; palynostratigraphy Polycingulatisporites 268, 276 Polycingulatisporites crenulatus 299, 300 Polycingulatisporites mooniensis 300, 305 Polygrapta chatangensis 370 Polygrapta subovata 362 –363 Polypodiisporites ipsviciensis 302 Polypodiisporites polymicroforatus 268, 276, 277, 292 Poposaurus 461, 464, 465 Porcellispora longdonensis 267, 274, 276, 288 Postosuchus 462, 464, 465, 469, 473 Potonieisporites 269 Poulpus 175 Praecirculina granifer 267, 288, 289, 301 Praecitriduma 183, 185, 192 Praegomberellus 173 Praeheliostaurus 176, 187 Praehexasaturnalis 193 Praemesosaturnalis 182, 185 Praenanina 179 Praeorbiculiformella 179 Praeprotunuma 179, 188 precession cycles 120, 125, 129, 130, 131, 133 Predazzo igneous complex, 40 Ar/39Ar dating 44 Preondactylus 470, 475 primary production see bioproductivity crises Probainognathus 460 Probelesodon 460 Procolophon 452, 453, 454–455 Procolophonichnium 422, 423, 424– 425, 427, 430, 432, 437, 438, 439, 440 Proexeraetodon 467 Proganochelys 470 Prohungarites mckelvei 239 Prolacerta 452, 453, 454 Promastodonosaurus 467 Promyalina 207 Proparvicingula 182, 185, 189, 193 Proptychites 68 Proptychites candidus 29, 238 Prorotodactylus 422 Protecovasaurus 461 Proterochampsa 467 Proterochersis 470

Proterosuchus 454 Protochirotherium 423, 438, 439, 481 Protochirotherium wolfhagense 422, 423, 424 Protodiploxypinus 300 Protodiploxypinus doubingeri 266, 273, 274 Protodiploxypinus fastidiodes 266, 273 Protodiploxypinus gracilis 266 Protohaploxypinus 264, 269, 273, 294 Protohaploxypinus pantii see Lundbladispora obsolete-P. pantii zone Protomonocarina 379 Protosuchus 448, 452, 472, 474 Protrachyceras 243, 253 Protrachyceras gredleri 245 Protrachyceras longobardicum 245 Protrachyceras margariosum 245 Protrachyceras neumayri 245 Protuberum 460 Psephoderma 470 Pseudacanthocircus 183, 186, 189, 193 Pseudhesperosuchus 473 Pseudoenzonalasporites summus 288, 289, 293, 301 Pseudoeucyrtis 183, 185 Pseudogodia 178, 185 Pseudohagiastrum 182, 185 Pseudoheliodiscus 180, 186, 188, 193 ‘Pseudomonotis’ occidentalis 206 Pseudopalatus 462, 464, 469, 471, 474 Pseudosaturniforma 177 Pseudosepsagon 171, 183 Pseudostylosphaera 171, 183, 190 Pseudotetrasauropus 431– 432, 434 Pseudotriassocampe 172 Pseudotriconodon 469 Pseudovum 179 Psiloceras 51, 89, 90 Psiloceras pacificum 92 Psiloceras planorbis 89, 92 Psiloceras psilonotum 332 Psiloceras sampsoni 332 Psiloceras spelae 89, 90, 92, 111, 315 Psiloceras spelae tirolicum 251 Pterospongus 176, 187 Ptychites 243 Puesto Viejo Formation, K– Ar age 43 Punctatisporites leighensis 300 Punctatisporites triassicus 273 Punctatisporites walkomii 300 Pylostephanidium 175 Quadraeculina anellaeformis 268, 276, 292, 293, 294, 295, 297 Quadriremis 170 Quadrisaturnalis 180, 184, 188 radio-isotope ages 3, 41– 55 early work 42–44 Lower Triassic 45–46

511 Middle Triassic 46–49, 50, 52, 54 Permian–Triassic boundary 44– 45, 53 recent advances and problems 43– 44 Triassic– Jurassic boundary 51, 52 uncertainties 54–55 Upper Triassic 49– 51, 52 radiolarians 5, 163 –194 Anisian 170–174, 183, 187, 189–190 Anisian– Ladinian boundary 190 Carnian 171, 173– 181, 187 –188, 189–190 Carnian– Norian boundary 191 genera 169, 170–183 genera revision project 168–169 Hettangian 185– 187, 191 –193 Induan 169, 170 Ladinian 171– 178, 187 mass extinction 189 Norian 171, 173– 182, 188 Olenekian 169, 170– 171 Permian–Triassic boundary, extinction 189 problems at Triassic boundaries 189–194 Rhaetian 171, 174, 176, 178, 180–187, 188–189, 191–192 Triassic– Jurassic boundary 184–187, 191–192 zonation 10–12, 164, 165, 166, 167–168, 167 Radium 176 Rauisuchus 467 Re– Os dating 55 Recoaroella 174, 175 Redonda Formation 356, 360 Redondasaurus 360, 462, 469, 472–474, 480 biochronology 448 Redondasuchus 462, 472, 474 Redondavenator 472 Redondestheria 318, 320, 390 Redondestheria grovetonensis, zone 381, 387 –388 Redondestheria novomexicoensis 322, 356, 388 Reitziites 254 Reitziites reitzi 149, 190, 240 Relindella 173 Renila 178 Renzium 181 Reticulatisporites dolomiticus 278, 292, 302 Retisulcites perforatus 291 Retitriletes 268, 297 Retitriletes austroclavatidites 300 Retitriletes semimuris 288, 293 Revueltian LVF 448, 449, 462, 464, 469–472 correlation with Norian Stage 475, 480

512 Revueltosaurus 462, 469, 471–472 Revueltosaurus callenderi 464, 471 Revueltosaurus hunti 464 Rhabdoceras 246 Rhacophyllites 246 Rhaetavicula contorta 32, 158, 209, 212, 213 Rhaetian Stage 3, 12, 24, 25, 26, 32, 34, 84 ammonoids 236, 246, 251, 252 carbon isotope excursion 110–113 conchostracans 360– 361, 381, 388 correlation with Apachean LVF 475, 480 duration 34 magnetostratigraphy 87, 88, 89, 90 radio-isotope dating 52 radiolarians 171, 174, 176, 178, 180–187, 188–189, 191–192 see also Norian–Rhaetian boundary Rhaetian– Hettangian boundary, carbon isotope excursion 111– 113 Rhaetipollis 293 Rhaetipollis germanicus 89, 92, 267, 268, 275, 276, 278, 288, 292, 294, 303 zones 270, 275– 276, 295– 296 Rhaetipollis –Limbosporites zone 296– 297 Rhaetipollis –Porcellispora zone 296– 297 Rhaetogonyaulax rhaetica 89, 90, 92 Rhombocythere wicheri 331 Rhynchosauroides 422, 423, 424, 427, 430, 431, 432, 437, 438, 439, 440 Ricciisporites 267, 293 Ricciisporites tuberculatus 89, 92, 268, 275, 276, 277, 288, 289, 292, 294, 303 Ricciisporites umbonatus 267, 275, 276 Richmond Basin 339, 342 conchostracans 347 Richthofen, Ferdinand von (1833– 1905), Alpine marine Triassic 20 Rieber, H., ammonoids 226–227 Rieppelites shevyrevi 241 Rigalites 428, 430 Rikivatella 174 Rioarribasuchus 462, 464, 469 Riograndia 467 Riojasaurus 462, 472, 473, 474 Riojasuchus 473 Rogalkaisporites 300 Rogalskaisporites 268 Rohillites rohilla 66, 67

INDEX Rotodactylus 423– 424, 425, 428, 438, 439, 481 Ruesticyrtium 179 Rugaletes awakinoensis 300 Rugulatosporites 300 Rutiodon 448, 461, 462, 463, 464, 465 Rutiodon carolinensis 344 Rutiodon manhattanensis 344 Saccasaurus 467 Sagenites 246 Sagenites quinquepunctatus 252 Sagenites reticulatus 252 Saitoum 187 Saltopus 467 Samaropollenites 286 Samaropollenites speciosus 278, 286, 288, 289, 291, 292, 299, 302, 305 Sanfilippoella 179 Sanford Sub-basin, conchostracans 338, 339, 346 Santa Maria Formation conchostracans 358 tetrapods 467 Sarla 180, 187 Sasenisaurus 456, 475, 477 Saurosuchus 467 Scalenodon 453, 458, 459 Scalenodontoides 473 Scaloposaurus 452, 453 Scaphonyx 467, 468 Schilfsandstein 329, 333, 462, 467 Scleromochlus 467 Scottsburg Basin 339, 341 Scutispongus 177, 187 Scythian Series 26, 27, 28 sea-level, Triassic 121–122, 131– 132 seawater, stable isotopes 103–106, 109 sedimentation, carbonate platforms, cyclicity 132–133 Sedovia fecunda 378 seedferns 9, 266 Sellaspora rugoverrucata 278 Sellosaurus 470 Semiretisporis 267, 276, 292 Semiretisporis gothae 276, 292 Semiretisporis wielichoviensis 276 Senelella 180 Sepsagon 173, 183 Serilla 183, 185, 189, 192 Sertasaturnalis 178 Setalella 176 Sevatian substage 30, 32 SGCS, correlation with Triassic LVFs 474– 481 Shangsi section, Permian–Triassic boundary 44 magnetostratigraphy 64, 65 Shansiodon 453, 455, 457–458, 459 Shipingia 318, 356, 390, 401 Shipingia hebaozhaiensis 319, 320, 360 zone 381, 387

Shipingia mcdonaldi 331, 401, 403 N.barnaschi-S. mcdonaldi zone 381, 386–387 Shipingia olseni 318, 319, 322, 356, 387 zone 381, 387 –388 Shuvosaurus 469, 473 Siberia, ammonoids 231– 232 Siberian Traps carbon isotope excursion 107– 108, 109 radio-isotope dating 45 Silberling, N. J. (b. 1928), American timescale 23, 226 Silberlingtoides cricki 241 Silenticeras 242 Silesaurus 383, 467 Silicarmiger 174, 183 Sillosuchus 467 Simeonella brotzenorum alpina 332 Simeonella nostorica 332, 383 Sinokannemeyeria 453, 458, 459 Sirenites nanseni 247 Sloss, L.L., super-sequences 121 Smilosuchus 464, 465, 468 Smith, James Perrin (1864– 1931), ammonoids 22 Smithian substage 25, 29 ammonoids 232, 233, 234 stable-isotope record 107, 108– 109 Solling Formation 328–329, 332–333 Somerset Dam Gabbro, K–Ar age 44 Spath, Leonard Frank (1888– 1957) Triassic timescale 18, 22– 23, 226 work on ammonoids 225– 226 Spathian substage 25, 29 ammonoids 232, 233, 234 stable-isotope record 107, 108– 109 Spathicuspus 143 Spathicuspus spathi 145, 146, 147 sphenopsids 9 Spheripollenites 300 Sphingopus 427, 428, 429, 438, 439, 440, 481 Sphingopus ferox 427 Spinicaudata 317 Spinohollisella 176 Spinolobocyrtium 176 Spinomersinella 175 Spinopoulpus 180 Spinosicapsa 181, 186 Spinosuchus 465 Spinotriassocampe 173, 183 Spongolophophaena 172 Spongopallium 173 Spongosaturnaloides 178, 188 Spongoserrula 177, 187 Spongosilicarmiger 172, 183 Spongostephanidium 169, 171 Spongotortilispinus 177, 187 Spongoxystris 172

INDEX spores 264– 268 see also palynoflora; palynology; palynomorphs; palynostratigraphy stable-isotope record 103–113 Lower Triassic 106–109 Middle Triassic 107, 109– 110 Permian–Triassic boundary 106–108 Upper Triassic 107, 110– 113 stage boundaries 139, 140, 277– 278 stages, short v. long 34 Norian 355– 356 Stagonolepis 383, 462, 463, 464, 465, 467, 468 Stahleckeria, biochronology 449, 453, 460–461 Stampfliella 180 Stauracanthocircus 187 Staurikosaurus 467 Stauromesosaturnalis 187 Stauropylissa 174 Staurosaccites 299 Staurosaccites quadrifidus 291, 302 Steigerispongus 176, 187 Steinernes Meer, cyclostratigraphy 127 Stellapollenites 266 Stellapollenites thiergartii 73, 266, 273, 274, 277 zone 270, 273– 274 Stenaulorhynchus 459 Stenopoanoceras 234 Stereisporites 268 Stereisporites antiquasporites 300 Stikinoceras 246 Stikinoceras kerri 248 Stockton Formation 341, 344 cyclostratigraphy 129 Stolleyites 242 Striatella seebergensis 300 Striatoabieites 269 Striatoabieites balmei 267, 274, 278 Striatopodocarpites 264 Striatotriassocampe 175 strontium isotopes 104, 106 Strotersporites 264, 266, 269 Stubensandstein 331, 360, 467, 470 Subcommission on Triassic Stratigraphy (STS) timescale 25, 27, 32 nomenclature 25, 32– 33 sulphate, carbonate-associated 106 sulphur isotopes 106 Permian–Triassic boundary 108 Sweetospathodus kummeli 140, 142 Synaptichnium 422, 423, 425, 427, 428, 429, 438, 439, 481 Tamonella 170, 190 Tandarnia 171 Tanystropheus 460 Tanytrachelos 465 Tanzania, Late Triassic palynology 300

Tarjadia 460 Tauridastrum 181 taxonomy 33 ammonoids 252– 254 conchostracans 317–319 Corollina-Circulina-Classopollis complex 267, 306 see also ichnotaxonomy Taylorsville Basin 339, 342 conchostracans 347 Tecovasaurus 465, 469 Tecovasuchus 468 Teratosaurus 383, 467, 470 Tertrema 456, 475, 477 Tethyan marine fossils 34–35 ammonoids 230– 231, 238, 236, 240, 244, 249, 250, 252 conchostracans 323–324 Tethys Ocean 204– 206 Tetracapnuchosphaera 177 Tetragonias 449 Tetrapaurinella 170, 190 tetrapods biostratigraphy and biochronology 9, 447– 481 early work 447–450 footprints 7, 419– 440 biochronology 438–440 biostratigraphy 480– 481 Lower Triassic 421–422 Middle Triassic 422 –430 Upper Triassic 430–438 ichnotaxa 421 PT boundary extinction 70, 71–72 Tetraporobrachia 178, 184, 192 Tetrarchiplagia 174 Tetrasauropus 431, 434 Tetraspinocyrtis 173, 183 Tetraspongodiscus 178 Tetratholura 172 Thaisphaera 170, 190 Thecodontosaurus 470 Therapsipus 424, 427 Thisbitidae 246 Thrinaxodon 452, 453, 454 Thurstonia 186 Tiborella 169, 171, 190 Tigjanium borchgrevinki 378 Tipperella 183, 186, 189, 193 Tirodella 178 Tirolites 27, 68 Tirolites cassianus 70, 73 Tirolites spinosus 70, 73 Todisporites 266, 278 Todisporites marginales 278 Toogoolawah Group, K– Ar ages 42 Tosapecten efimovae 209, 212, 213 Tozer, E. Timothy (b. 1928), American timescale 23–25, 27– 29, 35, 226 Tozerium 186 Trachyceras 246, 253, 254

513 Trachyceras aon 31, 148, 249 Trachyceras aonoides 31, 249 Trachyceras desatoyense 149 Trachyceras longobardicus 21 Trachyceras pseudo-Archelaus 21 Trachysagenites 246 Trachysporites 292 Trachysporites fuscus 276, 288, 303, 305 Trachysporites–Heliosporites assemblage zone 296–297 Trachysporites–Pinuspollenites assemblage zone 296–297 Trachysporites–Porcellispora assemblage zone 296–297 Trans-Caucasia, ammonoids 231 Transdanubian Range, cyclostratigraphy 127 Traversodon 460 Trematosaurus 453, 456, 477 Trematosuchus 453, 456, 457 Trexus 187 Triadispora 77, 78, 266, 267, 273, 274, 289, 293, 294, 300 Triadispora crassa 266, 267, 273, 277, 278 Triadispora crassa– Verrucosisporites zone 270, 273 Triadispora obscura 82 Triadispora plicata 266, 273 Triadispora verrucata 267, 274, 275 Triadosphaera 181 Triaenosphaera 170 Trialatus 179, 188 Triancoraesporites 267, 268, 293 Triancoraesporites ancorae 276, 297 Triancoraesporites reticulatus 276, 292 Triarcella 181, 188 Trias, work of Alberti 18, 19–20 Triassic, Lower 25, 26–29 ammonoids 232 –234, 235, 236, 237 zonation 206, 234 bivalves 206, 207 –208, 208 conchostracans 357– 358 zonation 367–376 magnetostratigraphy 66–71 nonmarine 70, 71–73 palynoflora 265– 266, 277 stable-isotope record 106– 109 tetrapod footprints 421–422 U– Pb dating 45– 46 Triassic, Middle 25, 29–31 ammonoids 234 –236, 240–241, 243, 244, 245 zonation 209, 243, 246 bivalves 208, 209, 210 conchostracans 358 zonation 376–380 magnetostratigraphy 76–78 nonmarine 77, 78, 80 palynoflora 266– 267, 277

514 Triassic, Middle (Continued) radio-isotope dating 47– 49, 50, 54 stable-isotope record 107, 109–110 tetrapod footprints 422– 430 Triassic, Upper 25, 31–32 ammonoids 239, 243, 246, 247–252 zonation 209, 246, 251 bivalves 209, 210 –211, 212, 213 conchostracans 358– 361 zonation 380–389 miospores 292 –298 southern hemisphere 298– 300, 302, 305 palaeoclimate 285 –286 palaeogeography 285 palynoflora 267 –268, 278 radio-isotope dating 49, 51, 52 stable-isotope record 107, 110–113 tetrapod footprints 430– 438 Triassic–Jurassic boundary 296, 315 carbon isotope excursion 111–113, 286 GSSP 193–194 magnetostratigraphy 89, 90, 92– 93 mass extinction, radiolarians 189, 192 palynology 268, 286–307 correlation 296–298 radio-isotope dating 51 radiolarians 184– 187, 191–194 Triassistephanidium 174 Triassoastrum 179 Triassobipedis 172, 183 Triassobullasphaera 174 Triassocampe 171, 183, 188 Triassocingula 177, 187 Triassocrucella 180, 184, 188 Triassocyrtium 173, 183 Triassospathodus homeri 144, 146, 147 Triassospongocyrtis 169, 170, 190 Triassospongosphaera 175 Triassothamnus 173 Tricornicyrtium 182, 184 Trilophosaurus 465 Trilophosaurus buettneri 461, 464 Trilophosaurus jacobsi 464 Trimiduca 179 Trinodus 326 Tripedocassis 172, 183, 187 Tripedocorbis 169, 170, 190 Tripedurnula 172, 183, 187 Tripemphigus 324, 326, 362 Tripemphigus khovorkiliensis 362 Tripemphigus minutus 362 Tripemphigus sibiricus 362 Trirachodon 453, 456, 457, 459 Trisauropodiscus 437

INDEX Tritortis 176, 187 Tropites 246 Tropites dilleri 250, 251 Tropites subbullatus 31, 250 Tropites welleri 247 Trossingen Formation 330–331, 334, 470 Tsugaepollenites oriens 266, 273, 274 Tsugaepollenites pseudomassulae 288, 294, 295, 303 Tubospongopallium 175 Tubotriassocyrtis 175 Tuchodiceras poseidon 245 Tupilakosaurus 70, 71, 452, 453, 454, 475, 477 Turospongus 174 Tuvalian substage 30, 31 conchostracans 359, 382– 385 Typothorax 462, 471 Typothorax antiquum 464, 468, 471 Typothorax coccinarum 448, 463, 464, 469, 471 U– Pb dating 3, 44–45, 54 Lower Triassic 45–46 Middle Triassic 46– 49 Permian– Triassic boundary 44–45 Triassic–Jurassic boundary 51 Upper Triassic 49, 51 Uatchitodon 465 Udalia 186 United States of America ammonoids 230, 237, 238, 241, 243 Newark Supergroup, conchostracans 321–322, 338– 351 southwestern, conchostracans 351– 357, 354 Urals foreland basin, tetrapods 451– 452, 454, 456, 457, 459 Uvaesporites argenteaeformis 293, 303 Uvaesporites gadensis 278 Uvaesporites reissingerii 292 Vallasporites ignacii 267, 274, 278, 288, 289, 291, 292, 293, 294, 298, 301, 305 Vallesaurus 470 Van Houten cycles 129 Vancleavea 464, 472, 473 Vavilovites sverdrupi 29, 66, 67, 236 Veghia 179 Veghicyclia 178, 184 Verrucosiporites 265 see also Triadispora crassaVerrucosisporites zone Vertexia 318, 319 Vertexia tauricornis 367 –368, 368 Vertexia tauricornis tauricornis 367, 369, 371

Vertexia tauricornis transita 367, 369, 371 Verticiplagia 169, 171, 190 Veryhachium 269 Vesicaspora 300 Vesicaspora fuscus 302 Vileginea dorofeevi 378 Vileginea tuberculata 378 Vinassaspongus 175 Vittatina 71, 264, 269 volcanism CAMP 51, 112–113, 285, 286, 390 carbon isotope excursion 104, 107– 108, 109 Voltziaceaesporites heteromorphus 266, 273, 277, 300 Waagen, Wilhelm Heinrich (1841–1900), Triassic timescale 20, 22, 29, 224 Wadiasaurus 459 Wannerestheria 344, 399 Wannerestheria pennsylvanica 344, 384, 399, 400, 401 zone 381, 385 Wasatchites 29, 68, 70, 72, 234 Welirella 175 Wellesaurus 456, 457 Wellesaurus peabodyi 352 Wetlugasaurus 70, 71, 452, 453, 454, 475, 477 Weverella 179 Weylandites magmus 278, 292 Whitakersaurus 473 Wuranella 178 Wutonggou Formation, conchostracans 326 ‘Xenacanthus’ moorei 469 Xenoprotrachyceras 253 Xenorum 180, 188 ‘Xestorrhytias perrini’ 478 Xiangxiella acuta 379 Xiangxiella bicostata zone 377, 379 Ximolzas 171 Xipha 181, 188 Xiphothecaella 177 Xiyukannemeyeri 459 Yeharia 173, 183 Zanclodon silesiacus 478 Zebrasporites 267, 292, 293, 299 Zebrasporites interscriptus 268 Zebrasporites laevigatus 276 Zeldacria 172 Zhamojdasphaera 177 Zieglericonus rhaeticus 157 zircon, U– Pb dating 45, 46– 49, 54 Zupaysaurus 473

The Mesozoic Era begins with the approximately 50-million-year-long Triassic Period, a major juncture in Earth history when the vast Pangaean supercontinent completed its assembly and began its fragmentation, and the global biota diversified and modernized after the end-Permian mass extinction, the most extensive biotic decimation of the Phanerozoic. The temporal ordering of geological and biotic events during Triassic time thus is critical to the interpretation of some unique and pivotal events in Earth history. This temporal ordering is mostly based on the Triassic timescale, which has been developed and refined for nearly two centuries. This book reviews the state of the art of the Triassic timescale and includes comprehensive analyses of Triassic radio-isotopic ages, magnetostratigraphy, isotope-based and cyclostratigraphic correlations and timescale -relevant marine and non-marine biostratigraphy.