South Caspian to central Iran basins

Contents Acknowledgements

vi

Preface

vii

BRUNET , M.-F., GRANATH , J. W. & WILMSEN , M. South Caspian to Central Iran basins: introduction

1

MUTTONI , G., MATTEI , M., BALINI , M., ZANCHI , A., GAETANI , M. & BERRA , F. The drift history of Iran from the Ordovician to the Triassic

7

ZANCHI , A., ZANCHETTA , S., BERRA , F., MATTEI , M., GARZANTI , E., MOLYNEUX , S., NAWAB , A. & SABOURI , J. The Eo-Cimmerian (Late? Triassic) orogeny in North Iran

31

ZANCHETTA , S., ZANCHI , A., VILLA , I., POLI , S. & MUTTONI , G. The Shanderman eclogites: a Late Carboniferous high-pressure event in the NW Talesh Mountains (NW Iran)

57

GAETANI , M., ANGIOLINI , L., UENO , K., NICORA , A., STEPHENSON , M. H., SCIUNNACH , D., RETTORI , R., PRICE , G. D. & SABOURI , J. Pennsylvanian–Early Triassic stratigraphy in the Alborz Mountains (Iran)

79

FU¨ RSICH , F. T., WILMSEN , M., SEYED -EMAMI , K. & MAJIDIFARD , M. R. Lithostratigraphy of the Upper Triassic– Middle Jurassic Shemshak Group of Northern Iran

129

SHEKARIFARD , A., BAUDIN , F., SCHNYDER , J. & SEYED -EMAMI , K. Characterization of organic matter in the fine-grained siliciclastic sediments of the Shemshak Group (Upper Triassic –Middle Jurassic) in the Alborz Range, Northern Iran

161

WILMSEN , M., FU¨ RSICH , F. T. & TAHERI , J. The Shemshak Group (Lower –Middle Jurassic) of the Binalud Mountains, NE Iran: stratigraphy, depositional environments and geodynamic implications

175

FU¨ RSICH , F. T., WILMSEN , M., SEYED -EMAMI , K. & MAJIDIFARD , M. R. The Mid-Cimmerian tectonic event (Bajocian) in the Alborz Mountains, Northern Iran: evidence of the break-up unconformity of the South Caspian Basin

189

TAHERI , J., FU¨ RSICH , F. T. & WILMSEN , M. Stratigraphy, depositional environments and geodynamic significance of the Upper Bajocian– Bathonian Kashafrud Formation, NE Iran

205

EGAN , S. S., MOSAR , J., BRUNET , M.-F. & KANGARLI , T. Subsidence and uplift mechanisms within the South Caspian Basin: insights from the onshore and offshore Azerbaijan region

219

GREEN , T., ABDULLAYEV , N., HOSSACK , J., RILEY , G. & ROBERTS , A. M. Sedimentation and subsidence in the South Caspian Basin, Azerbaijan

241

ZANCHI , A., ZANCHETTA , S., GARZANTI , E., BALINI , M., BERRA , F., MATTEI , M. & MUTTONI , G. The Cimmerian evolution of the Nakhlak –Anarak area, Central Iran, and its bearing for the reconstruction of the history of the Eurasian margin

261

BALINI , M., NICORA , A., BERRA , F., GARZANTI , E., LEVERA , M., MATTEI , M., MUTTONI , G., ZANCHI , A., BOLLATI , I., LARGHI , C., ZANCHETTA , S., SALAMATI , R. & MOSSAVVARI , F. The Triassic stratigraphic succession of Nakhlak (Central Iran), a record from an active margin

287

WILMSEN , M., FU¨ RSICH , F. T., SEYED -EMAMI , K. & MAJIDIFARD , M. R. An overview of the stratigraphy and facies development of the Jurassic System on the Tabas Block, east-central Iran

323

Index

345

South Caspian to Central Iran Basins

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) PHIL LEAT (UK) NICK ROBINS (UK) JONATHAN TURNER (UK) Society Books Advisors

MIKE BROWN (USA) ERIC BUFFETAUT (FRANCE ) JONATHAN CRAIG (ITALY ) RETO GIERE´ (GERMANY ) TOM MC CANN (GERMANY ) DOUG STEAD (CANADA ) RANDELL STEPHENSON (NETHERLANDS )

Geological Society books refereeing procedures The Society makes every effort to ensure that the scientific and production quality of its books matches that of its journals. Since 1997, all book proposals have been refereed by specialist reviewers as well as by the Society’s Books Editorial Committee. If the referees identify weaknesses in the proposal, these must be addressed before the proposal is accepted. Once the book is accepted, the Society Book Editors ensure that the volume editors follow strict guidelines on refereeing and quality control. We insist that individual papers can only be accepted after satisfactory review by two independent referees. The questions on the review forms are similar to those for Journal of the Geological Society. The referees’ forms and comments must be available to the Society’s Book Editors on request. Although many of the books result from meetings, the editors are expected to commission papers that were not presented at the meeting to ensure that the book provides a balanced coverage of the subject. Being accepted for presentation at the meeting does not guarantee inclusion in the book. More information about submitting a proposal and producing a book for the Society can be found on its web site: www.geolsoc.org.uk.

It is recommended that reference to all or part of this book should be made in one of the following ways: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) 2009. South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312. BRUNET , M.-F., GRANATH , J. W. & WILMSEN , M. 2009. South Caspian to Central Iran basins: introduction. In: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) 2009. South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 1 –6.

GEOLOGICAL SOCIETY SPECIAL PUBLICATION NO. 312

South Caspian to Central Iran Basins

EDITED BY

M.-F. BRUNET CNRS-INSU and Universite´ Pierre et Marie Curie, France

M. WILMSEN Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Germany

and J. W. GRANATH Granath & Associates Consulting Geology, USA

2009 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 2009. 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, 90 Tottenham Court Road, London W1P 9HE. Users registered with the Copyright Clearance Center, 27 Congress Street, Salem, MA 01970, USA: the item-fee code for this publication is 0305-8719/09/$15.00. British Library Cataloguing in Publication Data A catalogue record for this book is available from the British Library. ISBN 978-1-86239-271-7 Typeset by Techset Composition Ltd., Salisbury, UK Printed by MPG Books Ltd., Bodmin, 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]

Acknowledgements The volume editors would like to acknowledge the following colleagues, who kindly helped with reviewing the papers submitted for this volume. N. Abdullayev M. Allen V. Bachtadse A. Baud M. Berberian D. Bosence M. Brookfield S. Egan M. Gaetani

C. Henderson J.-P. Houzay J. Ineson R. Littke B. Lombardo F. Lucazeau B. Natal’in A. Okay A. Robertson

A. Saidi K. Seyed-Emami M. D. Simmons H. Stollhofen J. Tait B. R. Turner T. Voigt

The Geological Survey of Iran is warmly thanked for its support of the fieldwork performed in Iran.

The following companies are thanked for their participation as sponsors to the MEBE (Middle East Basins Evolution) Programme and for their contributions towards colour printing costs.

Preface The Middle East Basin Evolution Programme (MEBE) was a 4-year consortium (2003–2006) funded by the major oil companies (BP, ENI, PETRONAS, SHELL and TOTAL) and by the French research organizations (INSU-CNRS and UPMC). This multidisciplinary study of the Middle East, spanning the Arabian-Peri-Arabian and Caucasian – Caspian areas, was led by E. Barrier (CNRS –Universite´ Pierre et Marie Curie, Paris, France) and M. Gaetani (University of Milan, Italy). Its focus was the geodynamic evolution of the area, particularly since the late Palaeozoic, and emphasized dating; providing considerable details on the regional kinematics, plate tectonic models and geodynamic evolutions of the area. MEBE brought together about 300 scientists from 28 countries, representing 100 universities and research organizations. In order to prepare regional syntheses, MEBE established eight working groups comprising MEBE participants and external regional specialists. The MEBE working groups were focused regionally (Zagros, South Caspian Basin –central Iran, Caucasus, Black Sea, Levantine and East Arabian margins) and on products (stratigraphic comparisons, lithospheric cross-sections). From 2003 to 2005, 26 scientific projects were funded in 14 countries of the Middle East, including the Black Sea, Caucasus, northern Iran, the Zagros, the Arabian margins and Levant domains. About half of the MEBE projects were located in Iran because of the pivotal importance of several regions of this country to both scientific and hydrocarbon interests, as well as the continuous support by the Geological Survey of Iran (GSI). Workshops were a key element in the MEBE Programme. The first one was held in Kiev in February 2006 by the Black Sea Working Group; the Caucasus Working Group met in Ankara in September 2006. A programme-wide workshop in Milan in December 2006 gathered the Zagros, East Arabian margin, South Caspian Basin–central Iran and Stratigraphic Comparisons MEBE Working Groups. The last workshop was held in Paris in December 2006 by the Levant Working Group. An important documentation of the MEBE activities was given during the 2007 EGU General Assembly in Vienna, where a special MEBE session was held. The results of the programme, as well as the MEBE GIS Database, were presented by way of 70 communications during the final MEBE meeting held at the Universite´ Pierre et Marie Curie (Paris) in December 2007. The principal MEBE products are an atlas of 14 palaeotectonic maps (published by the CGMW) showing the geodynamic and tectonic evolution of the Middle East

between the Late Triassic and the Present, and a series of four regional Geological Society of London Special Publications presenting the results of the regional MEBE working groups. The four volumes cover the Black Sea –Caucasus, the South Caspian–Central Iran, the Zagros–East Arabian margin and the Levant. This present volume deals with the geodynamic evolution of the region from the South Caspian Basin to North and Central Iran. The volume covering the Black Sea to the Caucasus, entitled Sedimentary Basin Tectonics from the Black Sea and Caucasus to the Arabian Platform, will be edited by M. Sosson (CNRS, University of Nice, France), N. Kaymakci (Middle East Technical University, Ankara, Turkey), R. Stephenson (Vrije Universiteit, Amsterdam, The Netherlands), V. Starostenko (Institute of Geophysics, National Academy of Sciences of Ukraine, Kiev, Ukraine) and F. Bergerat (CNRS –Universite´ Pierre et Marie Curie, Paris, France). The volume will present detailed results of new fieldwork, as well as syntheses of the tectonic evolution of the Black Sea –Caucasus area (Greater Caucasus, Lesser Caucasus, South and East Anatolia, margins of the Black Sea and the Black Sea) constrained by an integration of the newly obtained and previously published data. The volume entitled Tectonic and Stratigraphic Evolution of Zagros and Makran During the MesoCenozoic, and edited by P. Leturmy (University of Cergy-Pontoise, France) and C. Robin (University of Rennes 1, France), will present new data and results on sedimentology, stratigraphy, biostratigraphy, tectonics and kinematics in the Zagros fold belt and the adjacent Makran accretionary prism. The teams involved were sponsored either by MEBE or by other programmes. The fourth volume, Evolution of the Levant Margin and Western Arabia Platform Since the Mesozoic, edited by C. Homberg (Universite´ Pierre et Marie Curie, Paris, France) and M. Bachmann (University of Bremen, Bremen, Germany), will cover improvements in our knowledge of the tectonic, stratigraphic and environmental evolution of the Levant basin and its margins since the Mesozoic. In summary, the MEBE Programme provides a significant contribution of high-quality geological data at interregional scales and a considerable advance in the knowledge of the regional geology. These results will strongly contribute to new interpretations of the geodynamic evolution of the whole Middle East. Marie-Franc¸oise Brunet (MEBE Bureau member) Paris, June 2008

The Upper Triassic to Middle Jurassic succession of the Shemshak Group (up to 4000 m in thickness in the Alborz Mountains, Northern Iran) contains key information about the closure of the Palaeotethys Ocean, the rise and denudation of the Cimmeride Mountains, and the succeeding opening of the South Caspian Basin. Here at Emamzadeh–Hashem Pass (NW of Tehran, Iran), the Shemshak Group is embraced between Upper Palaeozoic–Middle Triassic (foreground) and Upper Jurassic carbonates (background). Photography by: Markus Wilmsen.

South Caspian to Central Iran basins: introduction MARIE-FRANC¸OISE BRUNET1 *, JAMES W. GRANATH2 & MARKUS WILMSEN3,4 1

UPMC Universite´ Paris 06, CNRS-INSU, UMR Tectonique, case 129, 4, place Jussieu, F-75005 Paris, France

2

Granath & Associates Consulting Geology, 2306 Glenhaven Drive, Highlands Ranch, CO 80126, USA

3

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany 4

GeoZentrum Nordbayern (GZN) Fachgruppe Pala¨oUmwelt, FAU Erlangen-Nu¨rnberg, Loewenichstrasse 28, D-91054 Erlangen, Germany *Corresponding author (e-mail: [email protected]) Abstract: The structurally and stratigraphically complex area of northern and central Iran holds the key to understanding the plate tectonic evolution of the South Caspian– Central Iran area. The closure of the Palaeotethys, the opening of the Neotethys, the rise and demise of the Cimmerian mountain chain, as well as the onset of Neotethys subduction and large-scale Neotethyan back-arc rifting all predated the formation of the more than 20 km-thick fill of the South Caspian Basin. This volume brings together work by specialists in different disciplines of the geosciences (tectonics, geophysics, sedimentology, stratigraphy, palaeontology, basin modelling and geodynamics) in order to elucidate the complex Late Palaeozoic– Cenozoic geodynamic history of the Iran area and the birth of the South Caspian Basin.

The South Caspian Basin (Fig. 1) is among the deepest sedimentary basins in the world, yet its timing of formation, subsidence pattern and history are poorly understood, as are the mechanisms responsible for its geometry and evolution. For example, the age of the oldest sediments of the South Caspian Basin is unknown due to their inaccessibility at great depth of burial under Upper Cenozoic sediments. Consequently, understanding of the pre-Neogene, and particularly the Mesozoic history, is dependent on inference from the geology of surrounding areas. One aim of the Middle East Basin Evolution (MEBE) Programme was to perform new fieldwork on the margins of the South Caspian Basin, where inversion during the Cenozoic stages of the Arabia–Eurasia collision has exposed the rock record in the Alborz, Koppeh Dagh and Binalud mountains in northern Iran, as well as the eastern extent of the Greater Caucasus. Because the eastern edge of the South Caspian Basin and its relationship with other Iranian basins is unknown, studies of the central Iranian blocks where the remnants of Mesozoic deep oceanic basins are preserved are key to understanding the complex geological evolution of north-central Iran and the greater South Caspian area.

MEBE South Caspian to Central Iran Basins Working Group This Special Publication contains mainly the results of the MEBE Programme by the South Caspian to Central Iran Basins Working Group (SCWG), led by Marie-Franc¸oise Brunet of CNRS – Universite´ Pierre et Marie Curie, Paris. Two additional papers (the second paper by Wilmsen et al. and Green et al.) were funded outside MEBE and complete the set of papers presented in this volume. Results of the programme were presented during a workshop held on 4–5 December 2006 in Universita` degli studi di Milano-Bicocca, Milan, Italy. Some of the papers were also presented in Session TS 10.3, ‘Middle East Basins Evolution’, of the European Geosciences Union (EGU) General Assembly in Vienna, April 15–20 2007. Some work funded by MEBE on the geodynamics of the Alborz and Koppeh Dagh (Brunet et al. 2007; Shahidi 2008) are not presented here but, instead, will be published elsewhere. The SCWG comprised six MEBE projects involving interdisciplinary groups from mainly nine European research institutes and universities, as well as the Geological Survey of Iran, University of Tehran and the Geology Institute of Azerbaijan.

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 1– 6. DOI: 10.1144/SP312.1 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Situation of the South Caspian area and Iran (topography: digital elevation model SRTM30 from NASA Shuttle Radar Topography Mission- 30 arc-second grids mosaic; http://www.jpl.nasa.gov.srt). Ag, Agdharband; Af, Afghanistan; Bi, Binalud Mountains; EC, Eastern Caucasus; M, Mashhad; T, Tehran; Ta, Talesh Mountains; TB, Tabas Block; LB, Lut Block; YB, Yazd Block. The numbers refer to the approximate areas concerned by the papers of this volume: 1, Muttoni et al. 2, Zanchi et al. (their first paper in this volume); 3, Zanchetta et al.; 4, Gaetani et al.; 5, Fu¨rsich et al. (their first paper in this volume); 6, Shekarifard et al.; 7, Wilmsen et al. (their first paper in this volume); 8, Fu¨rsich et al. (their second paper in this volume); 9, Taheri et al.; 10, Egan et al.; 11, Green et al.; 12, Zanchi et al. (their second paper in this volume); 13, Balini et al.; 14, Wilmsen et al. (their second paper in this volume).

General evolution of the area The geodynamic evolution and sedimentation of the area ranging from the South Caspian Basin–eastern Caucasus to Central Iran is mainly governed by the evolution of two subduction systems closing successively the Palaeotethys and the Neotethys oceans. To the south of the Palaeotethys, a long strip of microcontinental blocks rifted away from north of

Gondwana from the end of the Palaeozoic. Consisting of parts of Turkey and Iran, these blocks stretch to Southeast Asia, and are collectively known as Cimmeria, the Cimmerian continent or, preferably, simply the Cimmerian blocks as their original contiguity is questionable. The Permian–Middle Triassic advance and successive Late Triassic collision of some Cimmerian blocks with Laurasia along the north-directed

INTRODUCTION

subduction zone closed the Palaeotethys Ocean during the Triassic in the so-called Early or EoCimmerian orogeny defined by Stille (1910). It was followed by Neotethys subduction, culminating in the collision of Arabia and Eurasia in the Cenozoic. This geodynamic history can be subdivided into pre-Eo-Cimmerian (late Palaeozoic – Middle Triassic) and post-Eo-Cimmerian intervals (Late Triassic –Cretaceous), separated by the collision of the Iran Plate with the southern margin of the Turan Plate. The pre-Eo-Cimmerian interval comprises the opening and closure of the Palaeotethys. It marks the history of the detachment of the Iran Plate from the northern margin of Gondwana, its drift as a part of the Cimmerian continental collage northwards toward Laurasia, and the simultaneous opening of the Neotethys in its wake. The post-Eo-Cimmerian interval comprises the Eo-Cimmerian collision of the Iran and Turan plates, and events in its aftermath: the uplift and denudation of the Cimmerides; the onset of northwards-directed Neotethys subduction below the Iran Plate, which continued for some 170 Ma; and the development of large-scale back-arc basins along the Palaeotethyan suture. That late rifting stage opened the South Caspian Basin; it began as early as the Late Triassic in the Alborz margin for Brunet et al. (2007) and Shahidi (2008), whereas Fu¨rsich et al. (in their first paper in this volume) prefer an onset in Liassic times. Several compressional tectonic events occurred during the post-rifting period of thermal subsidence, especially in the Early Cretaceous and the latest Cretaceous –Palaeogene. These compressional phases are evidenced by regional unconformities and deposition of clastic sediments. In the Eocene, a strong extension, associated with east– west-trending normal faulting, originated the thick volcanoclastic Karaj Formation in the southern part of the Alborz and the thick series of Talesh (NW Iran and SE Azerbaijan). Brunet et al. (2007) and Shahidi (2008) assume that this regional submeridian extension is related to the back-arc opening behind the northwards subduction of the Neotethys oceanic lithosphere beneath the Eurasian margin. The first compressional deformations, related to the Arabia –Eurasia collision, started in the Late Eocene (Barrier et al. 2008a, b). This early stage of the collision involved only the Eurasian and Arabian margins, and lasted until the Early Miocene, after which continent– continent collision generated the main orogenic belts of the Middle East (i.e. the Caucasus –Alborz, southern Turkey and Zagros ranges). The thickest series in the South Caspian Basin is represented by the Pliocene–Quaternary, with up to 10 km of sediments deposited in less than 5 Ma. Accumulation occurred within the framework of

3

strong crustal shortening. The orogen’s uplift (Caucasus, Talesh, Alborz, Koppeh Dagh) surrounding the South Caspian Basin date to this time, isolating the South Caspian Basin. Their erosion provided an important source of sediment, filling up the subsiding South Caspian Basin and the foreland basins of the belts.

Main outline of the papers The first paper by Muttoni et al. discusses previously published palaeomagnetic data, as well as new Late Ordovician and Triassic palaeomagnetic data. They provide insights into the drift history of Iran as part of Cimmeria during the Ordovician –Triassic. Results indicate that in the Late Ordovician –earliest Carboniferous Iran was part of Gondwana, residing in the Late Permian and early Early Triassic at subequatorial palaeolatitudes, clearly disengaged from the parental Gondwanan margin in the southern hemisphere. Since the late Early Triassic Iran has been located in the northern hemisphere close to the Eurasian margin after crossing much of the Palaeotethys in approximately 35 Ma, at an average plate speed of about 7 –8 cm year21. Muttoni et al. argue that caution is warranted with the 1358 counterclockwise rotational model for central Iran with respect to Eurasia since the Triassic.

Northern Iran Zanchi et al. start their first paper with a complete account of the geological setting of the Alborz since the Precambrian before describing the Eo-Cimmerian orogenic structures resulting from the Late Triassic collision between Iran and the southern margin of Eurasia. Eo-Cimmerian structures are discontinuously exposed below the siliciclastic deposits of the Shemshak Group (Upper Triassic –Jurassic). In eastern Alborz (Neka Valley), south of Gorgan, the external part of the belt is preserved. Mesozoic successions overlie a pre-Jurassic Eo-Cimmerian thrust stack with a sharp unconformity. The stack is overthrusted southwards above a strongly deformed Upper Palaeozoic–Middle Triassic succession belonging to North Iran. Zanchi et al. also investigate the Shanderman Complex, exposed in the Talesh Mountains, western Alborz. According to their new data and analyses, Zanchetta et al. consider the Shanderman Complex as a metamorphic complex containing well-preserved eclogitic-phase assemblages and not as an ophiolitic fragment of the Palaeotethys Ocean along the Palaeotethys suture, as it was previously interpreted. New dating suggests an

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equilibration in high-pressure conditions of a deeply subducted continental crust during the late Carboniferous, tentatively in a Variscan orogenic event, thus pre-dating the Eo-Cimmerian orogeny by more than 100 Ma. Zanchetta et al. suggest that the Shanderman Complex represents a fragment of the Upper Palaeozoic European continental crust (Transcaucasia) stacked southwards on the northern edge of the Iran Plate during the Eo-Cimmerian orogeny, later to become finally exhumed at the end of the Eo-Cimmerian orogeny. Zanchi et al. also record the occurrence of an Eo-Cimmerian metamorphic event south of the Shanderman Complex, in Upper Palaeozoic slates and carbonates occurring below the Lower Jurassic Shemshak Group. Based on these new data, the Eo-Cimmerian structures exposed in the Alborz appear to be remnants of a collisional orogen consisting mainly of deformed continental crust without preservation of ophiolites. Gaetani et al. describe in detail the sedimentary succession from the Pennsylvanian (Carboniferous) to the Lower Triassic in the central and eastern Alborz on the basis of new fieldwork, and propose a regional synthesis based on sedimentary analysis and the collection of new palaeontological data. They discuss the main factors controlling the sedimentation: the global sea-level curve in the Moscovian and from the Middle Permian upwards; passive margin conditions; and change in the climate during the fast northwards drift of Iran blocks. It passed from a southern latitude through the aridity belt in the Middle Permian, across the equatorial humid belt in the Late Permian and reached the northern arid tropical belt in the Triassic. Fu¨rsich et al., in their first paper in this volume, give a complete new lithostratigraphic description of the Upper Triassic –Middle Jurassic Shemshak Group of northern Iran. This up-to-4000 m-thick siliciclastic unit is widespread across the Iran Plate of northern and central Iran; it was deposited after a drastic change in sedimentation following the deposit of the Lower–Middle Triassic platform carbonates of the Elikah and Shotori formations. The passage to the Shemshak Group is marked by karstification and bauxite/laterite deposits. The top of the group corresponds also to a sharp change from siliciclastic rocks to Middle–Upper Jurassic carbonates of a platform –basin system. In the Alborz, a Triassic and a Jurassic unit separated by an unconformity (in part angular in the north and less conspicuous towards the south) comprise the Shemshak Group. Fu¨rsich et al. distinguish two major facies belts approximately parallel to the strike of the Alborz. The northern one starts with deep-marine sediments (the Ekrasar Formation) representing prodelta –delta front environments, succeeded by fluvial–lacustrine, coal-bearing formations passing to the Jurassic conglomerates of

the Javaherdeh Formation. The southern facies belt begins with fluvial, coastal plain and shallow –marginal marine environments, followed by rocks of fluvial–lacustrine origin with coal. It continues with deep-marine shales and ends with sediments of deltaic –coastal plain origin. Fu¨rsich et al. interpret this lithostratigraphic scheme with the tectono-sedimentary evolution of the Shemshak Foreland Basin of the Alborz. A relict marine basin was gradually infilled in the north during the Late Triassic, whereas largely non-marine–marginalmarine conditions prevailed in the south. The basin was filled during the early Lias by erosional products of the uplifting Cimmerian orogen. Rapid subsidence prevailed during the Toarcian–early Middle Jurassic in the south, resulting in the deepening and subsequent infilling of a marine basin. Shekarifard et al. present results of bulk organic geochemical and microscopic studies carried out on the fine-grained sediments of the Shemshak Group in the Alborz. They show a dispersion of the organic matter that experienced high temperatures (Tmax values average 500 8C) during deep burial and active post-sedimentary tectonics. Total organic carbon (TOC) indicates a generally poor –moderate organic carbon content. The Lower Shemshak Group (Upper Triassic) was mainly deposited in marine/lake settings under dysoxic–anoxic conditions, and is an important effective petroleum source rock in the Alborz Range. In contrast, the Upper Shemshak Group (Toarcian –Aalenian) was deposited under comparatively deeper marine oxic–dysoxic conditions and has a low potential to produce petroleum. The Shemshak Group is also recognized in the Binalud Mountains, in NE Iran by Wilmsen et al. in their first paper in this volume. A Lower–lower Middle Jurassic non-marine sedimentary succession rests with angular unconformity on a metamorphic basement deformed during the Late Triassic Eo-Cimmerian orogeny. The overall succession with many conglomerates fines upwards owing to erosion, which eventually exposed metamorphic basement of a high-relief source area. Wilmsen et al. suggest that the strata are intramontane deposits of the Cimmerian Mountain chain in NE Iran, and that the position of the NW–SE-trending Eo-Cimmerian suture in NE Iran should be placed further SW than formerly assumed. An intra-Bajocian unconformity terminates the sedimentary cycle of the Shemshak Group. It is observed in many places all across the Iran Plate (Alborz Mountains of northern Iran, northeastern Iran, east-central Iran) and is sometimes called the Lutian unconformity (Seyed-Emami & Alavi-Naini 1990) or Mid-Cimmerian unconformity. In their second paper in this volume Fu¨rsich et al. describe the Mid-Cimmerian unconformity in the Alborz Mountains, northern Iran. A first

INTRODUCTION

distinct unconformity (near the Lower –Upper Bajocian boundary) is the sedimentary expression of rapid basin-shallowing due to uplift and erosion. Another unconformity is developed in the early Upper Bajocian, expressed as an angular unconformity due to block rotation, overlain by a thin transgressive conglomerate, followed by silty marls of the deep-marine Upper Bajocian –Callovian Dalichai Formation. This upper unconformity signals a rapid subsidence pulse. On the Tabas Block of east-central Iran (Wilmsen et al. in their second paper), a single unconformity (fusion of the lower and upper Mid-Cimmerian unconformities) can be documented. The Mid-Cimmerian tectonic event is thought now to be confined to the Bajocian. Fu¨rsich et al. explain this Bajocian phase as the expression of the change from the rifting- to the spreading-stage of the South Caspian Basin situated to the north of the present-day Alborz Mountains. In their interpretation, the Mid-Cimmerian event reflects a break-up unconformity, i.e. the onset of sea-floor spreading of the South Caspian Basin. Also along the southern margin of the Koppeh Dagh Mountains (northeastern Iran), one phase of uplift and erosion took place followed by a pronounced pulse of subsidence in the Late Bajocian that initiated the opening of the strongly subsiding Kashafrud Basin. This basin stretched in a NW– SE direction between the Koppeh Dagh and the Binalud Mountains, and is regarded as an eastwards extension of the South Caspian Basin (Brunet et al. 2007; Shahidi 2008; Taheri et al.). Taheri et al. describe the Kashafrud Basin, which rifted from the Late Bajocian onwards. It accumulated more than 2000 m of Upper Bajocian –Upper Bathonian siliciclastic sediments (termed the Kashafrud Formation) that vary rapidly in thickness and facies: basal non-marine alluvial fans and braided-river deposits grade into slope and deep-marine sediments at the top. The deep-marine shales of the Kashafrud Formation are effective hydrocarbon source rocks in the Koppeh Dagh.

South Caspian Basin subsidence Two papers, Egan et al. and Green et al. deal with subsidence modelling of the South Caspian Basin. They present different versions of models using similar sets of data and modelling software. Both papers conclude with a scenario of a Jurassic creation of thin and dense crust, probably of oceanic type, allowing the deposition of around 20 km of sediments in the main basin after thermal and flexural subsidence. Egan et al. combine fieldwork in the eastern part of Caucasus and offshore data provided by BP to investigate the role, as well as the timing, of the subsidence–uplift mechanisms that have affected

5

the Azerbaijan region of the South Caspian Basin from the Mesozoic to Recent. Model results support the tectonic scenario that South Caspian Basin crust has a density that is compatible with an oceanic composition and is being underthrust beneath the central Caspian region with an increase of the flexural subsidence in the northern part of the South Caspian Basin. Green et al., with a more complete oil exploration data set, give more emphasis on the Pliocene– Recent .7 km-thick sequence. Six kilometres of this sequence form the principal hydrocarbon reservoirs in the basin, the fluvial–lacustrine and deltaic sediments of the Pliocene Productive Series. According to Green et al., modelling shows that the deposition of this thick sequence can be explained without an additional Tertiary tectonic event, driven only by sediment loading and compaction on a thermally subsiding late Mesozoic crust. Their main conclusion is that deposition of the Pliocene Productive Series in a topographic depression was controlled by local factors, with a local base-level isolated from the global oceanic system, similar to the Mediterranean Messinian basin at the same time.

Central Iran Three papers deal with the central Iranian area. Zanchi et al., in their second paper, enumerate the models and the related blocks or terranes proposed in the literature to compose the Iran Plate. They discuss new fieldwork data and analyse results from the region of Nakhlak–Anarak (Central Iran). These structural, sedimentological, petrological and palaeomagnetic data provide important constraints on the Cimmerian evolution of Central Iran. Balini et al. give more details on the Triassic strata exposed at Nakhlak. They provide a lithostratigraphic revision of the succession on the basis of bed-by-bed sampling for ammonoids, conodonts and bivalves, as well as thin sections for limestone and sandstone petrographic analysis. Provenance for Olenekian–Anisian and Upper Ladinian sandstones is from a volcanic arc setting, while the sandstones of the ?Upper Anisian–Ladinian are related to exhumation and erosion of a metamorphic basement possibly located in the present-day Anarak region (see also the second paper by Zanchi et al.). The Anarak Metamorphic Complex was repeatedly deformed during post-Triassic times. The combination of various parameters (provenance, huge thickness, fast change of facies, abundance of volcaniclastic supply and orogenic calc-alkaline affinity of the volcanism) favours the interpretation that sedimentation of the Olenekian– Upper Ladinian succession occurred along an active margin, most probably in a forearc setting. The Nakhlak–Anarak units represent an arc–trench

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M.-F. BRUNET ET AL.

system developed during the Eo-Cimmerian orogenic cycle. Balini et al. make a comparison between the Triassic successions of Nakhlak and the one exposed in the erosional window of Aghdarband (Koppeh Dagh, northeastern Iran). The similarity pointed out by previous authors was an argument for the model of 1358 counterclockwise rotation of the central Iran blocks (Davoudzadeh et al. 1981; Soffel et al. 1996, among others), placing the Nakhlak area in a palaeoposition close to Aghdarband before rotation. According to Balini et al., these two successions, even if deposited along active margins, were possibly not deposited along the same margin. Indeed, their characteristics present more differences than similarities from a lithofacies point of view, and a preliminary palaeobiogeographical comparison is made based on new ammonoid data. Zanchi et al., in their second paper, discuss different tectonic scenarios that can account for the evolution of the region and for the occurrence of the Anarak orogenic wedge in its present position within central Iran, as well as its relationship to a presumed ‘Variscan’ metamorphic event. These results, added to the discussion of Muttoni et al. on palaeomagnetic data, lead Zanchi et al. and Balini et al. to conclude that large anticlockwise block rotations for Central Iran with respect to Eurasia since the Middle Triassic are doubtful. Timing, mechanism and amount of relative displacements between the different blocks of Iran are still obscure. Wilmsen et al., in their second paper in this volume, give an overview of the stratigraphy and facies development of the Upper Triassic –Jurassic sequences on the Tabas Block, east-central Iran. Unconformities related to the Cimmerian tectonic events subdivide the succession into major tectono-stratigraphic units. As in northern Iran, the Middle Triassic platform carbonates (Shotori Formation) are followed, with a drastic facies change, by the siliciclastic rocks of the Shemshak Group that reflect the onset of Eo-Cimmerian deformation in northern Iran. Marine sediments of the Norian –Rhaetian Nayband Formation pass to nonmarine, coal-bearing siliciclastic rocks around the Triassic– Jurassic boundary, related by Wilmsen et al. to the main uplift phase of the Cimmerian orogeny. Toarcian –Aalenian condensed limestones, indicating transgression, are followed by rapid lateral facies and thickness variations during the Lower Bajocian that culminate in the middle Bajocian compressional –extensional Mid-Cimmerian Event. The succeeding Late Bajocian –Bathonian onlap of the Baghamshah Subgroup was caused by increased subsidence of the Tabas Block, followed by the development of a large-scale platformto-basin carbonate system during the Callovian–

Kimmeridgian (Esfandiar Subgroup) which was destroyed by block-faulting that started in the Kimmeridgian (Late Cimmerian event). The same type of general evolution occurred here and in the Alborz Range of northern Iran, suggesting a common behaviour in the same structural unit. The data and models presented in this volume result in a detailed picture and a much more comprehensive understanding of the late Palaeozoic– Cenozoic geodynamic evolution of the South Caspian northern and central Iran sedimentary basins. They clarify some relationships of the region and add detail to its Palaeozoic–Mesozoic evolution. Notably, the work in the MEBE Programme will rearrange widely held notions of the early Mesozoic palaeogeography of this part of the Tethyan realm. We thank the staff at the Geological Society Publishing House for helping to oversee the project through publication, especially J. Pollitt (our production editor), J. Turner and A. Hills.

References B ARRIER , E., V RIELYNCK , B., B RUNET , M.-F., B ERGERAT , F. & S OSSON , M. 2008a. Toward a model of tectonic evolution of the Middle East since Mesozoic. In: Abstract International Geological Congress, Oslo, August 6– 14 2008. http://www. cprm.gov.br/33IGC/1343751.html B ARRIER , E. & V RIELYNCK , B. (Contributors: B ERGERAT , F., B RUNET , M.-F., M OSAR , J., P OISSON , A. & S OSSON , M.). 2008b. Palaeotectonic Maps of the Middle East. Tectono-Sedimentary– Palinspastic Maps from Late Norian to Piacenzian. Commission for the Geological Map of the World (CGMW)/UNESCO. (http://www.ccgm.org) Atlas of 14 maps, scale 1/18 500 000. B RUNET , M.-F., S HAHIDI , A., B ARRIER , E., M ULLER , C. & S AI¨ DI , A. 2007. Geodynamics of the South Caspian Basin southern margin now inverted in Alborz and KopetDagh (Northern Iran). Geophysical Research Abstracts, European Geosciences Union, Vienna, 9, 08080. http:// www.cosis.net/abstracts/EGU2007/08080/EGU2007J-08080.pdf D AVOUDZADEH , M., S OFFEL , H. & S CHMIDT , K. 1981. On the rotation of Central-East-Iran microplate. Neues Jahrbuch Geologie und Pala¨ontologie Monatshefte, 3, 180–192. S EYED -E MAMI , K. & A LAVI -N AINI , M. 1990. Bajocian stage in Iran. Memorie descrittive della carta geologica d’Italia, 40, 215 –221. S HAHIDI , A. 2008. Tectonic evolution of Northern Iran (Alborz and Kopet Dagh) since the Mesozoic. PhD thesis, Universite´ Pierre et Marie Curie, Paris (in French). S OFFEL , H. C., D AVOUDZADEH , M., R OLF , C. & S CHMIDT , S. 1996. New palaeomagnetic data from Central Iran and a Triassic palaeoreconstruction. Geologische Rundschau, 85, 293– 302. S TILLE , H. 1910. Die mitteldeutsche Rahmenfaltung. Jahresbericht des Niedersa¨chsischen geologischen Verein, 3, 141–170.

The drift history of Iran from the Ordovician to the Triassic GIOVANNI MUTTONI1*, MASSIMO MATTEI2, MARCO BALINI1, ANDREA ZANCHI3, MAURIZIO GAETANI1 & FABRIZIO BERRA1 1

Dipartimento di Scienze della Terra, Universita` di Milano, via Mangiagalli 34, I-20133 Milano, Italy, and ALP – Alpine Laboratory of Paleomagnetism, via Madonna dei Boschi 76, I-12016 Peveragno (CN), Italy 2

Dipartimento di Scienze Geologiche, Universita` di Roma-Tre, Largo San Leonardo Murialdo 1, I-00146 Roma, Italy

3

Dipartimento Scienze Geologiche e Geotecnologie, Universita` di Milano-Bicocca, Piazza della Scienza 4, I-20126 Milano, Italy *Corresponding author (e-mail: [email protected]) Abstract: New Late Ordovician and Triassic palaeomagnetic data from Iran are presented. These data, in conjunction with data from the literature, provide insights on the drift history of Iran as part of Cimmeria during the Ordovician– Triassic. A robust agreement of palaeomagnetic poles of Iran and West Gondwana is observed for the Late Ordovician–earliest Carboniferous, indicating that Iran was part of Gondwana during that time. Data for the Late Permian–early Early Triassic indicate that Iran resided on subequatorial palaeolatitudes, clearly disengaged from the parental Gondwanan margin in the southern hemisphere. Since the late Early Triassic, Iran has been located in the northern hemisphere close to the Eurasian margin. This northward drift brought Iran to cover much of the Palaeotethys in approximately 35 Ma, at an average plate speed of c. 7 –8 cm year21, and was in part coeval to the transformation of Pangaea from an Irvingian B to a Wegenerian A-type configuration.

According to Sengo¨r (1979), a strip of Gondwanan terranes called the Cimmerian Continent, which includes Iran, broke off the eastern Gondwanan margin during the Permian–Triassic, drifted northwards across the Palaeotethys in a windscreenwiper fashion and eventually collided with the Eurasian margin, giving birth to a conspicuous mountain chain, the Cimmerian orogen. This eminent and imaginative model of plates’ motion had a profound impact on Tethyan geology and was particularly appealing to palaeomagnetists insofar as it is testable by using the relationship between inclination of the time-averaged geomagnetic field and latitude. Indeed, quoting Van der Voo (1993), ‘the observed palaeolatitudes vindicate, in a very convincing way, the model of the Cimmerian continental motion of Sengo¨r & colleagues, at least insofar as Iran is concerned’, and this was due largely to the research efforts of Wensink et al. (1978), Wensink (1979, 1982, 1983) and, more recently, Besse et al. (1998). This premise would suggest that the drift model of Iran is grounded on a sound palaeomagnetic basis. However, in some instances, palaeomagnetically determined palaeolatitudes could be ambiguous for the reason that, without independent

estimates of polarity, the hemisphere of magnetization acquisition is undetermined. This ambiguity was resolved by Besse et al. (1998) in the Late Permian –Early Triassic by means of magnetostratigraphic correlations to sections of known polarity, but this exercise of polarity determination remains an exception in the Iranian database. Polarity (hemispheric) uncertainty can represent a problem when attempting to resolve the motion of Iran or other Cimmerian terranes because the Tethys Gulf was largely symmetric with respect to the palaeoequator. In this paper, new palaeomagnetic data from Iran are presented and a critical revision of palaeomagnetic data from the literature is attempted with the aim to refine the drift history of Iran over the Ordovician–Triassic.

Palaeomagnetic data Palaeomagnetic analyses were conducted on samples from sediments biostratigraphically constrained to the Late Ordovician, Late Permian, Early Triassic and late Early–late Middle Triassic (Olenekian–late Ladinian). Palaeomagnetic samples were drilled in the field with a water-cooled rock drill and oriented

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 7– 29. DOI: 10.1144/SP312.2 0305-8719/09/$15.00 # The Geological Society of London 2009.

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with a magnetic compass. All samples were thermally demagnetized and measured after each heating step on a 2G Enterprises dc SQUID (superconducting quantum interference device) cryogenic magnetometer located in a shielded room at the Alpine Laboratory of Paleomagnetism of Peveragno, Italy. An average of 50 8C steps from room temperature to a maximum of 680 8C were adopted. The component structure of the natural remanent magnetization (NRM) was monitored after each demagnetization step by means of vector end-point demagnetization diagrams (Zijderveld 1967), and steps of 25–10 8C were adopted close to critical unblocking temperatures. Magnetic components were obtained by standard least-square analysis (Kirschvink 1980) on linear portions of the demagnetization paths and plotted on equal-area projections. Standard Fisher (1953) statistics was applied to calculate site mean directions. Inclination-only statistics (McFadden & Reid 1982) were occasionally applied to sites characterized by declination scattering. The magnetic mineralogy of rocks was studied by means of isothermal remanent magnetization (IRM) acquisition curves, thermal decay of a three-component IRM (Lowrie 1990) acquired in 2.5 T (high), 0.5 T (medium) and 0.15 T (low) fields, or simply by

analysis of the maximum unblocking temperatures of the NRM.

Upper Ordovician Shirgesht Formation, Anarak A well-exposed Ordovician–Permian section crops out to the east of the town of Anarak at the edge of the Kavir loot (Sharkovski et al. 1984). The sampled site (Fig. 1, site IR12; 33.188N, 53.898E) is located at the base of lithological interval 3 of Sharkovski et al. (1984) and consists of strata dipping to the SW by 60 –708. They include red bioclastic sandy–silty limestones, silty marls and red shales pertaining to the first marine intercalation of the Shirgesht Formation. Schallreuter et al. (2006) attributed this interval to the Late Ordovician based on ostracoda biostratigraphy. Deformation in this general area took place broadly during the Mesozoic. The Neo-Cimmerian event is clearly recorded by a strong angular unconformity, metamorphism and the emplacement of magmatic bodies occurring between the Middle – Late Jurassic and the beginning of the Cretaceous. The possible occurrence of Late Triassic

Fig. 1. Structural sketch map of Iran with palaeomagnetic sites discussed in the text indicated. Modified after Angiolini et al. (2007).

DRIFT HISTORY OF IRAN

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Eo-Cimmerian compression is discussed by Zanchi et al. (2009b). A total of 19 samples have been cored in the field, yielding 28 standard (c. 11 cm3) specimens for palaeomagnetic analyses. Isothermal remanent magnetization acquisition curves show the occurrence of a high coercivity magnetic phase with a tendency to saturate only in fields higher than about 1.5 T (Fig. 2A, left panel). Thermal decay of the NRM shows that this phase consists essentially of hematite characterized by a narrow spectrum of maximum unblocking temperatures around approximately 660–680 8C (Fig. 2A, right panel). Most of the samples show the presence of pervasive A components with northerly and steep down directions in in situ co-ordinates (declination (Dec.) ¼ 2.18E, inclination (Inc.) ¼ 54.38, Fisher precision parameter (k) ¼ 79.9, Fisher angle of halfcone of 95% confidence about the mean direction (a95) ¼ 3.38) isolated between room temperature and approximately 175 8C (Fig. 3A, B and Table 1). These components are rather well defined (overall mean angular deviation (MAD) ¼ 2.28, standard deviation (St) ¼ 2.18) and are broadly aligned along a recent geocentric axial dipole (GAD) field direction (GAD inclination ¼ 528). In the temperature range between about 200 and c. 675 8C, characteristic C components are observed trending to the origin of the demagnetization axes with well-defined directions (overall MAD ¼ 4.78, St ¼ 2.18) oriented northeasterly and steeply down, which, upon correction for bedding tilt, become westerly and steeply down (Dec. ¼ 263.08E, Inc. ¼ 51.88, k ¼ 47.2, a95 ¼ 4.28; Fig. 3A, B and Table 1). No field test is available to constrain the age of these C components, which are nonetheless regarded as primary (Late Ordovician) based on palaeomagnetic poles analysis, as discussed in the next section.

c. 16 m of yellowish –brown laminated marls and fine-grained oolitic limestones. A laterally discontinuous 3–5 m-thick red–brown, fine-grained and massive bed overturned and dipping to the north by approximately 778 is present in this interval, and will hereafter be referred to as Aruh laterite. At the top, the residual profile grades into the marls and limestones of the basal Elikah Formation. Deformation in this general area took place essentially during the Cenozoic (e.g. Allen et al. 2003), although the effects of older deformational phases since the Late Triassic –Early Jurassic Eo-Cimmerian orogeny cannot be excluded. The studied site is located along the southern flank of a late Cenozoic anticline, which is crossed by the Mosha Fault, an active left-lateral strike-slip fault with large displacement (Allen et al. 2003). A total of 18 samples have been cored in the field in the Aruh laterite, yielding an equivalent amount of approximately 11 cm3 specimens for palaeomagnetic analyses. The Aruh laterite contains hematite of chemical origin that carries essentially two (and rarely three) magnetic components. Components oriented broadly along a recent GAD field direction in in situ co-ordinates dominate from room temperature up to about 250 8C. Characteristic C components (MAD ¼ 4.78, St ¼ 3.58) are isolated in the c. 350–620 8C temperature range. These are orientated essentially west and up in in situ co-ordinates, or NW and very shallow upon correction for bedding tilt (Dec. ¼ 331.88E, Inc. ¼ 21.38, k ¼ 18.5, a95 ¼ 10.48; Table 1) (see in Muttoni et al., submitted, for additional information). No field test is available to constrain the age of these C components, but, as for data from the Shirgesht Formation, we propose a primary (Late Permian) age of magnetization based on palaeogeographic considerations as discussed in the following section and in Muttoni et al. (submitted).

Upper Permian Aruh laterite

Lower Triassic Sorkh Shale, Tabas

A lateritic residual profile crops out near the Aruh village in the Alborz Mountains of north Iran (Fig. 1, site IR27; 35.668N, 52.408E). Age, chemical composition, palaeomagnetic properties and palaeogeographic implications of this residual profile and similar deposits from other Cimmerian units are discussed in Muttoni et al. (submitted). In brief, this profile is Late Permian in age, and is comprised between the carbonate platform limestones of the Ruteh Formation of Middle Permian (Capitanian) age and the carbonate platform limestones of the Elikah Formation of early Early Triassic (Induan) age (Gaetani et al. 2009 and references therein). The lowermost part of the profile consists of approximately 1 m of quartz-rich conglomerates and quartz arenites in a red shaly matrix, overlain by

Thin-bedded reddish oolitic–bioclastic limestones and marly limestones pertaining to the Sorkh Shale are exposed in the Shotori Range near the town of Tabas (Fig. 1, site SS). These sediments are Early Triassic in age (Bro¨nnimann et al. 1973) and have previously been studied for palaeomagnetism by Wensink (1982). In this same area we sampled two sites, namely IR34 (33.298N, 57.378E) and IR35 (33.378N, 57.258E), which yielded 19 and 21 standard c. 11 cm3 specimens for palaeomagnetic analyses, respectively. At both sites bedding attitude is highly variable, with strata dipping to the SE by approximately 60–858 at site IR34, and to the SSW by approximately 55 –758 at site IR35. Both sites are located along the northern branch of the North–South Nayband

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Fig. 2. Rock magnetic experiments on representative samples from this study. (A) Isothermal remanent magnetization (IRM) acquisition curve of sample IR12-14 from Upper Ordovician site IR12 (left panel) and histogram of maximum unblocking temperatures of the natural remanent magnetization (NRM) of samples from site IR12 (right panel). (B) Thermal decay of a three-component IRM (Lowrie 1990) acquired in 2.5 T (high), 0.5 T (medium) and 0.15 T (low) fields in samples from sites IR34 and IR35 from the Lower Triassic Sorkh Shale. (C) Thermal decay of a three-component IRM (same as in (C)) of samples from sites IR02 and IR04 from late Early Triassic (Olenekian) nodular limestones from Nakhlak; histograms of maximum unblocking temperatures of the NRM of samples from sites IR03 and IR17 (left panel), and IR01 and IR05 (right panel), from volcanoclastic sandstones from Nakhlak.

Fault, an active dextral strike-slip structure associated with complex folding. Thermal decay of the three-component IRM show that samples from both sites contain essentially hematite with maximum unblocking temperatures

approaching c. 680 8C; medium–high coercivity minerals with unblocking temperatures around 150–200 8C are also present and interpreted as iron hydroxide phases (Fig. 2B). Thermal decay of the NRM reveals the presence in most of the

DRIFT HISTORY OF IRAN

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Fig. 2. Continued.

samples from site IR34 of A component directions (MAD ¼ 5.68, St ¼ 2.18) isolated between approximately 100 and 320 8C. These are oriented northerly and steep down in in situ co-ordinates (Dec. ¼ 351.38E, Inc. ¼ 55.98, k ¼ 46.7, a95 ¼ 5.78), and are distributed broadly around a recent GAD field direction (GAD inclination ¼ 538). In samples from site IR35, similar components (MAD ¼ 5.48, St ¼ 2.18) with in situ GAD-aligned directions (Dec. ¼ 4.68E, Inc. ¼ 48.38, k ¼ 99.3, a95 ¼ 3.68) are present between room temperature

and approximately 250 8C (Fig. 4A, B and Table 1). Removal of these A components reveals the presence of characteristic C components isolated between approximately 450 and 650 8C at site IR34 (MAD ¼ 2.58, St ¼ 1.18), and approximately 525– 660 8C at site IR35 (MAD ¼ 2.68, St ¼ 1.38). The boundary between low-temperature A and hightemperature C components at both sites is frequently transitional, with directions moving along smallcircle paths. The C components at both sites show high degrees of dispersion, but, after correction for

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Fig. 3. Palaeomagnetic data from the Upper Ordovician Shirgesht Formation of site IR12. (A) Vector end-point demagnetization diagrams of samples IR12-01 and IR12-12B. Closed symbols are projections onto the horizontal plane and open symbols onto the vertical plane in in situ (geographical) co-ordinates. Demagnetization temperatures are expressed in 8C. (B) Equal-area projections before (in situ) and after bedding tilt correction of the A and C component directions. Closed symbols are projections onto the lower hemisphere and open symbols onto the upper hemisphere.

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Table 1. Palaeomagnetic mean directions from this study Comp.

T.unblock

N

In situ k

a95

MGDEC

Tilt corrected MGINC

k

a95

IR12 – Anarak (33.188N, 53.898E), Shirgesht Formation, Upper Ordovician A c. 030 –175 24 79.9 3.3 2.1 54.3 78.1 3.4 C c. 200 –675 26 52.6 3.9 35.0 58.2 47.2 4.2 IR27 – Aruh Laterite (35.668N, 52.408E), Upper Permian (Muttoni et al. submitted) A c. 100 –250 8 134.2 4.8 2.7 52.9 134.7 4.8 B1 c. 100 –325 4 70.3 11.0 299.9 81.9 70.3 11.0 C c. 350 –620 12 18.5 10.4 265.1 247.8 18.5 10.4 IR34 – Tabas (33.298N, 57.378E), Sorkh Shale, Lower Triassic A c. 100 –320 15 46.7 5.7 351.3 55.9 41.1 6.0 C c. 450 –650 12 9.4 21.0 + 15.6 25.3 IR35 – Tabas (33.378N, 57.258E), Sorkh Shale, Lower Triassic A c. 030 –250 17 99.3 3.6 4.6 48.3 58.7 4.7 C c. 525 –660 15 16.1 61.9 + 9.8 52.2 IR34 and IR35 combined (inclination only statistics) C c. 450 –660 27 4.3 43.4 + 15.2 36.5 IR01 – Nakhlak (33.538N, 53.848E), Alam Formation, Olenekian A c. 100 –450 17 30.6 6.6 10.8 45.6 30.5 6.6 B1 c. 450 –600 6 8.1 25.1 182.3 227.1 8.1 25.1 IR02 – Nakhlak (33.538N, 53.848E), Alam Formation, Olenekian A c. 030 –350 19 257.0 2.1 0.0 50.6 262.9 2.1 C c. 450 –675 11 35.8 7.7 167.2 74.9 33.0 8.1 IR04 – Nakhlak (33.588N, 53.838E), Alam Formation, Olenekian A c. 030 – 300 11 12.0 13.7 358.4 49.4 12.0 13.7 C c. 450 –665 14 38.2 6.5 184.5 24.7 33.9 6.9 IR02 and IR04 combined (inclination only statistics) C c. 450 –675 25 29.0 IR03 – Nakhlak (33.548N, 53.848E), Alam Formation, lower – middle Anisian A c. 100 –300 12 137.0 3.7 355.1 51.2 136.6 3.7 B1 c. 325 –510 4 84.3 10.1 17.2 220.5 84.2 10.1 IR17 – Nakhlak (33.548N, 53.848E), Alam Formation, lower – middle Anisian A c. 100 –300 18 54.2 4.7 6.4 47.5 45.0 5.2 B1 c. 250 –450 10 scattered B2 c. 450 –570 7 41.8 9.4 187.7 37.7 43.6 9.2 IR05 – Nakhlak (33.568N, 53.838E), Ashin Formation, upper Ladinian A c. 100 –600 19 117.4 3.1 2.2 55.3 56.8 4.5

MBDEC

MBINC

280.5 263.0

38.0 51.8

198.0 199.7 331.8

23.7 210.4 21.3

105.5

53.1 20.3 + 9.2

211.1

60.3 15.4 + 5.3 17.6 + 4.6

9.9 182.1

211.1 29.5

9.5 37.7

213.9 33.7

280.8 169.5

59.5 241.3 37.4 + 5.4

0.4 61.2

211.1 279.3

18.3

26.5

110.5

60.2

18.7

24.0

Comp., palaeomagnetic component; T.unblock, median values of component unblocking temperature (in 8C); N, number of standard 11 cm3 specimens; k, a95, Fisher (1953) precision parameter and angle of half-cone of 95% confidence about the mean direction, respectively; MGDEC, MGINC, mean geographical (in situ) declination and inclination, respectively; MBDEC, MBINC, mean bedding (tilt corrected) declination and inclination, respectively.

bedding tilt, their inclination values tend to cluster around a mean of 20.38 + 9.28 at site IR34 and 15.48 + 5.38 at site IR35, with the precision parameter k increasing by a factor of 2.7 and 3.2, respectively (Fig. 4B and Table 1). Combining C components from both sites results in a clustering of their inclination values around an overall mean of 17.68 + 4.68 after correction for bedding tilt, with the precision parameter increasing by a factor of 8.5 (Fig. 4C and Table 1). C components are therefore regarded as pre-folding in age. The dispersion in declination values that persists in bedding

co-ordinates is probably owing to uncertainties in the (simple) tilting correction applied and hence in the assumed deformation history of these highly deformed and fault-bounded outcrops.

Lower – Middle Triassic Alam and Ashin formations, Nakhlak A sedimentary sequence, late Early–late Middle Triassic (Olenekian–late Ladinian) in age, is exposed in the Nakhlak area of Central Iran (Fig. 1, site NAK; c. 33.558N, c. 53.848E). Six

14

G. MUTTONI ET AL.

Fig. 4. Palaeomagnetic data from the Lower Triassic Sorkh Shale of sites IR34 and IR35. (A) Vector end-point demagnetization diagrams of representative samples from both sites. Closed symbols are projections onto the horizontal plane and open symbols onto the vertical plane in in situ (geographical) co-ordinates. Demagnetization temperatures are expressed in 8C. (B) Equal-area projections before (in situ) and after bedding tilt correction of the A and C component directions at both sites; the C components are scattered in declination but show sign of clustering in inclination after correction for bedding tilt. Closed symbols are projections onto the lower hemisphere and open symbols onto the upper hemisphere. (C) Equal-area projections before (in situ) and after bedding tilt correction of the C component directions of both sites; an improvement in inclination clustering after correction for bedding tilt is evident. Symbols as in (B).

DRIFT HISTORY OF IRAN

sites (IR01, IR02, IR04, IR03, IR17 and IR05) have been sampled in stratigraphic order from the sequence base to top (Fig. 5A), yielding a total of 109 samples (115 standard, c. 11 cm3, specimens) for palaeomagnetic analyses. The sampled units belong to a geologically complex structure consisting of a WNW–ESE-trending fold-and-thrust belt. Strata at sampling sites generally dip to the NNE by approximately 508 –708, except for site IR04 where strata dip to the SW by approximately 508. Gently dipping strata of Upper Cretaceous –Palaeogene limestones and red sandstones unconformably cover the deformed Triassic succession (Zanchi et al. 2009a). Palaeomagnetic sites consist of volcanoclastic fine-grained sandstones and siltstones, except for sites IR02 and IR04 which consist of reddish nodular limestones. Sites IR01, IR02, IR04, IR03 and IR17 are located in the upper Lower– lower Middle Triassic (Olenekian –middle Anisian) Alam Formation. In particular, sites IR02 and IR04 are located in the Red Nodular Limestone of the Alam Formation, dated to the Olenekian using ammonoid biostratigraphy (Balini et al. 2009). Site IR05 is located in the lower part of the Ashin Formation of late Middle Triassic (late Ladinian) age (Balini et al. 2009). These samples contain magnetite and/or hematite in variable proportions. In samples from site IR02, hematite dominates with a narrow spectrum of maximum unblocking temperatures approaching approximately 680 8C, as visible in the thermal decay of the medium–high coercivity components of the IRM (Fig. 2C, samples IR02-15 and IR02-20). In samples from sites IR04, maximum unblocking temperatures consistent with both magnetite (c. 570 8C) and hematite coexist, as visible in the thermal decay of the low and medium–high coercivity components of the IRM, respectively (Fig. 2C, samples IR04-03 and IR04-17). Analysis of the maximum unlocking temperatures of the NRM seems also to suggest the coexistence of magnetite and hematite in samples from sites IR01 and IR05, whereas samples from sites IR03 and IR17 appear dominated by magnetite (Fig. 2C, lower panels). Palaeomagnetic components have been grouped into four categories labelled A, B1, B2 and C. Components A, B1 and B2 are regarded as postdepositional overprints of different origin and age, whereas the characteristic C components are tentatively regarded as primary (syn-depositional) in age, as discussed later. Most of the samples from all sites show the presence of pervasive A components (MAD ¼ 3.58, St ¼ 2.68) with northerly and steeply down directions in in situ co-ordinates, isolated from room temperature–100 8C to approximately 300–450 8C, occasionally up to about 600 8C (Figs 5B & 6 and Table 1). The A component

15

directions are clearly grouped around a recent GAD field direction in in situ co-ordinates (Fig. 5C). The overall mean of N ¼ 6 sites (Dec. ¼ 2.38E, Inc. ¼ 50.18, k ¼ 263, a95 ¼ 4.18) is statistically undistinguishable from the local GAD inclination of 538. Upon correction for bedding tilt, the sites’ mean directions become scattered, the overall precision parameter k reducing by a factor of 42 (Fig. 5C). Removal of these A components reveals the presence in samples from sites IR02 and IR04 of characteristic C component directions (MAD ¼ 7.08, St ¼ 3.98) isolated between approximately 400 and 675 8C, and trending to the origin of the demagnetization axes (Figs 5B & 6 and Table 1). At site IR04 these C components show dual polarity. The C component site mean directions are 99.78 apart along a great circle in in situ co-ordinates, and after correction for bedding tilt they tend to cluster, reducing their great-circle separation to 38.58 (Fig. 5D). Inclination-only statistics applied to C components from both sites resulted in a clustering of their inclination values around an overall mean of 37.48 + 5.48 after correction for bedding tilt (Fig. 5D and Table 1). Finally, B1 and B2 component directions were also observed in samples from sites IR01, IR03 and IR17 in the c. 250–600 8C temperature range after removal of the A components (Figs 5B & 6 and Table 1). At site IR01, the B1 components in in situ co-ordinates are quasi-antipodal to the in situ A components found in the same site at lower temperatures (Fig. 5B). These B1 components may thus represent recent overprints of reverse polarity. For samples from sites IR03 and IR17, the B1 components are more obscure and difficult to interpret. They are not compatible with recent field directions in in situ co-ordinates, and their site mean directions show no sign of clustering upon correction for bedding tilt. At site IR17, removal of these B1 components reveals the presence of B2 components with a tilt-corrected mean inclination value of 60.28 (a95 ¼ 9.28; Table 1) that is significantly different from, and hence apparently inconsistent with, the inclination values of the characteristic C components isolated at sites IR02 and IR04 (average inclination of 37.48). In conclusion, post-depositional remagnetization events of thermochemical origin pervasively overprinted all volcaniclastic sandstone and siltstone samples, generating recent GAD-aligned A components and possibly also scattered B1 and B2 components. Nodular limestones at sites IR02 and IR04 survived overprinting, showing apparently pre-folding C components that are possibly primary (syn-depositional) in age. The dispersion in their site mean declination values that persists after correction for bedding tilt is probably due to

16

G. MUTTONI ET AL.

Fig. 5. Stratigraphic and palaeomagnetic data of Triassic sediments from Nakhlak. (A) Stratigraphic log with location of palaeomagnetic sampling sites. (B) Equal-area projections before (in situ) and after bedding tilt correction of the A, B1, B2 and C component directions isolated in each site. (C) Equal-area projections before (in situ) and after bedding tilt correction of the A component site mean directions, showing an evident grouping along a recent geocentric axial dipole (GAD) field direction in in situ co-ordinates. (D) Equal-area projections before (in situ) and after bedding tilt correction of the C component site mean directions from sites IR02 and IR04, showing a relative improvement in inclination grouping after correction for bedding tilt. Closed symbols are projections onto the lower hemisphere and open symbols onto the upper hemisphere.

DRIFT HISTORY OF IRAN

17

Fig. 6. Vector end-point demagnetization diagrams of representative samples from sites IR01, IR02, IR03, IR04, IR05 and IR17 from Nakhlak with indication of magnetic components isolated. Closed symbols are projections onto the horizontal plane and open symbols onto the vertical plane in in situ (geographical) co-ordinates. Demagnetization temperatures are expressed in 8C.

post-folding differential rotation on vertical axes within the highly deformed Nakhlak structure.

The drift history of Iran Several studies have been conducted since the 1970s on Palaeozoic– Cenozoic rock units from Iran. Data from the literature summarized in Van

der Voo (1993) and Besse et al. (1998), augmented by data from this study, have been used to calculate north palaeomagnetic poles and palaeolatitudes for Iran from the Ordovician to the Triassic (Table 2). In our reconstructions, the various Iranian blocks, e.g. Alborz, Yadz, Tabas, Lut and Sanandaj– Sirjan (Fig. 1), have not been considered separately because the available palaeomagnetic results are too

Table 2. Summary of palaeomagnetic data from the literature and this study from Iran 18

Site IR12 RB GL

RL H IR27 HE

Rock unit Site IR12, Shirgesht Formation Red beds combined

Sorkh Shale, sites IR34 þ IR35 combined SS2 Sorkh Shale NAK2 Site IR02, Alam Formation NAK4 Site IR04, Alam Formation NAK Sites IR02 þ IR04 combined W Waliaband limestone, Abadeh SS1

Rel.age

33.18 53.89 Late Ordovician 30.5 57.2 Early Devonian 36 51.5 Earliest Carbon†

L.age H.age

N

ktc 1

a95tc Dec.tc Inc.tc NPlat.tc NPlong.tc A95,dp/dm

460.9 443.7

26

47.2

416.0 397.5

132

18.0 10.1

385.3 318.1

2

5 380.0

4.2 263.0

Wensink (1983)

3.9 211.0

67.0

0.0

212.0

5.3/6.5*

250

+5

Wensink et al. (1978)

–

22.6

19.3

289.9

1.9

212

+2

24.3 243.4

109.9

2.2/4.3

–

–

124.4

1.5/3.0

–

–

þ9?

+2.5 This study

35.66 52.40 Late Permian

260.4 249.7

121

18.5 10.4 331.8

L. Perm./ Induan

124.7

4.3 142.0

2.9 134.5 215.2 241.7

4.6

251.0 245.0

271

36.5

33.3 57.3 Early Triassic 33.53 53.84 Olenekian

251.0 245.0 249.7 245.0

52 111

30.0 14.2 288.9 33.0 8.1 37.7

33.58 53.83 Olenekian

249.7 245.0

141

Olenekian

249.7 245.0

251

216.5 210.0

1

52.9

Early Norian

–

23.7

33.33 57.31 Early Triassic

31.0

21.3

1

260.4 249.7 101

This study

+5

42.0

53.2

+4

21

31

30.9

232

5.1/10.1

260.4 251.0

52.0

3.9/5.7

196.3

1

32.0

175.0

51.3

131

–

Reference

21.3

270.6 260.4

52.2

errPlat

24.2

l. Mid Permian‡ Late Permian

36.3

51.8 212.0

Plat

–

Besse et al. (1998) 22 +2 Besse et al. (1998) 21/ þ 1 +5 Muttoni et al. (submitted) 28 +1.5 Gallet et al. (2000)

17.6

–

20.8 33.7

– 53.2

– 158.4

– 5.3/9.2

þ11?§ þ18

+8 +6

Wensink (1982) This study

33.9

6.9 169.5 241.3

76.5

279.5

5.1/8.4

þ24

+5

This study

29.0

5.4

–

37.4

–

–

–

þ21

+4

This study

35 266.9

4.6

–

51.9

–

–

–

þ32.5

+4.5 Besse et al. (1998)

Site, palaeomagnetic site abbreviation (in extended form under Rock unit); Slat, latitude of palaeomagnetic site in 8N; Slong, longitude of palaeomagnetic site in 8E; Rel.age, relative age of palaeomagnetic site (Induan is early Early Triassic, Olenekian is late Early Triassic and Norian is middle Late Triassic); L.age, older limit age of palaeomagnetic site in Ma according to the timescale of Gradstein et al. (2004); H.age, younger limit age of palaeomagnetic site in Ma according to the timescale of Gradstein et al. (2004); N, number of 1sites or 2samples; ktc, Fisher (1953) precision parameter of the palaeomagnetic direction in tilt corrected co-ordinates; a95tc , Fisher (1953) angle of half-cone of 95% confidence about the palaeomagnetic direction in tilt-corrected co-ordinates; Dec.tc/Inc.tc, declination/inclination in tilt-corrected co-ordinates of the palaeomagnetic direction; NPlat.tc/NPlong.tc, latitude (8N)/longitude (8E) of the palaeomagnetic North pole in tilt-corrected co-ordinates; A95,dp/dm, Fisher (1953) angle of half-cone of 95% confidence (A95) and oval of 95% confidence (dp/dm) about the palaeomagnetic pole in tilt-corrected co-ordinates; Plat, palaeolatitude (8N when positive, 8S when negative); errPlat, error of palaeolatitude expressed in degrees (8). *The oval of confidence reported by Wensink et al. (1978, table 2) is 3.98/5.38, but if calculated from the mean direction from table 1 in same paper (i.e. Dec. ¼ 2118E, Inc. ¼ 678, a95 ¼ 3.98) the oval of confidence is 5.3/6.5. † Wensink et al. (1978) attributed these lavas to the Upper Devonian –Lower Carboniferous but according to a recent study by Wendt et al. (2005), they are most probably earliest Carboniferous in age. ‡ The Ruteh lavas at Bear Gully are Capitanian (late Middle Permian) in age (Gaetani et al. 2009 and references therein). § Wensink (1982) reported a palaeolatitude of 118S, but because of the substantial lack of constraints on the polarity of magnetization of the Sorkh Shale, a palaeolatitude of 118N cannot be excluded; the same applies to sites IR34 and IR35 of this study.

G. MUTTONI ET AL.

Geirud Formation lavas, Alborz Mountains Ruteh lavas, Alborz Mountains Hambast Formation, Shahreza Site IR27 laterite, Alborz Mountains Hambast – Elikah formations, Abadeh

Slat Slong

DRIFT HISTORY OF IRAN

sparse to resolve potential palaeolatitude differences among them. In order to compare poles from Iran with coeval poles from Gondwana and Europe, apparent polar wander (APW) paths for West Gondwana (in NW Africa co-ordinates) and Europe have been constructed using selected data from the literature, as illustrated in Table 3. The APW path of Gondwana during the Palaeozoic has been the subject of much debate because some of the paleomagnetic poles for the Silurian and Devonian were derived from suspect terranes from Australia (McElhinny et al. 2003 and references therein). The Ordovician –Carboniferous Gondwana APW path adopted in this study is from McElhinny et al. (2003), whereas the coeval APW path for Europe is from Van der Voo (1993). The Permian–Triassic segments of the Gondwana and Europe APW paths are derived from palaeomagnetic poles of generally good quality (sensu Van der Voo 1993) and are listed in Table 3. Poles for West Gondwana have been rotated from NW Africa into Arabian co-ordinates using an Euler pole derived from Lottes & Rowley (1990) (26.228N, 11.158E, V ¼ 14.178), and have been plotted together with the European poles in Figure 7. These poles have been used to calculate Ordovician–Triassic palaeolatitudes expected at a nominal point located at 338N, 538E in central Iran for comparison with observed Iranian palaeolatitudes (Fig. 8). Data have been grouped in three time windows: namely, Late Ordovician–Early Carboniferous; Middle Permian–early Early Triassic; and late Early Triassic –Late Triassic.

Late Ordovician – Early Carboniferous: Iran as part of Gondwana Data from Iran pertaining to this time window come from the Upper Ordovician site IR12 of this study (Fig. 1 and Table 2; site IR12), the Lower Devonian Red beds of Wensink (1983) (site RB) and the Geirud lavas of Wensink et al. (1978) (site GL). The Geirud lavas were attributed to the Late Devonian –Early Carboniferous by Wensink et al. (1978), but, according to Wendt et al. (2005), they are bracketed between upper Famennian sediments below and upper Tournaisian –Visean sediments above, and are therefore probably earliest Carboniferous in age. The polarity of magnetization at these sites is unknown. Hence, their hemisphere of magnetization acquisition and north palaeomagnetic poles (defined by magnetizations in normal polarity reference frame) cannot be unequivocally determined. Polarity has been chosen in order to obtain the simplest tectonic scenario derived from comparison with coeval north poles from West Gondwana and

19

Europe. Reverse, normal and reverse polarity has been assigned to the mean magnetization of, respectively, the Upper Ordovician site IR12 (Dec. ¼ 263.08E, Inc. ¼ 51.88), the Lower Devonian Red beds (Dec. ¼ 24.28E, Inc. ¼ 21.38) and the lowermost Carboniferous Geirud lavas (Dec. ¼ 211.08E, Inc. ¼ 67.08). The derived palaeomagnetic north poles (Table 2) fall relatively close to, respectively, the Late Ordovician, Early Devonian and Early Carboniferous north poles of West Gondwana (Fig. 7A). This substantial congruence of poles (within palaeomagnetic error resolution) over a period of remarkable polar wander implies significant tectonic coherence of central Iran (sites IR12 and RB) and the Alborz region (site GL) with West Gondwana during this time interval. The sampled sites were located at a palaeolatitude of, respectively, approximately 328S (+48) during the Late Ordovician, approximately 18S (+58) during the Early Devonian and approximately 508S (+58) during the earliest Carboniferous, close to the Arabian margin of West Gondwana (Fig. 8, sites IR12, RB and GL, respectively). These overall conclusions concerning the Arabian affinity of Iran in the Palaeozoic substantiate previous conclusions by Wensink and colleagues (e.g. Wensink et al. 1978), and rule out the hypothesis that the Alborz region was located close to the southern margin of Eurasia during the Early Carboniferous as proposed by Kalvoda (2002) using palaeobiogegraphic data.

Middle Permian– early Early Triassic: Iran in motion Data from Iran pertaining to this time window come from four rock units: the Middle Permian Ruteh lavas from the Alborz Mountains (Besse et al. 1998) (Fig. 1 and Table 2, site RL); the Upper Permian Hambast Formation from Shahreza (Besse et al. 1998) (site H); the Upper Permian laterite of site IR27 from the Alborz Mountains (this study; Muttoni et al. submitted); and the Upper Permian– lower Lower Triassic Hambast and Elikah formations from Abadeh (Gallet et al. 2000) (site HE). The Ruteh lavas studied by Besse et al. (1998), equivalent to the Ruteh lavas studied by Wensink (1979), have been attributed to the Late Permian following stratigraphic considerations from the literature (Besse et al. 1998). In the wellstudied Bear Gully section, these lavas are interbedded in the uppermost part of the Ruteh Formation, biostratigraphically dated to the late Middle Permian (Capitanian) (Gaetani et al. 2009). A Capitanian age for the Ruteh lavas sampled by Besse et al. (1998) is therefore adopted in this study.

20

Table 3. Palaeomagnetic reference poles from West Gondwana and Europe Long. (8E)

Lat. (8N)

A95 or dp/dm (8)

GONDWANA in NW Africa co-ordinates; rotation poles of Lottes & Rowley (1990) Ordovician –Carboniferous Early Permian 242.0 41.4 Middle Permian 239.9 44.1 Late Permian–Early Triassic 237.6 46.8 Middle Triassic 218.1 57.1 Late Triassic –Early Jurassic; based on the following palaeomagnetic poles: Bolivar Dykes, Venezuela 206 74 Paramillos, Cach-Uspallata Argentina 223 67 Zambia Red Sandstone 211 72 Atlas & Meseta Volcanic, Morocco 216 71 Late Triassic –Early Jurassic 214.8 71.1 EUROPE Ordovician –Carboniferous Early Permian 166.2 42.2 Middle Permian 164.5 46.4 Late Permian–Early Triassic; based on the following palaeomagnetic poles: North Sudetic Sed. Zechstein 168 51 Intrasudetic Sed. Zechstein 160 51 Lower Buntsandstein 166 51 Buntsandstein Holy Cross 155 49 Saint-Pierre pelites 163 50

K

N

4.4

122

10

3.1 6.3

317 67

8 9

Table 3, McElhinny et al. (2003) See* See† Muttoni et al. (1996) Muttoni et al. (2001) MacDonald & Opdyke (1974) Creer et al. (1970) Opdyke (1964) Bardon et al. (1973) same as in Muttoni et al. (1996)

5 13 5 7 4.3

458

4

3.1

126

18

4/7 2/4 3 2 5

Reference

Table 5.7 in Van der Voo (1993) Muttoni et al. (2003, online table 2) See† Nawrocki (1997) Nawrocki (1997) Szurlies et al. (2003) Nawrocki et al. (2003) Diego-Orozco & Henry (1993)

G. MUTTONI ET AL.

Age

51 50 53 50.8

4 12 6 2.0

53 49 52 53 53.5 52.4

12 6 6 3.5 3 5.0

50 52 50 48 50.0

8 6.5 5 4 6.6

Merabet & Daly (1986) Cogne´ et al. (1993) Torsvik et al. (1998) 764

8 Symons et al. (1995) Edel & Duringer (1997) Nawrocki & Szulc (2000) Hounslow & McIntosh (2003) The´veniaut et al. (1992)

234

5 Edel & Duringer (1997) Hounslow et al. (2004) Briden & Daniels (1999) Hounslow et al. (2004)

198

4

Age, relative age of palaeomagnetic pole; Long. (8E), longitude of palaeomagnetic pole in 8E; Lat. (8N), latitude of palaeomagnetic pole in 8N; A95, dp/dm, Fisher (1953) errors about the palaeomagnetic pole in degrees (8); K, Fisher (1953) precision parameter; N, number of palaeomagnetic poles used to calculate the overall mean palaeomagnetic pole. *Palaeomagnetic pole obtained by averaging N¼9 palaeomagnetic poles listed in Muttoni et al. (2003, online table 2) and the palaeomagnetic pole from the Jebel Nehoud ring complex, Sudan (280+2 Ma) of Bachtadse et al. (2002). † Obtained from linear interpolation of bracketing Early Permian and Late Permian –Early Triassic palaeomagnetic poles. In this table, all West Gondwana palaeomagnetic poles are in NW Africa co-ordinates (Euler poles of rotation from Lottes & Rowley (1990)). In Fig. 7, the West Gondwana overall mean palaeomagnetic poles are plotted in Arabian co-ordinates using a NW Africa to Arabian Euler pole of rotation derived from Lottes & Rowley (1990) (26.228N, 11.158E, rotation angle 14.178).

DRIFT HISTORY OF IRAN

Massif des Maures pelites 161 St Affrique sediments 167 Lunner dykes 243+5 Ar/Ar 164 Late Permian–Early Triassic 162.9 Middle Triassic; based on the following palaeomagnetic poles: Mushelkalk, Silesia 123 Gipskeuper, Germany 131 Mushelkalk, Poland 143 Sherwood Sandstones, UK 139 Heming Limestone, France 141 Middle Triassic 135.3 Late Triassic; based on the following palaeomagnetic poles: Rhaetian sediments, Germany, France 112 Blue Anchor Formation, UK 109 Mercia Mudstone, UK 128 Twyning Formation, UK 114 Late Triassic 115.9

21

22

G. MUTTONI ET AL.

Fig. 7. (A) The Ordorician–Triassic APW paths of West Gondwana (Table 3, rotated into Arabian co-ordinates) and Europe (Table 3) are compared with palaeomagnetic poles from Late Ordovician site IR12, the Early Devonian Red beds of Wensink (1983) (RB) and the earliest Carboniferous Geirud lavas of Wensink et al. (1978) (GL). The congruence with West Gondwana poles lO (Late Ordovician) eD (Early Devonian) and lD–eC to eC (Late Devonian – Early Carboniferous), respectively, is evident. (B) The Permian–Triassic segments of the APW paths of West

DRIFT HISTORY OF IRAN

23

Fig. 8. Palaeolatitudes of the Iranian sites (acronyms and values in Table 2) compared with the palaeolatitudes expected at a nominal site located at 338N, 538E in central Iran from the West Gondwana APW path (Table 3, rotated into Arabian co-ordinates) and the Europe APW path (Table 3). Ages according to Gradstein et al. (2004).

The polarities of magnetization of the two entries from the Hambast and Elikah formations are known from magnetostratigraphic correlations with sections from the literature of believed known polarity (Besse et al. 1998; Gallet et al. 2000). Their north palaeomagnetic poles (Fig. 7B, poles H and HE) are significantly displaced with respect to both the West Gondwana and Europe APW paths. This displacement could be owing to variable combinations of plate motion and local vertical-axis tectonic rotations. In order to

discriminate between these two sources of displacement, the palaeomagnetic poles have been projected laterally along colatitude small circles, and have been found to intercept younger (Triassic) portions of the West Gondwana APW path (Fig. 7B). This suggests that the observed polar displacement has a main contribution from clockwise vertical-axis tectonic rotations (see also Besse et al. 1998) and a subordinate contribution probably from the plate motion of Iran relative to West Gondwana. Clockwise rotations explain the offset between Iran and

Fig. 7. (Continued) Gondwana and Europe are compared with palaeomagnetic poles from the late Middle Permian Ruteh lavas and the Late Permian Hambast Formation of Besse et al. (1998) (RL and H, respectively), and the Late Permian– early Early Triassic Hambast and Elikah formations of Gallet et al. (2000) (HE). Colatitude small circles of poles are indicated. (C) The Permian–Triassic segments of the APW paths of West Gondwana and Europe are compared with palaeomagnetic poles and associated colatitude small circles from late Early Triassic sites IR02 and IR04 from Nakhlak (NAK2 and NAK4, respectively) and the colatitude small circle of the Upper Triassic (lower Norian) Waliand limestone pole of Besse et al. (1998) (W). Age of poles as follows: e– mO, Early–Middle Ordovician; lO, Late Ordovician; mO–eS, Middle Ordovician–Early Silurian; m–lS, Middle– Late Silurian; lS– eD, Late Silurian–Early Devonian; eD, Early Devonian; mD, Middle Devonian; e– mD, Early– Middle Devonian; m– lD, Middle– Late Devonian; lD–eC, Late Devonian–Early Carboniferous; lD–mC, Late Devonian –Middle Carboniferous; eC, Early Carboniferous; lC1, early Late Carboniferous; lC2, late Late Carboniferous; lC–eP, Late Carboniferous–Early Permian; eP, Early Permian; lP– eTr, Late Permian–Early Triassic; mTr, Middle Triassic; lTr, Late Triassic; lTr–eJ, Late Triassic –Early Jurassic.

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West Gondwana poles along colatitude small circles, whereas relative plate motion explains the offset between the projected Iran poles and the coeval Late Permian –Early Triassic West Gondwana reference pole. According to this analysis, the sampled units were located at subequatorial palaeolatitudes (c. 28S + 28 for site H; c. 88S + 1.58 for site HE; Table 2) north of the Arabian margin of West Gondwana in Late Permian –early Early Triassic times (Fig. 8, sites H and HE), whereas the timing of postdepositional clockwise rotation is at present difficult to ascertain. Similar conclusions can also be drawn for the Upper Permian laterite of site IR27. Here, polarity is unknown. Hence, hemisphere of magnetization, and sense and amount of tectonic rotation, cannot be unequivocally determined. However, this constitutes a negligible problem in view of the fact that the characteristic magnetizations therein isolated clearly indicate equatorial palaeolatitudes of either approximately 18S or approximately 18N (+58) (Table 2, site IR27). This also suggests that the Alborz region, within which site IR27 is located, was already detached from the Arabian margin in Late Permian times (Fig. 8, site IR27). The notion derived from the present analysis that the Iranian block(s) had Arabian affinity in the Late Ordovician–earliest Carboniferous and were drifting to the north with respect to Arabia, attaining subequatorial palaeolatitudes in the Late Permian–early Early Triassic, is used to interpret the hemisphere of magnetization acquisition of the late Middle Permian Ruteh lavas, the polarity of which is unknown. The simplest scenario is to assign reverse polarity to its mean magnetization (Dec. ¼ 124.78E, Inc. ¼ 22.68; Table 2). The derived north palaeomagnetic pole is rotated counterclockwise with respect to the West Gondwana polar wander path, and its colatitude small circle intercepts slightly younger portions of the same path (Fig. 7B, pole RL). This implies similar conclusions as those for the Hambast and Elikah formations, i.e. that the observed polar displacement can be accounted for by a combination of vertical-axis tectonic rotations and plate motion. The Alborz region, within which the Ruteh lavas are located, resided at a palaeolatitude of approximately 128S + 28 (Table 2, site RL), slightly to the north of the Arabian margin in late Middle Permian times (Fig. 8, site RL).

Late Early Triassic – Late Triassic: Iran approaching Europe Data from Iran pertaining to this time window come from the characteristic components of Lower Triassic sites IR34 and IR35 from the Sorkh Shale of this study and Wensink (1982) (Fig. 1 and Table 2, site

SS), upper Lower Triassic (Olenekian) sites IR02 and IR04 of this study (site NAK), and the Upper Triassic (lower Norian) Waliaband limestone of Besse et al. (1998) (site W). According to the previous analysis, the Iranian block(s) moved northwards from the southern hemisphere to subequatorial palaeolatitudes during the Permian– early Early Triassic. Hence, the simplest scenario is to envisage a continuation of this northwards motion into the northern hemisphere during the Triassic. Polarity has been assigned to the mean characteristic components of sites IR02 and IR04 according to a northern hemisphere origin of magnetization, and was used to calculate north palaeomagnetic poles (Fig. 7C and Table 2; NAK2 and NAK4, respectively). Pole NAK2 falls very close to the Late Permian– Early Triassic reference pole of Europe. Pole NAK4 is rotated counterclockwise with respect to the APW path of Europe, and, when traced laterally along a colatitude small circle, it falls close to the Late Permian– Early Triassic reference pole of Europe, where pole NAK2 is also located (Fig. 7C). This suggests substantial European affinity of the sampled units and local vertical-axis tectonic rotations of the eastern part of the Nakhlak structure where site IR04 is located. A palaeolatitude of approximately 218N + 48 is derived by combining inclination-only data from both sites (Table 2, pole NAK). This palaeolatitude places the sampled units close to the European margin in late Early Triassic (Olenekian) times (Fig. 8, pole NAK). In a recent analysis, Bagheri & Stampfli (2008) interpreted the Naklakh–Anarak area as part of a ‘Variscan accretionary complex’ pertaining to the Turan domain. If this were the case, the European palaeolatitudes from the Naklakh structure would not be utilizable for the definition of the younger part of the Cimmerian drift history outlined in this study. Data from the Early Triassic Sorkh Shale are more difficult to interpret in terms of hemisphere of magnetization acquisition. Inclination-only data from combined sites IR34 and IR35 indicate a palaeolatitude of approximately 98 + 2.58 (Table 2, site SS1). Wensink (1982) reported a similar palaeolatitude (c. 118 + 88) for Sorkh Shale sites from the same general area of this study, and opted for a southern hemisphere origin of deposition (c. 118S). The substantial lack of constraints on the polarity of magnetization does not preclude, however, a northern hemisphere palaeolatitude, as suggested by Besse et al. (1998). The general drift trend of the Iranian block(s), as described above, is of moderate help in discerning between the two options. Palaeolatitudes from the Hambast and Elikah formations and site IR27 span from approximately 88S to equatorial in the Late Permian–early Early Triassic; late Early Triassic palaeolatitudes from Nakhlak are in the

DRIFT HISTORY OF IRAN

order of about 218N. Palaeolatitudes from the Sorkh Shale are tentatively interpreted as pertaining to the northern hemisphere, in agreement with Besse et al. (1998) (Fig. 8, site SS1), although caution on this interpretation should be maintained essentially

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because of the poor age constraints of the sampled units. Finally, Besse et al. (1998) showed the presence, in sites from the Upper Triassic (lower Norian) Waliaband limestone at Abadeh, of pre-folding

Fig. 9. Palaeomagnetic-based palaeogeographic reconstruction of West Gondwana and Pangaea over (A) the Late Ordovician, Early Devonian and Late Devonian– Early Carboniferous (B) Early Permian, (C) Middle Permian, (D) Late Permian–early Early Triassic and (E) Early– Middle Triassic, obtained using palaeomagnetic poles from West Gondwana and Europe (Table 3). Africa is internally treated as three plates (NW, NE and S Africa), with Adria (AD in panel B) tectonically coherent with NW Africa (Muttoni et al. 2003). Internal Gondwana plates are reconstructed using Euler poles of Lottes & Rowley (1990). Internal Laurasia plates are reconstructed using Euler poles of Bullard et al. (1965). Palaeolatitudes from Iranian sites discussed in the text and summarized in Table 2 are indicated by error bars and associated acronyms. Reconstructions made with PaleoMac 6.1 (Cogne´ 2003).

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palaeomagnetic directions variably rotated about vertical axes. Their analysis of inclination-only data produced an overall mean inclination value of 51.98 (a95 ¼ 4.68; Table 2). By assigning normal polarity to this mean inclination, according to a northern hemisphere origin of magnetization acquisition, a colatitude small circle containing the north palaeomagnetic pole has been calculated. This small circle intersects the Middle –Late Triassic portion of the European APW path (Fig. 7C, small circle W) suggesting European affinity for the Waliaband limestone, as originally proposed by Besse et al. (1998). These sites resided at a palaeolatitude of approximately 32.58N + 4.58, close to the European margin in early Norian times (Fig. 8, small circle W). In conclusion, data from this study and the literature suggest that the Iranian block(s) were located close to the Arabian margin of Gondwana in the Late Ordovician–earliest Carboniferous. They have drifted off this margin since the Permian, attained subequatorial palaeolatitudes at Late Permian–early Early Triassic times and approached the Eurasian margin by the late Early Triassic. From then on, the Iranian block(s) maintained European affinity, as deduced from early Norian data and, more generally, from Jurassic–Cretaceous data (e.g. Garedu Red beds, Bidou Beds and Dehuk Sandstones; Wensink 1982), as extensively discussed in Besse et al. (1998).

Palaeogeography The drift history of Iran, outlined in the previous section, has been placed in a general palaeogeographical context by reconstructing the configuration of continents using palaeomagnetic poles from West Gondwana and Europe (Table 3) and Euler poles from the literature (Fig. 9). According to the Ordovician–Carboniferous Gondwana APW path of McElhinny et al. (2003), the Arabian margin as part of Africa experienced significant plate motion, drifting from high southern latitudes in the Late Ordovician to subequatorial latitudes in the Early Devonian, and back to intermediate southern latitudes in the Late Devonian– Early Carboniferous (Fig. 9A). As anticipated in the previous section, Iran tracked closely this motion. The palaeolatitudes expected from the West Gondwana reference palaeopoles at Late Ordovician site IR12, Early Devonian site RB (Red beds) and earliest Carboniferous site GL (Geirud lavas) (Fig. 9A, stars) are in relatively close congruence with the palaeolatitudes calculated from the respective characteristic magnetizations (Fig. 9A, error bars). In the Carboniferous, Gondwana and Laurasia coalesced to form Pangaea. The Early Permian and Late Permian– Early Triassic

Pangaea reconstructions of Figure 9B, D, virtually identical to those of Muttoni et al. (2003), bracket a Middle Permian reconstruction (Fig. 9C) obtained by linear interpolation of Early Permian and Late Permian– Early Triassic poles of Table 3. At these times, palaeomagnetic data reveal a grand-scale tectonic scenario consisting of a transformation of Pangaea from an Irvingian B configuration (Irving 1977) in the Early Permian to a Wegenerian A-type configuration by the Late Permian–Early Triassic, as extensively discussed in Muttoni et al. (2003 and references therein). This transformation is coeval with the onset of northward drift of the Iranian block(s) across the Palaeotethys by way of Neotethys opening, depicting altogether an internally consistent plate motion scenario that is the subject of a parallel paper (Muttoni et al. submitted). Key elements of this northward drift are palaeolatitude estimates from the late Middle Permian Ruteh lavas (RL), the Late Permian – early Early Triassic Hambast and Elikah formations (H, HE), and the Late Permian site IR27 (Fig. 9C, D). Finally, palaeolatitudes from the late Early Triassic sites IR02 and IR04 from Nakhlak support an incipient approach of the Iranian block(s) with the active margin of Eurasia in the frame of an Early– Middle Triassic Pangaea configuration of the A-type (Fig. 9E, NAK).

Conclusions Palaeomagnetic data from this study and the literature have been used to reconstruct the drift history of the Iranian block(s) from the Ordovician to the Triassic. Key elements of this reconstruction are: † a robustly determined palaeomagnetic affinity of Iran with Arabia (Gondwana) during the Late Ordovician–earliest Carboniferous; † clear palaeomagnetic evidence for the residence of Iran on subequatorial palaeolatitudes during the Late Permian –early Early Triassic, noticeably disengaged from both the parental Gondwanan margin in the southern hemisphere and the espousal Laurasian margin in the northern hemisphere. On these two elements is based the interpretation of pre-Late Permian latitudes from Iran as pertaining to the southern hemisphere, and post-early Early Triassic latitudes as pertaining to the northern hemisphere. The resulting scenario is that of a relatively rapid northward drift that brought the Iranian block(s) to cover approximately 2500–3000 km of Palaeotethys in about 35 Ma (from c. 280 Ma in the Early Permian to c. 245 Ma at the end of the late Early Triassic), at an average plate speed of approximately 7 –8 cm year21. This motion was largely coeval to, and kinematically linked with,

DRIFT HISTORY OF IRAN

the transformation of Pangaea from an Irvingian B to a Wegenerian A-type configuration (Muttoni et al. 2003, submitted). The spatio-temporal resolution of the observed northward drift does not allow distinguishing components of relative motion between crustal blocks (e.g. Alborz v. Tabas v. Sanandaj–Sirjan). Despite these and other uncertainties (e.g. the possible Turanian rather than Cimmerian affinity of Naklakh: Bagheri & Stampfli 2008), the overall drift history outlined above is in broad agreement with geological data. Several authors pointed out the Proterozoic –Palaeozoic Gondwanan ancestry of north and central Iran based on the nature of their basement and the overlying sedimentary succession (Sto¨cklin 1968, 1974; Berberian & King 1981; Wendt et al. 2005). Zanchi et al. (2009b) pointed out that north and central Iran are entirely located south of the Eo-Cimmerian belt, which marks the collision of Iran with the southern Eurasian margin (e.g. Sengo¨r 1990; Zanchi et al. 2006). The upper age limit of this collision is Late Triassic, as the Upper Triassic– Middle Jurassic Shemshak Group seals unconformably EoCimmerian structures affecting the Palaeozoic– Middle Triassic sequences of northern Alborz (Zanchi et al. 2006). In this paper we opted to avoid discussing palaeomagnetic directions in terms of tectonic rotations within the assorted Iranian fold-and-thrust belts, a subject exhaustively treated by Besse et al. (1998). Nevertheless, our data from Nakhlak advise caution with regard to the hypothesis by Soffel et al. (1996) that a 1358 counterclockwise block rotation has occurred between the internal part of central Iran and Europe since the Middle Triassic, an hypothesis that is based on sparse data from the Triassic sequence of Nakhlak. Multiple remagnetization events pervasively overprinted the volcanoclastic sandstones and siltstones at Nakhlak, whereas magnetization components of presumed primary age were observed only in a solitary nodular limestone interval. These primary components, when compared with reference European directions, suggest no rotation of the western part of the Nakhlak structure (site IR02) and a moderate counterclockwise rotation of the eastern part of the Nakhlak structure since the late Early Triassic (site IR04). A complex and largely poorly known pattern of tectonic rotations around vertical axes has been active in central Iran since the Neogene (Walker & Jackson 2004) as a consequence of crustal-scale north–south shearing. Undoubtedly, more data are required to unravel in detail the complex tectonic history of Iran, a country of over 1.6  106 km2 and less than 30 palaeomagnetic entries listed in the IAGA Global Palaeomagnetic Database 2007.

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The MEBE (Middle East Basin Evolution) Programme funded the field work. Helpful comments by reviewers J. Tait and V. Bachtadse improved the manuscript. We thank M.-F. Brunet and J. Granath for assistance during preparation of this manuscript.

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DRIFT HISTORY OF IRAN implications for Oman. In: R OBERTSON , A. H., S EARLE , M. P. & R IES , A. C. (eds) The Geology and Tectonics of the Oman Region. Geological Society, London, Special Publications, 49, 797–831. S HARKOVSKI , M., S USOV , M. & K RIVYAKIN , B. 1984. Geology of the Anarak Area (Central Iran). Explanatory Text of the Anarak Quadrangle Map. Geological Survey of Iran, 19. S OFFEL , H. C., D AVOUDZADEH , M., R OLF , C. & S CHMIDT , S. 1996. New palaeomagnetic data from Central Iran and a Triassic palaeoreconstruction. Geologische Rundschau, 85, 293–302. S TO¨ CKLIN , J. 1968. Structural history and tectonics of Iran: a review. AAPG Bulletin, 52, 1229–1258. S TO¨ CKLIN , J. 1974. Possible ancient continental margins in Iran. In: B URK , C. A. & D RAKE , C. L. (eds) The Geology of Continental Margins. Springer, Berlin, 873–887. S YMONS , D. T. A., S ANGSTER , D. F. & L EACH , D. L. 1995. A Tertiary age from paleomagnetism for Mississippi Valley-type zinc–lead mineralization in Upper Silesia, Poland. Economic Geology, 90, 782–794. 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 HE´ VENIAUT , H., B ESSE , J., E DEL , J. B., W ESTPHAL , M. & D URINGER , P. 1992. A Middle Triassic Paleomagnetic Pole for the Eurasian Plate from Heming (France). Geophysical Research Letters, 19, 777– 780. T ORSVIK , T. H., E IDE , E. A., M EERT , J. G., S METHURST , M. A. & W ALDERHAUG , H. J. 1998. The Oslo rift: new palaeomagnetic and 40Ar/39Ar age constraints. Geophysical Journal International, 135, 1045–1059. V AN DER V OO , R. 1993. Paleomagnetism of the Atlantic, Tethys and Iapetus Oceans. Cambridge University Press, Cambridge. W ALKER , R. & J ACKSON , J. 2004. Active tectonics and late Cenozoic strain distribution in central and

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eastern Iran. Tectonics, 23, TC5010, doi:10.1029/ 2003TC001529. W ENDT , J., K AUFMANN , B., B ELKA , Z., F ARSAN , N. & B AVANDPUR , A. K. 2005. Devonian/Lower Carboniferous stratigraphy, facies patterns and palaeogeography of Iran Part II. Northern and central Iran. Acta Geologica Polonica, 55, 31– 97. W ENSINK , H. 1979. The implications of some palaeomagnetic data from Iran for its structural history. Geologie en Mijnbouw, 58, 175. W ENSINK , H. 1982. Tectonic inferences of paleomagnetic data from some Mesozoic formations in Central Iran. Journal of Geophysics, 51, 12– 23. W ENSINK , H. 1983. Paleomagnetism of red beds of Early Devonian age from Central Iran. Earth and Planetary Science Letters, 63, 325–334. W ENSINK , H., Z IJDERVELD , J. D. A. & V AREKAMP , J. C. 1978. Paleomagnetism and ore mineralogy of some basalts of the Geirud formation of Late Devonian to Early Carboniferous age from southern Alborz, Iran. Earth and Planetary Science Letters, 41, 441– 450. Z ANCHI , A., B ERRA , F., M ATTEI , M., G HASSEMI , M. & S ABOURI , J. 2006. Inversion tectonics in Central Alborz, Iran. Journal of Structural Geology, 28, 2023– 2037. Z ANCHI , A. ET AL . 2009a. The Cimmerian evolution of the Nakhlak–Anarak area, Central Iran, and its bearing for the reconstruction of the history of the Eurasian margin. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 261– 286. Z ANCHI , A. ET AL . 2009b. The Eo-Cimmerian (Late? Triassic) orogeny in North Iran. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 31– 55. Z IJDERVELD , J. D. A. 1967. A.C. demagnetization of rocks – analysis of results. In: C OLLINSON , D. W., C REER , K. M. & R UNCORN , S. K. (eds) Methods in Paleomagnetism. Elsevier, New York, 254– 286.

The Eo-Cimmerian (Late? Triassic) orogeny in North Iran ANDREA ZANCHI1, STEFANO ZANCHETTA1, FABRIZIO BERRA2, MASSIMO MATTEI3, EDUARDO GARZANTI1, STEWART MOLYNEUX4, AMIR NAWAB5 & JAFAR SABOURI5 1

Dipartimento Scienze Geologiche e Geotecnologie, Universita` di Milano-Bicocca, Piazza della Scienza 4, Milano, 20126 Italy (e-mail: [email protected]) 2

Dipartimento di Scienze della Terra, Universita` di Milano, Via Mangiagalli 34, 20133 Milano, Italy

3

Dipartimento di Scienze Geologiche, Largo San Leonardo Murialdo 1, 00146 Roma-I, Universita` Roma TRE, Italy

4

British Geological Survey, Kingdom Kingsley Dunham Centre Keyworth, Nottingham NG12 5GG, UK 5

Geological Survey of Iran, Azadi Square, Meraj Avenue, 13185-1494, Tehran, Iran Abstract: The Eo-Cimmerian orogen results from the Late Triassic collision of Iran, a microplate of Gondwanan affinity, with the southern margin of Eurasia. The orogen is discontinuously exposed along the northern side of the Alborz Mountains of North Iran below the siliciclastic deposits of the Shemshak Group (Late Triassic– Jurassic). A preserved section of the external part of the belt crops out in the Neka Valley (eastern Alborz) south of Gorgan. Here the Mesozoic successions (Shemshak Group–Upper Cretaceous limestones) overlay a pre-Jurassic Eo-Cimmerian thrust stack with a sharp unconformity. The stack includes the Gorgan Schists, an Upper Ordovician– Lower Silurian low-grade metamorphic complex, overthrusted southward above a strongly deformed Late Palaeozoic–Middle Triassic succession belonging to north Iran. In the Talesh Mountains (western Alborz), the Shanderman Complex, previously interpreted as an ophiolitic remnant isolated along the Eo-Cimmerian suture, is considered an allochthonous nappe of deeply subducted continental crust. The new evidence for this is the occurrence of previously unknown eclogites dating to the Carboniferous, and probably related to the Variscan history of Transcaucasia. South of the Shanderman Complex, Upper Palaeozoic slates and carbonates occurring below the Lower Jurassic Shemshak Group also record the occurrence of an Eo-Cimmerian metamorphic event. Based on our new data, the Eo-Cimmerian structures exposed in the Alborz appear to be remnants of a collisional orogen consisting mainly of deformed continental crust where no ophiolites are preserved.

The Late Triassic–Early Jurassic Cimmerian orogeny, which produced the final closure of the Palaeotethys Ocean and the accretion of Gondwana-derived microplates to southern Eurasia, was first recognized in Iran (Sto¨cklin 1974). This event, which reshaped the eastern Eurasian margin from Timor to Turkey, is still unclear concerning its location, extent, timing and evolution. Since Sto¨cklin (1974), the continental blocks forming north and central Iran (Fig. 1) have been considered to be of Gondwanan affinity for two major reasons: † their pre-Palaeozoic basement is thought to be related to the ‘Pan-African’ orogeny; † Precambrian–Cambrian successions are comparable across the Zagros suture. In addition, the

Lower Cambrian Laloon Formation has several equivalent units in the Salt Range of Pakistan, in Arabia (Sag Sandstone), Jordan (Lower Quweira Sandstone) and in Turkey (Cardac YaylaCalaktepe Formation). The region lacks a late Palaeozoic (Variscan) deformation (Saidi et al. 1997), whereas its northern portion, now occupied by the Alborz belt, was involved in the Cimmerian orogeny (Sengo¨r 1990). According to most authors, Iran began to drift northward in Early Permian times as part of the Cimmerian blocks (Sengo¨r 1979, 1984, 1990; Angiolini et al. 2007) or Mega-Lhasa (Dercourt et al. 2000) following the opening of Neotethys. Collision with Eurasia sensu latu occurred during

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 31– 55. DOI: 10.1144/SP312.3 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Tectonic scheme of Iran with the main tectonic subdivisions. The location of the Eo-Cimmerian units is shown with a vertical dash. Black: Mesozoic ophiolites along the Main Zagros and other sutures. AMC, Anarak Metamorphic Complex. Modified from Angiolini et al. (2007).

the Late Triassic, giving origin to the EoCimmerian deformation in North Iran, followed by a strong but poorly known Neo-Cimmerian compressional event during Late Jurassic times mainly affecting central Iran. Available palaeomagnetic data (Wensink et al. 1978; Besse et al. 1998; Muttoni et al. 2009) also suggest that at least some portions of central Iran (e.g. Abadeh) and the Alborz area moved in conjunction from southern latitudes in the Early Permian to northern latitudes in the Early –Middle Triassic, crossing subequatorial latitudes at Middle –Late Permian times. A major northwards subduction zone developed during Permo-Triassic times below the Eurasian margin, which is now located at the latitude of the southern Caspian coast (Berberian & King 1981; Alavi 1991; Boulin 1991; Ruttner 1993; Alavi et al. 1997). Remnants of one major Palaeotethys suture between Eurasian sensu latu and the Iranian block were first recognized in NE Iran around Mashad and eastward, where thin slices of

ultramafic rocks associated with pillow lavas, phyllites and siliciclastic turbidites were interpreted as an accretionary wedge formed during PermoTriassic times (Alavi 1991; Ruttner 1993). The flat-lying Middle Jurassic Kashaf Rud Formation, post-dating the final accretion of Iran to the southern margin of Eurasia, definitively seals the collision zone, although undated conglomerates and sandstones (Geological Survey of Iran 1986, 1996; Alavi 1991) exposed below the Kashaf Rud Formation around Mashad might constrain the time of deformation to the end of the Triassic or the beginning of the Jurassic. The continuation of the suture zone to the west, proposed by Alavi (1996), across central and western Alborz (Talesh Mountains) is questionable. Docking of the Iranian block to Eurasia occurred in the Late Triassic times. The collision is marked by an angular unconformity between the Upper Precambrian –Middle Triassic successions of north Iran and the Norian–Lower Jurassic Shemshak

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Fig. 2. Structural map of the Alborz belt based on Geological Survey of Iran (1989) and on our observations. The two squares refer to the location of Figures 3 and 8.

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Formation (Seyed-Emami 2003; Fu¨rsich et al. 2005), a thick terrigenous clastic unit generally considered to be post-orogenic clastics, which has been now defined as the Shemshak Group (Fu¨rsich et al. 2009). In this paper, we present a first systematic description and reconstruction of the tectonic setting of the Alborz region during the Eo-Cimmerian event, based on new data collected in the Alborz belt in the region south of Gorgan and in the Talesh Mountains (Fig. 2).

Geological setting of the Alborz The Alborz belt is a 1500 km-long mountain system extending from Azerbaijan to Afghanistan, flanking in its central part the southern coast of the Caspian Sea (Fig. 2). The belt was affected by several successive tectonic events, from the Eo-Cimmerian orogeny to the Late Tertiary– Quaternary intracontinental transpression (Allen et al. 2003a), and is still strongly seismically active (Ritz et al. 2006). Important compressional events also occurred at the end of the Cretaceous (Guest et al. 2006a) in connection with retro-arc compression induced by the subduction of the Neotethys below the Sanandaj–Sirjan zone and by the subduction of the Nain –Baft ocean directly below Central Iran (Ghasemi & Talbot 2006). Extensive mapping by the Geological Survey of Iran (Geological Survey of Iran 1977, 1985, 1987, 1988, 1991a –c, 1997, 2001, 2004) and other recent studies now provide an accurate stratigraphical and structural general framework of the belt (Alavi 1996; Allen et al. 2003a, b; Guest et al. 2006a, b; Zanchi et al. 2006).

Stratigraphic framework The stratigraphic successions preserved in the Alborz (Assereto 1966b; Alavi 1991) are overall more than 12 km thick, spanning from the latest Precambrian to the Holocene. The uppermost Precambrian–Middle Triassic sedimentary successions were deposited along a passive margin affected by the opening of the Palaeotethys Ocean during Early Palaeozoic times and, successively, by the opening of the Neotethys Ocean in the Late Palaeozoic (Stampfli et al. 1991). No Precambrian crystalline basement crops out in the belt, with the exception of the Lahijan Granite east of Rasht, which has given a late Neoproterozoic–Cambrian U– Pb zircon crystallization age (Guest et al. 2006b). The terrigenous Shemshak Group overlies with angular unconformity the aforementioned successions deformed during the Middle–Late Triassic Eo-Cimmerian orogenesis.

About 3000 m of Precambrian and Cambrian shallow-marine sandstones and dolostones, with Lower Cambrian continental deposits, form the base of the succession. Lower– Middle Ordovician glauconitic shales and siltstones follow, and are covered by Silurian lava flows around Ramsar in the central Alborz. A thick succession of volcanic deposits including lava flows, intercalated with Orthoceras-bearing limestones, occurs in the Talesh Mountains (western Alborz) and east of Gorgan, where basaltic –andesitic lava flows are associated with ignimbrites and intrusive rocks including syenite, diorite and gabbro (Davies et al. 1972; Geological Survey of Iran 1991b, 1997). Thick lava flows are also widespread at the base of the Devonian units. Wendt et al. (2005), observing the strong difference of the Palaeozoic successions of the Talesh Mountains from the ones of central Alborz, suggested the possibility that they represent the distal portion of the Iranian Plate or, alternatively, that they may belong to the Turan Plate. The entire volcanic succession was interpreted as representative of an important rifting event (Berberian & King 1981; Stampfli et al. 1991). The Devonian–Middle Triassic succession is 1300–1500 m thick and includes shallow-marine ramp carbonates, evolving to carbonate platforms in the Triassic (Elika Formation). Intercalated siliciclastic units (Qezel Qaleh Formation) and unconformities occur within the Lower Carboniferous carbonates (Mobarak Formation) at the base of the Permian units (Dorud Formation) and, in eastern Alborz, between the Ruteh and Elika Formations close to the Permo-Triassic boundary. They mark several uplift episodes, possibly related to the effects of faraway tectonic events (Gaetani et al. 2009). The whole pre-Middle Triassic succession is sealed with angular unconformity by the Upper Triassic –Middle Jurassic sediments of the Shemshak Group, locally reaching 4000 m in thickness. Its base is markedly heterochronous, ranging from Carnian in central Alborz to Toarcian in the Talesh Mountains around Masuleh in the internal part of the Eo-Cimmerian belt (Clark et al. 1975; Seyed-Emami 2003; Ghasemi-Nejad et al. 2004; Fu¨rsich et al. 2005). This unit includes deltaic– shallow-marine sandstone, conglomerate, shale and coal layers that seal the Eo-Cimmerian orogen and its foreland. The unconformity is subtle, commonly passing to a paraconformity in the northern and axial part of the central Alborz, whereas in the Shemshak area the basal beds of the Shemshak Group lie on Permian–Middle Triassic units. In the Talesh and Gorgan regions, west and east of the study region, the Shemshak Group non-conformably seals the Eo-Cimmerian nappes. Thick basaltic lava flows commonly occur above the unconformity, especially in the southern Alborz.

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Upper Jurassic – Cretaceous shallow-water limestones and marls, intercalated with uppermost Jurassic basaltic flows, are locally preserved. In the Gorgan region, Upper Cretaceous limestones definitively rest on Triassic and earlier rocks, sealing the Eo-Cimmerian relief with spectacular angular unconformity (Berra et al. 2007). The Cretaceous limestones commonly show closed folds and thrust faults sealed by the Palaeocene Fajan conglomerates. The Eocene volcanic and volcanoclastic rocks of the Karaj Formation, more than 3000 m thick, succeed the conglomerates. This unit, which is much thicker south of the Talesh Mountains, is related to the growth of an intracontinental volcanic arc in an extensional setting related to the last phases of the Neotethys subduction along the Zagros suture zone (Alavi 1996). The Neogene successions show different facies and thickness, including continental deposits in the Talesh area and shallow-sea coastal finegrained terrigenous bodies with gypsum layers and bioclastic limestones in southern central Alborz. These basins possibly formed in a foreland basin developed at the front of the growing Alborz belt. Quaternary alluvial deposits form thick accumulations in front of the belt, where they are strongly deformed and uplifted by active faults (Astaneh– Firuzkuh–Mosha and N Teheran–Qazvin faults to the south, and Khazar and South Talesh faults to the north).

Tectonic setting The Alborz belt can be divided in two main portions that show a different structural setting: western Alborz, forming the Talesh Mountains; and centraleastern Alborz east of Rasht. The separation between the two parts of the belt is at the longitude of Rasht, where lateral ramps and transverse faults occur along the Sefid Rud. The Talesh Mountains are narrower than 50 km, whereas central and eastern Alborz reach 100 km. In the analysed part of the Talesh Mountains, a set of parallel SW-verging Late Neogene thrust faults form huge ramp anticlines uplifting the Palaeozoic succession in their hanging walls. N – S-striking dextral strike-slip faults form the western lateral ramps of the main thrusts. A poorly exposed N –S fault associated with earthquakes related to thrusting along a horizontal fault plane (Jackson et al. 2002) marks the Eastern margin of the belt. The Talesh Mountains include a few isolated strips of metamorphic rocks, the Gasht and Shanderman complexes, exposed along the northern slopes of the belt along the South Caspian coast. Recent thrusts and strike-slip faults, which obscure their original relationships with the surrounding PalaeoMesozoic successions, define the boundaries of

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the two units. The Shanderman Complex, cropping out just west of Rasht along the Caspian foothills of the Talesh Mountains, is described as a deformed association of slate, phyllite, gneiss and amphibolite, with small patches of serpentinized peridotite affected by a medium-grade metamorphism (Davies et al. 1972; Clark et al. 1975). Based on presumed lithological similarities, Alavi (1991, 1996) suggested that the Shanderman Complex is an ophiolitic unit possibly equivalent to the ones exposed near Mashad, tracing the Palaeotethys suture along the South Caspian coast from there to the Talesh region. The Gasht Complex consists of two units with a different metamorphic imprint. The lower succession mainly includes medium- to high-grade metapelites, quartzites and amphibolites intruded by granitoids later affected by a retrograde metamorphism. Whole rock Rb–Sr dating obtained on phyllites (Crawford 1977) gave a problematic Middle–Late Devonian age of 382 + 48 Ma and 375 + 12 Ma. The upper unit includes slate, phyllite and quartzite. Clark et al. (1975) also described very low-grade metapelites with Upper Palaeozoic limestone intercalations around Masuleh and along the Shah Rud. The Shemshak Group non-conformably covers the metamorphic units of the Talesh area. The lowermost marine beds of the formation date to the lower Lias, but its base could be older (Davies et al. 1972). Granitoids intruding these metamorphic units show a Jurassic age (175 Ma; Crawford 1977). Central-eastern Alborz is dominated from west to east by WNW– ESE-, E– W- and ENE – WSW-trending high-angle and vertical faults running parallel to the belt (Fig. 2). Right- and left-lateral strike-slip motions and oblique to dip-slip movements (Allen et al. 2003a, b; Guest et al. 2006a; Zanchi et al. 2006) occur in the internal part of the belt, whereas reverse and thrust faults characterize the external fronts. The uppermost Precambrian–Lower Mesozoic units form the backbone of the present-day mountain chain between the Kojour Fault to the north and the Kandevan Thrust to the south in central Alborz. The northern part of the central Alborz belt consists of poorly deformed thick Mesozoic marine successions, delimited to the north by the active Khazar Thrust running parallel to the Caspian Sea. In the Gorgan region a low-grade unit, the Gorgan Schists, forms the northernmost part of the belt. Although Alavi (1996) assigned a possible Palaeozoic–Triassic age to the Gorgan Schists, these rocks have been recently attributed to the Early Palaeozoic (Late Ordovician– Silurian?; Geological Survey of Iran 1997, palynology by J. Sabouri). The main structures of the internal part of the eastern Alborz are the poorly studied North Alborz

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and Badeleh faults, both trending ENE–WSW. The North Alborz Fault has been interpreted in different ways (Geological Survey of Iran 1991b, 1997; Allen et al. 2003a). According to recent maps, it is a high-angle north-dipping fault, which marks the boundary between the Palaeozoic–Middle Triassic successions to the north, deformed during the

Eo-Cimmerian event, and its undeformed foreland to the south (Geological Survey of Iran 1997). The southern part of the belt shows a complex tectonic setting characterized by pronounced shortening because of a complex interaction among strike-slip and reverse faults producing the uplift of the lower part of the Palaeozoic succession (Allen et al.

Fig. 3. (a) Simplified geological map of the Neka Valley from Geological Survey of Iran (1997) modified based on our data. Location within the Alborz belt, in Figure 2. (b) Stereographic projections concerning the Eo-Cimmerian structures exposed along the Neka Valley. Site Gorgan1 Neka1 includes data on the Gorgan Schists measured south of the town of Gorgan (368420 5900 ; 548) and in the Neka Valley (368360 1400 ; 538490 5100 ) just west of Figure 3, close to the bridge of the village of Sefecha. Small empty circle, pole to bedding S0; pole to axial plane slaty cleavage of the D1 phase; black small circles, fold axes of phase D2; grey small circles, pole to axial surface of phase D2. In the other plots thin lines are faults, black dots represent striations with relative sense of motion and square are poles to mylonitic foliation along the Radekan Fault. The other symbols refer to the same geometrical features but are all related to the D1 deformational event. Schmidt’s projection, lower hemisphere.

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Fig. 3. Continued.

2003a; Guest et al. 2006a). The exposure of the Palaeozoic units in the Shemshak region of the central Alborz is believed to represent the effect of a mainly Cenozoic inversion of Middle–Late Triassic extensional structures formed in the foreland Triassic basin (Zanchi et al. 2006). Important dextral strike-slip ENE – WSWtrending faults have been recognized by several authors (Axen et al. 2001; Allen et al. 2003a; Zanchi et al. 2006) and related to N– S compression during the Miocene. A recent change in the sense of shearing from dextral to sinistral along the main faults parallel to the belt has been related to the major plate reorganization following the Eurasia – Arabia collision (Jackson et al. 2002), inducing a northwestward motion of the South Caspian since the Middle Pleistocene (Ritz et al. 2006). Oblique convergence along the central-eastern Alborz is now absorbed by major active transtensional faults as the Astaneh–Firuzkuh –Mosha fault system (Ritz et al. 2006), showing a left-lateral offset of about 30 km west of Teheran (Allen et al. 2003a).

The Eo-Cimmerian structures in the Alborz The Gorgan region Fieldwork was carried out around the town of Gorgan and along the Neka Valley, where the Palaeo-Mesozoic successions of the Alborz are well

exposed (Figs 3 and 4). Three main different tectonostratigraphic units have been analysed between the South Caspian coast and the North Alborz Fault: (1) the Gorgan Schists forming the strongly forested high mountain ridge (2400 m) running parallel to the Caspian coast; (2) the Palaeozoic–Triassic sedimentary succession of North Iran, which has been deformed during the Eo-Cimmerian orogeny; and (3) an ?Upper Triassic–Cretaceous sedimentary succession unconformably covering the two previous units (Fig. 5). The Gorgan Schists. The Gorgan Schists form the high topographic reliefs that border the SE Caspian coast from the town of Neka to Aliabad. The active Gorgan (Khazar) thrust fault, which dips southward, bounds the Gorgan Schist to the north. The metamorphic complex is stacked southward on the Palaeozoic successions of North Iran along the Radekhan Fault, an Eo-Cimmerian highangle north-dipping fault, crossing the northern flank of the Neka Valley (Figs 3 and 4). The Gorgan Schists mainly consist of slate and phyllite, including a thick volcanic–volcanoclastic succession, intruded by small basic and acidic bodies exposed south of Gorgan. Thin marble layers with fossil ghosts commonly occur in the phyllites. The Shemshak Group non-conformably covers the unit. A subhorizontal foliation, S1, defined by very finegrained chlorite and sericite, forming a pervasive slaty cleavage, is present in the metapelites and is gently folded by a second deformational event (D2),

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Fig. 4. General composite sections across the Neka Valley; traces A– B and C–D are in Figure 3a. The Eo-Cimmerian structures formed along the frontal part of the belt are preserved below the Mesozoic successions. Neogene reverse faults are also shown. GA, Ghalimoran Limestone (Late Cretaceous). See text for details. Geometry of the folds in the Gorgan Schists are partially inferred. Scales are different in the two sections.

accompanied by the formation of an S2 spaced fracture cleavage (Figs 3b and 5g) related to kink bands and small-scale chevron folds. Where the metamorphic grade increases, the S1 foliation is marked by the preferred orientation of sericitic white mica, quartz and albitic plagioclase. A penetrative foliation is also present in the meta-volcanic rocks. The S1 foliation is generally subhorizontal, showing open ENE–WSW-trending D2 folds with a kilometric wavelength. Structural analyses performed along the Radekhan Fault (Fig. 3b) indicate a strong deformation of the Gorgan Schists in the hanging wall accompanied by S–C shear structures. Tight– isoclinal folds occur in the lower part of the Carboniferous Mobarak Formation along the footwall of the fault (Figs 3 and 4). Available K –Ar whole-rock radiometric data of samples from the Gorgan Schists (Delaloye et al. 1981) cluster between 250 and 200 Ma, confirming an ‘Eo-Cimmerian’ age for the metamorphism and deformation of the complex. The Gorgan Schists have yielded poorly preserved acritarchs, chitinozoa, scolecodonts and cryptospores (Fig. 6), the first three groups indicating deposition in a marine environment. Acritarchs from Kond –Ab towards the bottom of the valley, Mile Radkan from the Gorgan Schists in the Neka Valley, include Veryhachium sp., with representatives of the V. lairdii and V. trisipinosum groups, ?Actinotodissus crassus, Diexallophasis cf. denticulata, and a variety of acanthomorphs (samples 79MH164 and 2001, js, ksg, 1– 11). Also present

are chitinozoa, including ?Belonechitina sp., Desmochitina sp. (sample 2001, js, ksg, 9) and cryptospore tetrads (samples 79MH164, 2001, js, ksg, 3). Similar assemblages were recorded from a further two samples. One, from a locality close to the Radekan Tower (GO5: 368370 4500 ; 548060 1800 ), yielded the acritarchs ?Tylotopalla caelamenicutis and ?Cymatiosphaera or ?Dictyotidium, together with sphaeromorph and acanthomorph acritarchs and possible cryptospore dyads and tetrads. The second sample, from the Ziarat Valley, south of Gorgan (GO44: 368490 1900 , 548290 0500 ), yielded ?Veryhachium trispinosum and acanthomorph acritarchs, specimens of ?Belonechitina, scolecodonts and possible cryptospore dyads. Tylotopalla caelamenicutis ranges from the latest Ordovician (post-glacial Hirnantian Stage: Vecoli & Le He´risse´ 2004, fig. 6) to the mid Silurian; it has not been recorded above the Llandovery in the Middle East (Jordan, Saudi Arabia: Keegan et al. 1990: Le He´risse´ et al. 1995), but has been recorded from the Wenlock at lower palaeolatitudes (Welsh Borderland, Gotland: Dorning 1981; Le He´risse´ 1989). Its occurrence suggests a latest Ordovician –early or mid-Silurian age. The specimens of ?Neoveryhachium and Diexallophasis also support this age determination. Both genera appear in the Late Ordovician (Caradoc and Ashgill, respectively: Vecoli & Le He´risse´ 2004, fig. 6), but ?A. crassus is restricted to the Late Ordovician whereas Diexallophasis ranges into the Devonian.

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39

Fig. 5. Field views of the Eo-Cimmerian structures and of the Upper Cretaceous unconformity in the upper Neka Valley. Compare with the Tertiary structures depicted in Figure 6 for the different style of folding. (a) Eo-Cimmerian structures below the Ghalimoran Formation in the upper Neka Valley just north of the Tertiary Talanbar Fault; (b) folds and faults in the Elika Formation along the uppermost part of the Neka Valley. Folds are entirely sutured by the Upper Cretaceous limestones; (c) and (d) the Upper Cretaceous unconformity sealing the Eo-Cimmerian structures along the middle Neka Valley. Structural data concerning the Eo-Cimmerian folds developed in the Lower Carboniferous Mobarak Formation are reported in Figure 3 (NEKA3 and NEKA6); (e) the Ghalimoran Limestone lays on a folded succession (NEKA9B in Figure 3) of Palaeozoic Red Nodular Limestones of unknown age. A Tertiary (Neogene?) steep reverse fault is responsible for the uplift of the Permian Ruteh limestones along the hanging wall above the Ghalimoran Formation. Faults in plot NEKA9B (Figure 3) cross-cut the Eo-Cimmerian folds formed in the red limestones. Note the flat geometry of the unconformity surface; (f) a close view of the Upper Cretaceous unconformity. Note the flat geometry of the surface; and (g) structural features of the Gorgan Schists in the Neka valley; two deformational events are shown.

40

A. ZANCHI ET AL.

Fig. 6. Acritarchs (1– 6, 8 –10, 13), chitinozoans (11–12, 15– 16), cryptospores (14, 17) and a scolecodont (7) from the Gorgan Schists. (1, 2) Tylotopalla caelamenicutis? Loeblich 1970. MPK 13545, sample GO5, England Finder co-ordinates M9/3. (3, 8) Acanthomorph acritarchs. (3) MPK 13546, GO44, L29/0; (8) MPK 13547, GO5, Q18/4. (4, 5) Veryhachium trispinosum (Eisenack) Stockmans & Willie`re 1962. (4) 79MH164 (33). (5) MPK 13548, GO44, G19/0. (6) Diexallophasis sp. cf. D. denticulata (Stockmans & Willie`re) Loeblich 1970. 79MH164 (2). (7) Scolecodont. MPK 13552, GO44, G27/0. (9, 10) Cymatiosphaera? or Dictyotidium? MPK 13549, GO5, M61/0. Specimen with two orders of reticulate ornament resembling that of the Silurian species Dictyotidium biscutulatum Kiryanov. (11) Belonechitina? sp. MPK 13550, GO44, Q35/3. (12, 15) Belonechitina? sp. MPK 13551, GO44, T30/1. (13) ?Actinotodissus crassus Loeblich & Tappan 1978. 2001 js, ksg, sa (5). (14, 17) Cryptospore tetrads? (14) MPK 13553, GO5, R12/2. (17) 79MH164 (21). (16) Desmochitina sp. 2001, js, ksg, 9b. Bar (Fig. 3) ¼ 5mm for 6.15, 20mm for 6.16, and 10mm for the rest. Specimens labelled MPK are held in the type and figured specimens collection of the British Geological Survey, Nottingham, UK; the remaining specimens are held in the collections of the Palynological Laboratory, Geological Survey of Iran.

THE EO-CIMMERIAN OROGENY IN NORTH IRAN

The specimens of ?Belonechitina from GO44 are relatively small for chitinozoa in general, and for specimens of Belonechitina in particular, but they are similar in some respects, notably in their size and the nature and distribution of their ornament, to Belonechitina paravitrea, a species recorded from the Llandovery (Rhuddanian– Aeronian) of Saudi Arabia (Paris & Al-Hajri 1995; Paris et al. 1995). The genus Belonechitina ranges from the early Ordovician (Arenig Series) to the Upper Silurian (Ludlow Series). The palynological evidence thus suggests a possible Early Silurian depositional age for the Gorgan Schists. The Palaeozoic–Triassic succession of the Neka Valley. An intensively folded and faulted succession ranging in age from Devonian to Middle Triassic is exposed along the Neka Valley between the Radekhan and the North Alborz faults (Fig. 3a). Devonian rocks of the Kosh Yeilagh Formation include well-bedded limestone, shale, sandstone, dolostone and diabase, and are covered by the Missisipian Mobarak Formation with marly limestone at the base passing to well-bedded dark limestone at the top. The Quezel Qaleh Formation (Pennsylvanian) follows with carbonate and terrigenous facies. The Permian successions include, at the base, the Dorud Formation with red conglomerate, sandstone, silty limestone and limestone with fusulinids. Bedded–massive limestone of the Ruteh Formation and the dolomitic limestone and dolostone of the Elika Formation follow up-section. High-angle faults and intensive folding with steep axial planes intensely deform the whole Palaeozoic– Early/Middle Triassic succession exposed between the Gorgan Schists and the North Alborz Fault (Figs 3b, 4 and 5). The occurrence of a regional Mesozoic unconformity, which cuts all these structures, demonstrates that they formed during the Eo-Cimmerian orogeny. Flat-lying gently deformed Mesozoic deposits (Shemshak Group, Lar Limestone and Ghalimoran Formation) unconformably cover previous faults and folds affecting the Gorgan Schists and the Palaeozoic– Middle Triassic deformed succession. This unconformity is well exposed along the middle –upper Neka Valley, where we focused our attention on the description of the Cimmerian deformation. Folding style strongly depends on competence contrasts and bedding thickness of the folded multilayer, which included very different lithologies. Short decametric–heptometric tight–isoclinal folds with very steep axial planes developed within the well-bedded limestone of the Mobarak Formation. (Fig. 5c, d, f). Open ramp anticlines and chevron folds with a heptometric wavelength and amplitude formed in the carbonates of the Elika

41

Formation due to parallel folding (Fig. 5b). Overturned folds in the Ruteh Limestone are faulted against the Devonian units north of the Talambar Fault in the upper part of the Neka Valley (Fig. 5a). Parallel folding and a poorly developed axial plane disjunctive cleavage suggest that deformation occurred at shallow crustal levels. The Santonian Ghalimoran Formation, which only shows gentle folds (Fig. 7), unconformably covers everywhere these structures. South of the North Alborz Fault, where stratigraphic relationships are preserved and the effects of the Eo-Cimmerian event are less intense, the Shemshak Group forms a low-angle unconformity with the underlying successions. The Late Triassic –Early Jurassic Eo-Cimmerian unconformity and the post-Elika Mesozoic succession. The non-conformity between the Gorgan Schists and the Shemshak Group is well exposed SW of Gorgan, close to the village of Ramedan. Here it consists of flat-lying coarse conglomeratic layers rapidly passing to marine bioclastic sandstone, unaffected by metamorphism. The Upper Jurassic Lar Formation and the Ghalimoran Limestone also cover the metamorphic complex. Eastwards (Neka Valley), the Shemshak Group is absent, and both the Palaeozoic units and the Gorgan Schists are directly covered and sealed by the Ghalimoran Formation (Figs 4 and 5). The absence of the Mesozoic pre-Upper Cretaceous succession suggests the occurrence of topographic highs representing reliefs of the Eo-Cimmerian belt, which were surrounded by a shallow sea for most of the Mesozoic age (Berra et al. 2007).

The Talesh Mountains (western Alborz) Although evidence of the Eo-Cimmerian orogeny is clearly recognizable in the Talesh Mountains (western Alborz), the strong tectonic imprint acquired during the Tertiary and Quaternary evolution of the belt strongly obscures the geometrical relationships among the Eo-Cimmerian units. Primary contacts among the internal and external units of the belt are poorly preserved in comparison to the central-eastern Alborz. We have focused our work (Fig. 8) on two main Eo-Cimmerian tectonic units located west of Rasht (Geological Survey of Iran 1975, 1977). They include: (1) the Shanderman Complex; and (2) a low-grade metamorphic rock association with Late Palaeozoic slates and limestones. The non-metamorphic Shemshak Group, dating to the Jurassic (Davies et al. 1972), nonconformably covers both of the units. The Shanderman Complex. The Shanderman Complex crops out in small erosional windows

42

A. ZANCHI ET AL.

Fig. 7. Post-Cretaceous deformation in the middle part of the Neka Valley (sites NEKA4 and NEKA5). Cretaceous rocks cover folded marly limestones of the Qezel Qaleh Formation (Carboniferous). The Mesozoic unconformity is locally faulted (dashed great circle) along the hinge of the post-Cretaceous folds (NEKA5a).

THE EO-CIMMERIAN OROGENY IN NORTH IRAN

43

Fig. 8. Simplified geological map of the Talesh Mountains (western Alborz); modified from Geological Survey of Iran (1975, 1977) based on our fieldwork. Traces of the cross sections (Fig. 10) and location of the structural observations (Fig. 11) are shown in the map.

44

A. ZANCHI ET AL.

below the Mesozoic units between the Rud-e-Masal (Masal Valley) and Asalem, 50 km to the north (Figs 2 and 8). Conglomerates and sandstones of the Shemshak Group non-conformably cover the Shanderman Complex in most of the area. The basal conglomerates of the Shemshak Group contain clasts of serpentinites, eclogites and other metamorphic rocks coming from the Shanderman Complex, testifying to its exposure during the Early Jurassic, after the end of the Eo-Cimmerian orogeny. The complex is generally separated from the Palaeozoic sedimentary cover of North Iran by steep strike-slip faults. The Shanderman Complex mainly includes mica schists and metabasites, with only minor calcareous

schists, quartzites and phyllites (Lachur Rud). The metabasites of this complex are affected by two main deformational events. The first one produced a pervasive foliation under eclogitic conditions; the second one is responsible for the crenulation of the previous structures, and it is locally associated with an axial plane foliation developed at amphibolite-facies conditions. Gabbro –dioritic intrusive bodies rich in ultramafic cumulates intrude the metamorphic basement of the Shanderman Complex after the main deformational events. These intrusive rocks often display a layered structure, including cumulitic dunites and peridotites. WNW–ESE- and NNW–SSE-trending highangle shear zones, a few to several metres thick and

Fig. 9. Field photographs of the Eo-Cimmerian structures in the Talesh Mountains. (a) The Eo-Cimmerian unconformity along the Shah Rud Valley just north of the village of Bellache´, right-hand side of the valley. The Shemshak Group basal conglomerates, containing pebbles of low-grade metamorphic and granitic rocks, unconformably cover deformed slates intruded by andesitic–gabbroic dykes. (b) Mountain-scale exposure of the Eo-Cimmerian unconformity, about 1 km NE of the summer village of Yeylagh, Siava Rud. Brownish rocks on the right-hand side of the photograph possibly represent lateritic soils developed along the unconformity. (c) Strongly deformed carbonatic metabreccias are exposed below the unconformity.

THE EO-CIMMERIAN OROGENY IN NORTH IRAN

with, respectively, left- and right-lateral motions, cross the whole complex, causing deep fracturing and alteration of the intrusives. The ultramafic cumulates often display a pervasive serpentinization that masks the original magmatic fabric. Mafic and ultramafic cumulates are also heavily serpentinized close to the contacts with the surrounding eclogites and mica schists. Poor outcrop conditions and a thick forest cover do not allow definition of the relationships between igneous rocks and the enclosing metamorphic rocks, even though the contacts seem to be of intrusive type with only minor shearing along the main faults. Thermobarometric estimates made on eclogites point to peak metamorphic conditions of 600 –700 8C and P . 1.5 GPa, with a subsequent re-equilibration at garnet- to epidote–amphibolite facies conditions (Zanchetta et al. 2009). Ar –Ar dating of white mica separates gives a Pennsylvanian age (315 + 9 Ma; Zanchetta et al. 2009),

45

slightly younger than the Middle –Late Devonian age obtained by Crawford (1977) for the phyllites of the Gasht Complex. Whole-rock chemical data of the Shanderman eclogites and mafic intrusives indicate a transitional to continental affinity (Zanchetta et al. 2009). These data, together with the age of metamorphism and the non-oceanic character of the serpentinized peridotites, here interpreted as altered ultramafic cumulates of igneous origin, suggest that the Shanderman Complex does not correlate to the Eo-Cimmerian Palaeotethys suture zone, as it may represent a nappe of European Variscan continental crust stacked on the north Iran margin. The Upper Palaeozoic nappes. A continuous belt of very-low-grade metamorphic and nonmetamorphosed but strongly deformed rocks, possibly belonging to the northern margin of the Iran block, occurs south and west of the Shanderman

Fig. 10. Geological cross sections across the Talesh Mountains showing Eo-Cimmerian and Tertiary deformations; traces in Figure 8. Shanderman Complex: Garnet– biotite micaschists (ms) and eclogites (ec). Masuleh-Shah Rud Unit: Permo-Carboniferous limestones (PCca), slate and phyllite (PCph). Boghrov Dagh Unit: Silurian– ?Early Devonian spilitic basalts (Sil-vs) with Silurian limestones with Nautiloids (ols), Carboniferous massive limestone with brachiopods (Cals) with ?Permian thin-bedded marly limestone (Pe), undifferentiated Upper Palaeozoic volcanics (Cavs). Sureh Khani Unit: Upper Devonian volcanoclastic deposits (vs) with thick andesitic lava flows (lf1) and bioclastic limestones (bls), uppermost Devonian–Carboniferous lava flows (lf2), Carboniferous coral and brachiopod massive limestone (cls), Carboniferous lava flows and breccias (lf3), Carboniferous marly limestone (cm) and limestone (cl) with brachiopods, Uppermost Carboniferous– ?Permian, thin-bedded limestone with marlstone and shale. Pre-Jurassic intrusives: serpentinized gabbro and cumulitic ultramafics (Gb-Um). Post-Eo-Cimmerian succession: Shemshak Group (Sh) with conglomerates (Sh-cg) and volcanics (Sh-v), Middle– Upper Jurassic Shal Formation (Sl), Lower Cretaceous fine-grained limestone with ammonoids (Kln), Upper Cretaceous marly limestone (Kls) with sandstone (s), Upper Cretaceous tuffaceous volcanics, basalts and andesites with marlstone (m), Karaj Formation (Eocene): green acid tuffs and dacitic lavas (kd), andesitic tuffs (kds), andesitic lavas (kl) and andesitic tuffs with lava and volcaniclastics (kt), Upper Oligocene?– Pliocene? Red Formation: red marlstone with gypsum (a), red and grey conglomerates (b), and sandstones (c).

46 A. ZANCHI ET AL.

Fig. 11. Stereoplots comparing the Eo-Cimmerian structures with the ones related to the late Tertiary thrusting and folding. Note that the Eo-Cimmerian structures show a different trend (N– S) with respect to the younger folds (NW– SE). The location of the sites is given in Figure 8.

THE EO-CIMMERIAN OROGENY IN NORTH IRAN

Complex (Fig. 8). We distinguished here two main tectonic units: (1) the Boghrov Dagh Unit, including a non-metamorphic Palaeozoic succession rich in volcanic layers; and (2) the Masuleh-Shah Rud Unit, with low-grade metapelites and Upper Palaeozoic metacarbonates. Closed folds with a welldeveloped axial plane cleavage occur everywhere in this unit around Masuleh, and along the lower part of the Shah Rud. The Shemshak Group unconformably covers part of this belt, indicating that these rocks were deformed during the Eo-Cimmerian event. Mountain-scale exposures of the unconformity occur around the Shah Rud downstream of Shal (Fig. 9a), around the village of Masuleh, along the Rud-e-Eskilit north of Masuleh, where an undeformed pre-Shemshak dioritic body also occurs (Fig. 8), and in the upper reaches of the Siava Rud gorge (Fig. 9b). Two geological cross sections describe the structural setting of this area. The first one (Fig. 10) crosses the South Boghrov Dagh thrust sheet, consisting of a Neogene SW-vergent overturned faulted fold, complicated by minor thrusts, deforming a peculiar Palaeozoic succession that is thrust southwards on the Mesozoic units. The age of the frontal thrust of the Boghrov Dagh structure is Tertiary in age as Neogene sediments are involved in these structures to the west along the continuation of the main fault. The Palaeozoic succession forming the hanging wall includes a Silurian Orthoceras limestone and thick Devonian basic lava flows and limestone, covered by a thick massive Carboniferous limestone rich in brachiopods, passing to thin layers of a possibly Permian fusulina-bearing limestone. The Shemshak Group

47

covers the succession with a low-angle unconformity. In the eastern part of the cross section, the thick Mesozoic units mask the Shanderman Complex east of the Urma Rud. The sheared serpentinitic bodies exposed along the NW–SE faults following the eastern margin of the Boghrov Dagh Unit (Geological Survey of Iran 1975) belong to the layered intrusive complex described in the previous section. The second cross section (Fig. 10) shows a lowgrade metamorphic unit that is very different in composition from the Boghrov Dagh thrust sheet to the north and from the units exposed to the south. It mainly includes slates with strongly recrystallized massive carbonate layers, in which Davies et al. (1972) identified Permo-Carboniferous faunas. Folds measured in the slate just above the village of Masuleh below the Shemshak Group, and along the Shah Rud in a similar position, show north –south-trending axes and vertical axial planes with a pervasive slaty cleavage (Fig. 11). These trends strongly differ from the dominating NW –SE trends of the Tertiary folds (Fig. 11). The Eo-Cimmerian unconformity crops out along the upper part of the Siava Rud, close to the summer village of Yeilagh (Fig. 9b). The Mesozoic cover present along the axis of the Shah Rud Valley generally hides the boundary between the two units. A NW–SE-trending vertical fault zone with strongly deformed tectonic slices of the Shemshak Group separates this nappe from the southernmost unit, consisting of an Upper Palaeozoic succession – here named Sureh Khani Unit – thrust onto the Tertiary successions exposed along the Gezel Owzen Valley. An overturned fold with a steep

Fig. 12. Tentative sketch of the relationships between the extensional structures formed in the Eo-Cimmerian foreland and the Eo-Cimmerian belt; modified from Zanchi et al. (2006). MG, medium-grade; HG, high-grade.

48

A. ZANCHI ET AL.

Table 1. Detrital modes of Shemshak sandstones* Shemshak Group

Sample

Kandavan

IR4

Kandavan Kandavan

IR5 IR6

Latitude

Longitude

Provenance

Grain size

Q

KF

P

Lvf

Lvm

Lcc

368110 4400

518190 1000

Recycled with volcanic detritus Recycled with volcanics Recycled with volcanics Recycled with volcanics Quartz recycled Slate arenite with volcanic and carbonate detritus Slate arenite with volcanic detritus Phyllite arenite Slate arenite Quartz recycled

270

61

0

2

9

1

0

220

58

1

1

10

2

0

220

60

1

3

9

3

0

220

63

2

4

15

1

0

110 270

79 27

0 1

2 3

3 12

1 6

0 7

300

27

1

2

12

8

0

150 140 50

46 18 95

1 0 1

1 3 3

2 1 0

2 0 0

1 0 0

0

00

0

00

0

00

36811 44 36811 44

0

00

0

00

0

00

51819 10 51819 10

Kandavan

IR7

36810 37

Masule` Masule`

IR8 IR9

378090 2300 378090 2300

488590 2900 488590 2900

Shah Rud

TZ2

378170 2100

488420 4400

Niala-Ramedan Niala-Ramedan Niala-Ramedan

GO32 GO31 GO29

0

00

36836 53 368360 5300 368360 5300

51818 37

0

00

53847 24 538470 2400 538470 2400

*In each of 10 coarse-siltstone to medium-grained sandstone samples, 400 –450 points were counted according to the Gazzi –Dickinson method (Ingersoll et al. 1984). Reported values are percentages of the total counted points. A detailed classification scheme allowed us to collect quantitative information on metamorphic rank of rock fragments (MI index; Garzanti & Vezzoli 2003). Thin sections were stained with alizarine red to distinguish calcite from dolomite. Mean grain size of studied samples (in microns) was determined by ranking and direct measurement in thin section. All compositional parameters are defined in Garzanti et al. (2006). Co-ordinates are from GPS (WGS84). Q, quartz (Qp. polycrystalline); F, feldspar (KF, K-feldspar; P, plagioclase); L, lithic grains (Lvf, felsic volcanic; Lvm, intermediate and mafic volcanic; Lcc, limestone; Lcd, dolostone; Lp, shale/siltstone; Lch, chert; Lms, very-low to low rank metasedimentary; Lmv, very-low to low rank metavolcanic; Lmf, medium to high rank felsic metamorphic; Lmb, medium to high rank mafic metamorphic; Luc, lizardite serpentinite; Lus, antigorite serpentineschist); Mu, muscovite; Bi, biotite; Pyx, pyroxene; Amp, amphibole; & HM, other heavy minerals. Reported values are percentage of the total counted points.

NE-dipping axial plane complicated by minor thrusting and disharmonic parasitic folding occurs in the hanging wall of the main thrust, which is Late Tertiary in age, as suggested by the occurrence of progressively folded Neogene beds in the frontal part of the structure along the Qezel Owzen Valley.

The Eo-Cimmerian structures in the central Alborz foreland The North Alborz Fault marks the boundary between the Eo-Cimmerian orogen and the stable foreland of north Iran, which now forms most of the central Alborz. To the south of this fault, the Shemshak Group covers the pre-Middle Triassic successions with a low-angle unconformity (Assereto 1966a) or conformably (Ghasemi-Nejad et al. 2004). According to most of the authors, the deposition of the Shemshak Group on the undeformed foreland marks the erosion of the Eo-Cimmerian orogen, growing due to the progressive collision of Iran with the southern Eurasian margin. Change from the carbonate platform of the Elika Formation to continental/marine siliciclastic depositional systems occurred between the end of the Carnian and the beginning of the Norian (Ghasemi-Nejad et al. 2004). Eo-Cimmerian graben and half-grabens occur within central Alborz, predating the deposition of

the Shemshak Group in the Shemshak region (Zanchi et al. 2006). Here an E –W-trending isolated ridge consisting of the Permo-Triassic carbonates of the Ruteh and Elika formations, previously described as an isolated thrust sheet (the Shahrestanak klippe of Assereto 1966b), has been now interpreted as a reactivated extensional structure together with the Gajereh half-graben, which is present a few kilometres to the north. The fact that the Elika Formation is confined within the two grabens, whereas outside the Shemshak Group directly covers the Palaeozoic succession, suggests that normal faulting occurred after the deposition of the Elika Formation and before deposition of the Shemshak Group clastics. Coincidence of the extensional event with the EoCimmerian orogeny suggests that it could have been an effect of the peripheral bulging of the foreland induced by the collision (Fig. 12). These extensional structures were inverted during the Meso-Cenozoic tectonics that affected the Alborz, accounting for many of the peculiar characters of the belt (Zanchi et al. 2006).

Petrographic analysis of the Shemshak sandstones The basal beds of the Shemshak Group were sampled in different portions of the Eo-Cimmerian

THE EO-CIMMERIAN OROGENY IN NORTH IRAN

49

Table 1. (Continued ) Lcd

Lp

Lch

Lms

Lmv

Lmf

Lmb

Luc

Lus

Mu

Bi

Pyx þ Amp

& HM

MI

0

18

2

7

0

0

0

0

0

0

0

0

0

35

0

17

3

6

0

2

0

0

0

0

0

0

0

46

0

15

1

7

0

1

0

0

0

0

0

0

0

47

0

9

0

3

0

4

0

0

0

0

0

0

0

69

0 11 0

0 0 7

0 0 0

10 29 36

0 1 4

2 4 2

0 0 0

0 0 0

0 0 1

2 0 0

0 0 0

0 0 0

0 0 0

158 126 111

0 0 0

0 0 0

0 0 0

12 55 0

0 2 0

31 14 0

0 0 0

0 0 0

0 0 0

4 3 0

1 4 0

0 0 0

0 0 0

300 221 n.d.

belt. Descriptions of the samples are reported in the Appendix. The petrographical composition of the Shemshak sandstones (Table 1) is distinct in the three sampled areas: (1) Kandavan Pass (external foreland; samples IR4, IR5, IR6 and IR7); (2) Masuleh –Talesh (internal part of the belt; samples TZ2, IR8 and IR9); and (3) Neka Valley (internal part of the belt; samples GO29, GO32 and GO33).

Provenance interpretation Quartzolithic composition of all Shemshak Group sandstones (Table 1) indicates a ‘recycled orogenic provenance’ (Dickinson 1985; Garzanti et al. 2007). In the Kandavan area, the Shemshak sandstones represent the relatively distal remnant of the Eo-Cimmerian foreland basin (Fig. 13a). Their homogeneous quartzolithic composition with diverse types of rock fragments suggests deposition by a significant palaeoriver, draining a considerable part of the Eo-Cimmerian belt. The latter, during Shemshak times, largely consisted of cover units including volcanic, sedimentary and very-lowgrade metasedimentary rocks. Composition of the Kandavan sandstones differs from the quartzolithic Shemshak sandstones in the Shemshak type area (Assereto 1966a, p. 1144) because the latter are more quartzose and lack volcanic rock fragments, indicating provenance chiefly from sedimentary cover units. In the Talesh area, the Shemshak sandstones non-conformably overlie a thrust stack characterized by very-low- to low-grade metasedimentary rocks and are associated with coarse-grained conglomeratic deposits. Their lithical composition (Fig. 13b, c) indicates relatively local provenance from cover units, including very-low-grade metasedimentary rocks as well as unmetamorphosed

felsic volcanic and sedimentary rocks (dolostone, shale/sandstone and chert). In the Ramedan area (Fig. 13d), the Shemshak sandstones and very coarse-grained conglomerates non-conformably overlie the Gorgan Schists. Their composition is virtually pure metamorphiclastic. Metamorphic grade of metapelite/metapsammite source rocks varies from anchimetamorphic to lowergreenschist facies up-section. This is consistent with deposition in proximal settings by a minor transverse river. Compositional variability of Shemshak sandstones does not only depend on provenance, but also markedly on grain size. Analysed very-finegrained sandstone in the Talesh area (IR8) and coarse siltstone in the Ramedan area (GO29), in fact, display a strong to very strong enrichment in monocrystalline quartz (probably largely recycled from older quartzose sandstones) at the expense of rock fragments.

Discussion The data collected in the Alborz and Talesh mountains provide a new interpretation of the remnants of the Eo-Cimmerian orogen. In the Neka Valley, south of Gorgan, a complete section of the Eo-Cimmerian belt is exposed. The section includes the Gorgan Schists, a Lower Palaeozoic metasedimentary and metavolcanic succession that was deformed and metamorphosed around 200 Ma, as well as the Upper Palaeozoic and Lower –Middle Triassic successions of North Iran, which are strongly deformed but lack a significant metamorphism. The occurrence of generally flat-lying undeformed Mesozoic successions (from the Shemshak to the Santonian Ghalimoran Formation) unconformably overlying these units testifies that their deformation is Eo-Cimmerian in age. Our new interpretation is in disagreement with previous works, such as that of Alavi (1996), who suggested

50

A. ZANCHI ET AL.

Fig. 13. Lithic fragments in the Shemshak sandstones. (a) Quartzolithic sandstone containing both radiolarian chert (Ch), and felsic volcanic grains (Fv; Kandevan, IR4); (b) lithic sandstone rich in dolostone grains (Do; Masuleh, IR9) authigenic calcite is stained red; (c) lithic sandstone rich in shale/slate grains (IR8; Masuleh, TZ2; Shah Rud); (d) lithic sandstone rich in phyllite grains (Ph; Ramedan, GO31). Scale bar ¼ 250 mm. All photographs are taken with cross-polars.

the occurrence of a low-angle normal detachment between the Cretaceous undeformed formations and the underlying units. However, a basal sandy layer, rich in glauconitic grains, occurs in most of the analysed outcrops at the base of the Upper Cretaceous unit, and no evidence of regional faulting has been observed (Berra et al. 2007). Although Silurian volcanic rocks occur in the Palaeozoic units of the northernmost part of Iran (Jenny & Stampfli 1978), no thick terrigenous successions similar to the one forming the Gorgan Schists is known to the south, preventing a direct correlation with the Palaeozoic units of the Iranian margin. In the Talesh Mountains, the structural framework is much more complex, as Tertiary and Quaternary thrusting and faulting strongly dismembered the Eo-Cimmerian orogen. However, the occurrence of very-low-grade Upper Palaeozoic metasediments forming the Masuleh–Shah Rud tectonic unit and lying below the poorly deformed Shemshak

Group unequivocally indicates the occurrence of an important Eo-Cimmerian orogenic event also in the Talesh Mountains. This unit that crops out in a more external position with respect to the Boghrov Dagh thrust sheet – which does not record an Eo-Cimmerian metamorphic event – is problematic. Out-of-sequence thrusting or intensive faulting at the end of the Eo-Cimmerian event could explain this intricate structural relationship. Additional information comes from the analysis of the Shanderman Complex (see also Zanchetta et al. 2009), which includes gabbros with ultramafic cumulates intruding an eclogite-bearing basement. Magmatic rocks do not show a significant metamorphic overprint. The association of eclogitic rocks with a metapelitic complex point to continental crust rather than to subducted oceanic sea floor, as previously suggested by Alavi (1996). The exhumation of the Shanderman Complex definitively occurred at the end of the Eo-Cimmerian orogeny,

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as suggested by the Early Jurassic age of the Shemshak Group (Clark et al. 1972), which unconformably covers these rocks. In addition, Ar– Ar radiometric ages (Zanchetta et al. 2009) suggest a Pennsylvanian age for the eclogitic peak metamorphic conditions. High-pressure rocks with Late Devonian and Carboniferous ages (Philippot et al. 2001; Kazmin 2006; Saintot et al. 2006) occur in the Caucasian region. These authors relate the high-pressure event to the accretion of a microplate detached from Gondwana at the beginning of the Palaeozoic to the southern margin of the Scythian Platform, with the closure of the Protothethys Ocean. The colliding microplate included two main blocks separated by a back-arc basin. The Makera block was located to the north and the Pontian– Transcaucasian block to the south, the latter being the northern active margin of the Palaeotethys during the Middle Carboniferous (Kazmin 2006). High-grade metamorphics of the Chorchana –Utslevi zone intruded by ‘Variscan granitoids’ are described in the Dzirula Massif of the Transcaucasian region (Sengo¨r 1990). This unit also includes Silurian– Devonian phyllites and sheared serpentinites, possibly deriving from a dismembered ophiolite complex. Further evidence of a Carboniferous orogenic activity comes from the Khrami Salient and Loki Massif belonging to the same area, where, respectively, andalusite–sillimanite gneisses and greenschists –epidote amphibolite metamorphic rocks are intruded by Lower and mid-Carboniferous granitoids. The radiometric ages summarized by Sengo¨r (1990, and references therein) suggest that metamorphism, magmatism and deformation span the entire Carboniferous. We thus suggest that the Shanderman and Gasht Complex may represent fragments of large nappes of continental crust coming from the Transcausian area, or from its lateral extension along the presentday region of the South Caspian Sea. These nappes were overthrusted southwards on the northern margin of north Iran during the Eo-Cimmerian orogeny. The polyphase formation of the Caspian Sea (Brunet et al. 2003, 2007) following the Eo-Cimmerian collision and the Neogene SSW-vergent thrust stacking may account for the large displacement of these units from their initial position. N–S-trending dextral strike-slip faults (Figs 2 and 8), forming the western margin of the Talesh Mountains, have probably accommodated the later displacements of the thrust sheets. The Eo-Cimmerian compressive structures are restricted to the Talesh Mountains and to the Gorgan region. Central Alborz corresponds in large part to the foreland of the Eo-Cimmerian belt (Zanchi et al. 2006), where the formation of extensional structures during Middle– Late Triassic

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is related to peripheral bulging along the southern margin of the foreland basin (Fig. 12). Here the Shemshak Group conformably covers the Lower – Middle Triassic carbonates of the Elika Formation (Ghasemi-Nejad et al. 2004) or it overlies it with a low-angle unconformity. The petrographic composition of the Shemshak Group records the unroofing of the Eo-Cimmerian orogen since the Norian. Thickness variations, facies complexity and strong differences in the age of its base indicate a greater complexity for the Upper Triassic clastics, which are discussed by Fu¨rsich et al. (2009).

A comparison with the evolution of the Palaeotethys suture in NE Iran The evolution of the Alborz sector of the EoCimmerian orogen can be compared with the Palaeotethys suture zone (Fig. 1) extending between Mashad and Torbat-Jam (Sto¨cklin 1974; Geological Survey of Iran 1986, 1993, 1996; Alavi 1991; Boulin 1991; Ruttner 1993; Alavi et al. 1997). The Palaeotethys remnants are well exposed close to Mashad, in the Binalood Mountains, where an accretionary wedge with ophiolitic fragments unconformably rests below undated coarse conglomerates and sandstones, with plant remains dubitatively assigned to the Late Triassic–Early Jurassic time interval (Alavi 1991). The Middle Jurassic Kashaf Rud Formation in turn covers these deposits. To the south, the accretionary wedge extends eastwards to the Afghan Border (Fig. 2), forming an ESE– WNW-trending strip of pillow lavas, ultramafics, turbidites, Permian limestones and cherts intruded by Late Triassic –Early Jurassic granitoids (Geological Survey of Iran 1993). A thick succession of continental red conglomerates capped by latest Permian limestones rests on top of the external parts of the complex (Alavi et al. 1997). The northern part of the area includes the ‘Aghdarband basin’ (Figs 2 and 14), with Lower–Middle Triassic continental –deep-marine sediments and acidic volcanoclastic rocks at the top deposited in a deep fault-controlled intra- or back-arc basin (Ruttner 1993; Alavi et al. 1997). The Norian coal-bearing Miankuhi Formation covers the Aghdarband basin without a marked angular unconformity. The whole succession is strongly deformed in a north-verging thrust stack crossed by left-lateral strike-slip faults. The basement of the Aghdarband basin consists of Upper Palaeozoic very-low-grade metasediments and metavolcanics belonging to the Turan domain and related to the growth of volcanic arcs along the southern part of Eurasia (Alavi et al. 1997). The Middle Jurassic sediments of the Kashaf Rud Formation unconformably cover all the described units after their deformation.

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Fig. 14. Summary of the available data concerning the evolution of the Eo-Cimmerian orogen in the Alborz and of the Palaeotethys suture zone, based on our unpublished data and on Geological Survey of Iran (1986, 1993, 1996), Alavi (1991), Ruttner (1993), Alavi et al. (1997), Seyed-Emami (2003), Ghasemi-Nejad et al. (2004), Fu¨rsich et al. (2005) and Zanchi et al. (2006). Yellow indicates compression, grey indicates extension.

Evidence of a Palaeotethys suture with ophiolitic units sensu latu is thus restricted to the NE of Iran (Mashad region), whereas they are not preserved in the Alborz. Here the Late Triassic collision of North Iran with Eurasia resulted in the formation of a stack of metamorphic nappes entirely consisting of continental crust partially involved in a subduction process during the Carboniferous and possibly belonging to the southern portion of Transcaucasia. In addition, low-grade metamorphic rocks with Palaeozoic protoliths record a low-grade metamorphic event directly related to the Eo-Cimmerian orogeny. Radiometric ages constrain the event to between 250 and 200 Ma. The occurrence of a collisional orogen is also recorded by the deposition of the Shemshak Group, beginning in the foreland since the Norian (Fig. 14). From Mashad and to the east, accretionary phenomena are recorded from Permian to Middle Triassic. The last deformational event, possibly marking the collision of the northern part of Central Iran, occurs between the Rhaetian and the beginning of the Middle Jurassic, and is thus younger than in the Alborz. The different tectono-stratigraphic histories recognized in the Alborz and in the Mashad region can be mainly ascribed to the heterogeneous composition of the southern margin of Eurasia as well as to its irregular shape. The Late Palaeozoic

southern margin of Eurasia consisted of a mosaic of small and mobile continental blocks separated by accretionary wedges, back-arc basins and transform faults (Garzanti & Gaetani 2002). The occurrence of a rigid block possibly representing the southern part of Transcaucasia, which was accreted to Eurasia during the Late Palaeozoic, may account for an earlier Eo-Cimmerian collision of the western Alborz with the southern margin of Eurasia, causing a complete consumption of the Palaeotethys lithosphere. To the east, the mobile blocks of Late Palaeozoic volcanic arcs and back-arc basins south of Mangyslak, in front of the Mashad– Torbat Jam region, may have accommodated deformation in a wider collisional zone, which is still partially preserved. In this region the final deformation may have been delayed until the latest Triassic –earliest Jurassic, as it also affects the Norian Miankuhi Formation (Fig. 14). Other effects, such as the occurrence of a deep embayment along the northeastern edge of Iran, as well as a partially oblique collision, could also explain the different characters of the Eo-Cimmerian orogeny along the northern margin of Iran. This work was funded by the MEBE project (proposal 02-26, Meso-Cenozoic Evolution of the Alborz Mountain Range, Iran) and by an Italian MURST PRIN Project (2004– 2006, leader A. Zanchi) in collaboration with

THE EO-CIMMERIAN OROGENY IN NORTH IRAN the Geological Survey of Iran. Dr M. R. Ghasemi, Dr A. Saidi and their colleagues at the Geological Survey of Iran are warmly thanked for continuous help and assistance in the field. S. Molyneux’s contribution is by permission of the Executive Director, British Geological Survey (N.E.R.C.). I. Bollati counted sandstones at the Milano-Bicocca University. We thank the two reviewers (M. Allen and M. Brookfield) for useful suggestions and J. Granath for a complete final review of the manuscript. Comments by the editors of this special publication were also appreciated.

Appendix 1: Petrographic description of the Shemshak sandstones Kandavan sandstones Samples (IR4–IR7) were collected along the Karaj– Chalus road, 50 m above the basal layers of the formation in front of the coal mines located close to the road junction to Elika. The four upper-fine- to lower-medium-grained (2.1 + 0.1 F) sandstones were found to have homogeneous quartzolithic composition (Q 60 + 2, F 3 + 2, L 36 + 3; Qp/Q 32 + 7; Table 1). Volcanic (mostly felsic types; Lv 12 + 3, Lvm/Lv 9 + 4), sedimentary (mostly siltstone/shale with some radiolarian chert; Lp 15 + 4, Lch 2 + 1) and metamorphic lithic grains (mostly very-low- to low-rank metasedimentary; Lm 7 + 1, MI 173 + 45) were all significant (Fig. 13a). Plagioclase prevailed among the few feldspars (P/F 76 + 13). Micas and heavy minerals were found to be negligible.

Masuleh – Talesh sandstones The two lower-medium-grained sandstones (1.8 + 0.1 F), sampled along the road above Masuleh (IR8 and IR9) and along the Shah Rud downstream Kholur (TZ2), were found to have lithic composition (Q 27 + 1, F 3 + 1, L 70 + 0; Qp/Q 22 + 8). Low-rank metamorphic lithics (slate, phyllite; Lm 37 + 6, MI 177 + 21) were seen to prevail over volcanic (mostly felsic types; Lv 19 + 1, Lvm/Lv 17 + 4) and sedimentary lithics (Fig. 13b, c; dolostone or siltstone/shale, with minor chert; Lcd  17, Lp 7, Lch 1 + 0). Plagioclase prevailed among the few feldspars (P/F 75 + 0). Micas and heavy minerals were negligible. The very-fine-grained sandstone (3.2 F) sampled along the road above Masuleh was markedly enriched in monocrystalline quartz (Q 81, F 3, L 16; Qp/Q 11), and included dominant phyllite metamorphic lithics (Lm 12, MI 223), a few felsic volcanic lithics (Lv 4, Lvm/Lv 11) and muscovite (2%).

Neka Valley sandstones Samples were collected from the right-hand side of the Neka Valley, south of Gorgan, close to the village

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of Ramedan. The two lower-fine-grained sandstones (2.8 + 0.1 F) have lithic–quartzolithic composition (GO31 and GO32), dominated by low-rank (Fig. 13d; slate-arenite of Garzanti et al. (2006); Q 19, F 3, L 78; MI 227) to medium-rank metamorphic lithics (phyllite-arenite of Garzanti et al. (2006); Q 48, F 2, L 50; MI 320). Micas are common (muscovite 3.5 + 0%, biotite 2.5 + 2%), and a few volcanic grains occur (Lv 2 + 1). The coarse siltstone sample GO29 (4.3 F) is strongly enriched in monocrystalline quartz, and includes some plagioclase and no lithic grains (Q 96, F 4, L 0; Qp/Q 4, P/F 82).

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The Shanderman eclogites: a Late Carboniferous high-pressure event in the NW Talesh Mountains (NW Iran) STEFANO ZANCHETTA1*, ANDREA ZANCHI1, IGOR VILLA1,3, STEFANO POLI2 & GIOVANNI MUTTONI2 1

Dipartimento Scienze Geologiche e Geotecnologie, Universita` di Milano-Bicocca, Piazza della Scienza 4, Milano, 20126 Italy 2

Dipartimento di Scienze della Terra, Universita` di Milano, Via Mangiagalli 34, 20133 Milano, Italy

3

Institut fu¨r Geologie, Universita¨t Bern, 3012 Bern, Switzerland *Corresponding author (e-mail: [email protected])

Abstract: The Shanderman Metamorphic Complex, exposed along the Caspian foothills of the Talesh Mountain, western Alborz, Iran, has always been interpreted as an ophiolitic fragment of the Palaeotethys Ocean. According to our new data, this unit consists of metamorphic rocks mainly represented by garnet– staurolite micaschists with large bodies of metabasites containing well-preserved eclogitic-phase assemblages. The Shanderman Complex (SC) was later intruded at middle crustal levels by intermediate– basic intrusive bodies. New Ar/Ar ages of paragonitic white micas in equilibrium with the high-pressure assemblages have given a Late Carboniferous age (315 + 9 Ma). Our new data suggest that the SC was equilibrated in high-pressure conditions during an orogenic event that predates the Eo-Cimmerian orogeny by more than 100 Ma and that may be tentatively ascribed to the Variscan orogeny sensu latu. We suggest that the Shanderman Complex represents a fragment of the Upper Palaeozoic European continental crust. The occurrence of eclogites in these regions can be explained by two different hypotheses: (1) the SC high-pressure rocks can be related to the accretion of Gondwana-related Transcauscasian – Moesian microplate to the southern margin of Eurasia; or (2) the SC eclogites can represent a fragment of the Late Palaeozoic ‘Variscan belt’ sensu latu of central Europe, which has been translated eastwards during Permian along a dextral megashear zone taking from a Pangea-B to a Pangea-A plate configuration. This metamorphic unit was stacked southwards on the northern edge of the Iran Plate during the Eo-Cimmerian events occurring at the end of the Triassic. The eclogitebearing basement of the SC was finally exhumed at the end of the Eo-Cimmerian orogeny, as suggested by the composition of the basal layers of the Shemshak Group dated here Middle Jurassic, that cover the crystalline rocks of the SC along a regional non-conformity. The SC was probably displaced further southwards during the Mesozoic opening of the South Caspian Basin and the Tertiary thrust stacking and dextral shearing accompanying the formation of the Alborz intracontinental belt. Supplementary material: Available at http://www.geolsoc.org.uk/SUP18342.

Since the first reconstructions concerning the early Mesozoic evolution of Iran (Sto¨cklin 1968, 1974), the region between the northern flank of the Alborz Mountains and the South Caspian Sea coast has been interpreted as the trace of the Gondwana-related Iran Plate collision with the southern Eurasian margin (Sengo¨r 1979, 1984, 1990; Berberian & King 1981; Alavi 1996; Brunet et al. 2003). The collision of Iran and of other ‘Cimmerian’ blocks against southern Asia resulted in the Eo-Cimmerian orogeny, the most relevant collision process affecting Eurasia between the end of the Triassic and the beginning of the Jurassic, which led to the definitive closure of the Palaeozoic Palaeotethys Ocean.

The main evidence for the Palaeotethys suture is exposed in the southern part of the Kopeh Dagh belt, between the city of Mashad and the Afghan – Turkmenistan border to the east (Alavi 1991; Ruttner 1993; Alavi et al. 1997). Here a PermoTriassic accretionary wedge and Triassic volcanic arc successions separate Upper Palaeozoic slightly metamorphosed volcaniclastic units, belonging to the southern part of the Turan region, from the Gondwanan Palaeozoic succession of central Iran to the south. West of Mashad the suture zone is masked by Meso-Cenozoic sediments unconformably covering the Eo-Cimmerian structures. According to Alavi (1996), ‘Palaeotethys remnants’ discontinuously crop out close to the Iranian coast

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 57– 78. DOI: 10.1144/SP312.4 0305-8719/09/$15.00 # The Geological Society of London 2009.

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of the Caspian Sea in the Gorgan region and west of Rasht in Talesh Mountains, marking the position of the Palaeotethys suture and its general continuity along the northern edge of the Iran Plate. The Talesh Mountains of northwestern Iran are thus a key area to investigate the location and evolution of the Eo-Cimmerian collisional zone. The peculiar rock association of metabasites with serpentinized ultramafic rocks present in the Shanderman Complex (Davies et al. 1972; Clark et al. 1975), cropping out along the Caspian foothills west of the towns of Masal and Shanderman, has been generally interpreted as a possible fragment of the Palaeotethys lithosphere (Sengo¨r 1990; Alavi 1996) and has been considered as one of the main pieces of evidence of the Palaeotethys collision zone. In this work we revise in detail the meaning of the Shanderman Complex in terms of lithological composition, geochemistry and metamorphic evolution, in order to understand if it can be really considered as an ophiolitic fragment. In fact, our new findings of previously unknown eclogites associated to garnet metapelites, later intruded by intermediate-mafic intrusive bodies, suggest a continental-crustal affinity of this unit rather than an oceanic origin. New Ar/Ar radiometric ages obtained on high-pressure white micas have given a Late Carboniferous age for the eclogitic peak, which suggests that high-pressure conditions are related to an earlier orogenic cycle, before the Cimmerian event. The origin of the Shanderman Complex, interpreted as a possible allochthonous

nappe, from the Transcaucasian region and emplaced along the northern edge of the Iranian Plate during the Eo-Cimmerian orogeny is also discussed in this paper.

The geological setting of the Talesh Mountains and its tectonic evolution The Talesh Mountains, extending from Azerbaijan to the Sefid Rud south of Rasht, form the western part of the Alborz mountain system and flank the southwestern coast of the South Caspian Sea (Fig. 1). The belt has a sinuous shape, trending around east–west in Azerbaijan, turning north – south along the southwestern Caspian coast, and finally trending WSW –ENE west of Rasht. The Azerbaijan side of the belt, consisting of Tertiary – mainly Eocene – sedimentary and volcanic successions, has been described by Vincent et al. (2005). They interpret the Eocene evolution of this area in terms of back-arc extension or transtension following the Mesozoic opening of the South Caspian Basin to the north of the Neotethys subduction zone. The change to a compressional regime forming a high topographic relief occurred during the Oligocene and is marked by the deposition of coarse-grained siliciclastic deposits. The Iranian part of the belt, about 50 km in width, is more complex, and consists mainly of a strongly deformed and discontinuous sedimentary succession spanning the entire Phanerozoic time. A few slices of metamorphic rocks grouped into

Fig. 1. Simplified tectonic map of the South Caspian region.

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the Gasht and Shanderman complexes are exposed along the Caspian foothills of the belt. According to previous authors (Davies et al. 1972; Clark et al. 1975), this section of the Alborz belt records several distinct orogenic phases, from problematic Middle–Late Devonian tectono-metamorphic events to the Eo-Cimmerian orogeny, whose effects in the region are also described in detail in Zanchi et al. (2006, 2009). The Eo-Cimmerian event marking the collision of Iran with the southern margin of Eurasia is followed by a complex evolution, related to the opening of the ‘palaeo-Caspian basin’ through the Mesozoic and the beginning of the Cenozoic (Allen et al. 2002; Brunet et al. 2003, 2007). Compression dominated from Oligocene times, leading to the formation of the present-day mountain belt. The metamorphic basement includes two poorly known units showing different evolutions and compositions (Fig. 2). The Shanderman Complex, that is the main subject of this paper, is later described in more detail. According to previous authors (Davies et al. 1972; Clark et al. 1975), it consists of an intricate association of phyllites, gneiss and amphibolites, with small patches of serpentinized peridotite, showing a medium-grade metamorphism. The peculiar composition of this unit has led several authors (Sengo¨r 1984; Alavi 1996) to directly equate the Shanderman Complex to the ophiolitic succession of Mashad. The Gasht Complex shows a different composition and evolution from the Shanderman basement; it includes two separate units showing a different metamorphic evolution. Phyllites with slates and quartzite form the upper unit, which is believed to unconformably cover the lower unit. According to Davies et al. (1972) a metaconglomeratic layer with kyanite-bearing clasts separates the two units. The lower unit, which is generally poorly exposed, consists of amphibolite-facies gneisses with sillimanite, kyanite and staurolite, with minor amounts of quartzites and amphibolites; the basement of the lower unit has later been intruded by granitoids. A greenschist-facies retrogression generally affects the unit. Middle–Late Devonian whole-rock Rb –Sr ages of 382 + 48 and 375 + 12 Ma have been obtained by Crawford (1977) on phyllite sampled near Masuleh. The whole complex was later intruded by Cimmerian granitic rocks, which have given an Early Jurassic radiometric age (175 Ma: Crawford 1977). The Palaeozoic succession forms the axial part of the belt, consisting of Tertiary SW-vergent folds and thrust sheets, generally thrusted over the younger Meso-Cenozoic units. The Palaeozoic succession, which shows analogies with the units exposed in central-eastern Alborz and in Central Iran, has been related to the external part of the

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Gondwana-related Iranian plate by most authors (Davies et al. 1972; Clark et al. 1975; Stampfli et al. 1991), with a few exceptions (e.g. Wendt et al. 2005). A peculiar feature of the Talesh units is the occurrence of thick volcanic and volcaniclastic layers throughout the Palaeozoic and the occurrence of open-sea Orthoceras limestones in the Silurian beds. Permian units, where present, generally consist of fusulina-bearing limestone and marls. No Triassic carbonates are exposed in the Talesh Mountains. A thick succession of very-low-grade carbonates and metapelites with Upper Palaeozoic faunas (see also Zanchi et al. 2009), deformed during the Eo-Cimmerian orogeny, crops out in the central part of the belt. These units as well as the Gasht and Shanderman complexes are unconformably covered by the coal-bearing deltaic– shallow-marine siliciclastic deposits of the Shemshak Formation, that marks the growth and erosion of the Eo-Cimmerian orogen (Fu¨rsich et al. 2009). The basal Eo-Cimmerian ‘molasse’ is dated to the Norian in central Alborz, to the Sinemurian in the axial part of the Talesh Mountains, to the Aalenian northward (Clark et al. 1975; SeyedEmami 2003; Ghasemi-Nejad et al. 2004). The composition of the basal conglomeratic layers of the Shemshak Formation suggests that the metamorphic units were exposed at the surface at the end of the Eo-Cimmerian orogeny. The present-day complex structural setting was mainly acquired since Oligocene times and is defined by a system of imbricate SW-vergent thrust sheets exposing the Palaeozoic succession along ramp anticlines. North–south-trending rightlateral strike-slip faults favour the SW displacement of the thrust sheets, forming steep lateral ramps along the northern part of the belt. However, active faulting still occurs along the Caspian coast along a flat north –south thrust surface, which indicates eastwards thrusting of the belt on the South Caspian Sea (Allen et al. 2003). Left-lateral strikeslip motion to the south along a WNW –ESE active fault (Toram earthquake in Jackson et al. 2002) occurs in the Qezel Owzan Valley, south of the Talesh Mountains.

Geological setting of the Shanderman Complex The metamorphic and intrusive rocks of the Shanderman Complex crop out west of the towns of Shanderman and Masal along the deep valleys dissecting the eastern slopes of the Talesh Mountains. Isolated outcrops of metamorphic rocks are exposed northwards in small erosional windows below the Mesozoic sediments up to the town of Asalem. The structural position of this

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Fig. 2. Geological scheme of the NW Talesh Mountains region (detail from scheme of Fig. 1).

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complex is ambiguous owing to the lack of significant exposures, as the Caspian foothills of the Talesh Mountains are, in fact, densely forested and rocks are deeply weathered. Strike-slip and normal faults generally separate these rocks from the Palaeozoic successions cropping out to the west, whereas between the Masal and Lachur valleys the Palaeozoic sedimentary successions seem to tectonically cover the basement along low-angle surfaces. The Shanderman Complex mainly includes metabasites and micaschists, with only minor calcschists, quartzites and phyllites (Lachur Valley). Phyllitic rocks always show white mica porphyroclasts, suggesting that they have been originated by retrogression of the surrounding micaschists. In the Masal Valley the crystalline basement of the Shanderman Complex is represented by partly retrogressed garnet –staurolite micaschists with minor amphibolites, present as metric-scale boudins, and quartzite levels. The metamorphic basement was successively intruded by intermediate– mafic intrusive bodies, rich in ultramafic cumulates, mainly exposed in the Lachur Valley and in the Urma Valley, near the village of Shalerah, where the intrusives form a NW –SE narrow fault-bounded stripe within micaschists and amphibolites. Field structural analysis has shown the presence of at least three distinct deformational phases in the complex. The first phase produced a pervasive foliation under eclogitic conditions, the second one is responsible for crenulation of the previous structures, locally associated with an axial plane foliation, developed under amphibolite-facies conditions. The third deformation phase produces gentle –close folds with NW–SE-oriented axes and subvertical axial planes. The whole complex is cross-cut by high-angle and vertical WNW–ESE- and NNW –SSE-trending shear zones, a few to several metres thick, with, respectively, left- and right-lateral motions. Rocks are deeply fractured and altered close to shear zones, showing a pervasive serpentinization of the ultramafics. Ultramafic cumulates are also heavily serpentinized along the contacts with the eclogites and micaschists. Contacts between the igneous bodies and the metamorphic rocks were not observed owing to poor exposure. We suppose, based on general geometrical relationship, that the boundary is a partially faulted primary intrusive contact.

Structural evolution and petrography of the Shanderman metamorphic rocks The metamorphic rocks of the Shanderman Complex consist of mafic eclogites, garnet amphibolites and

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micaschists, with minor amounts of calcschists, phyllites and quartzites. Eclogites are present as bodies, several tens to hundreds of metres in size, cropping out within garnet–staurolite–kyanite micaschists near the village of Chadorpost (star in Fig. 1), along the road taking to Lachur. The metamorphic basement of the Shanderman Complex was affected by at least three different phases of ductile deformation. The first phase (D1) is responsible for the development of a pervasive axial plane foliation (S1) within the eclogites, marked by the preferred orientation of paragonite and clinopyroxene (Fig. 3a, d). The S1 foliation is clearly visible only within the eclogitic bodies, whereas it has been almost completely transposed during the successive deformation phases in the surrounding micaschists. Rare and poorly preserved relicts of the S1 foliation were observed at the microscale within garnet amphibolites, where pseudomorphs of plagioclase þ amphibole symplectite on former omphacitic clinopyroxene define a weak preferred orientation at a high angle to the S2 foliation developed during the amphibolite-facies retrogression. The successive event (D2) occurred at amphibolite-facies conditions. It is evident within micaschists and garnet amphibolites where an axial plane foliation, S2, represents the most pervasive fabric element at the meso- and the microscale. The D2 deformation phase did not deeply affect the eclogites, where only a weak crenulation of the S1 eclogitic foliation is locally present. The garnet amphibolites cropping out near the village of Varadeh and along the Lachur Valley are strongly foliated, with garnet present only as relicts. The foliation is defined by amphibole þ titanite þ zoisite and is pervasive on millimetric scale. In some samples a symplectitic association made of plagioclase þ amphibole has been observed, suggesting a former eclogitic mineral assemblage for the garnet amphibolites. The pervasive foliation shown by the amphibolites is probably correlated to the crenulation observed in the eclogites, which post-dates the eclogitic event. Structural relationships between eclogites and amphibolites have not been observed, but the presence of plagioclase þ diopside symplectites within the amphibolites suggest that the garnet amphibolites of Lachur and Varadeh have also experienced high-pressure conditions. Retrogression of the eclogitic assemblage in garnet amphibolites was accompanied by the development of a strong foliation, which is absent in well-preserved eclogite outcrops. Foliation S2 present in the garnet–kyanite– staurolite micaschists and the garnet amphibolites is locally intensively crenulated due to a successive deformation phase (D3). D3 is responsible for the development of tight– closed folds present both at

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Fig. 3. Microstructures of the Shanderman eclogites. (a) Garnet (grtII) and paragonite (pgI) porphyroblasts in a fine-grained matrix made of cpxII þ pgII þ czoII. (b) Garnet porphyroblast preserving a core (grtI) with inclusions of bluish-green amphibole (amI), quartz, chlorite and rutile of the M1 phase assemblage. (c) Large paragonite (pgI) and omphacitic clinopyroxene (cpxI) porphyroblasts of the M2 eclogitic assemblage. (d) Clinozoisite porphyroblasts (czoII) synkinematic with respect to S1 foliation; inclusions forming the internal foliation within the czoII crystals are of rutile, quartz and paragonite.

the cm- and the m-scale that are not associated to an axial plane foliation. The contacts between the eclogites and the surrounding micaschists and amphibolites are generally covered. Thus, structural relationships between the eclogitic bodies and the surrounding basement remain unclear, even if the pressure peak recorded by the eclogites should predate the main deformation event, D2, easily recognizable in micaschists and amphibolites. Nonetheless, a common tectono-metamorphic evolution of eclogites, garnet amphibolites and micaschists can be assumed, based on the occurrence of the S1 relicts and of diopside þ plagioclase symplectites within the garnet amphibolites, suggesting the derivation of these from former eclogites such as those exposed near Chadorpost. The eclogites display a S-tectonitic texture owing to the preferred orientation of white mica and clinopyroxene that define a pervasive foliation, the S1 in the above reconstruction. A poorly defined mineralogical layering is formed by mm-thick

metamorphically differentiated garnet- or clinopyroxene-rich layers. Eclogites have a rich mineralogy with clinopyroxene, garnet, paragonite, phengite, amphibole, clinozoisite, epidote, quartz, rutile, titanite, calcite, plagioclase, and minor apatite and chlorite. The texture is generally inequigranular owing to the occurrence of porphyroblasts, some millimetres in size, of garnet, omphacite, clinozoisite and paragonite, surrounded by a pervasive foliation made of the same phases, with the exception of garnet and the addiction of quartz and rutile. Garnet porphyroblasts display an inclusion-rich core and a almost inclusion-free rim, that had partly overgrown the S1 foliation. In the garnet cores inclusions of bluishgreen amphibole, epidote, white mica, quartz, rutile and calcite were recognized. Microstructural analyses allow the reconstruction of at least three different equilibrium-phase assemblages in the mafic eclogites of the Shanderman Complex. Mineral abbreviations in the text follow IUGS recommendations.

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The first assemblage (M1) is preserved within garnet cores and includes: amI þ ep þ phg þ rt þ qtz þ cal + chlðM1 Þ: This paragenesis is thought to predate the eclogite-facies stage and possibly indicates highpressure– low-temperature conditions within the epidote–blueschist facies. Garnet cores (grtI) preserving this phase assemblage (Fig. 3b) display a rounded shape with irregular rims, often characterized by clusters of tiny rutile and quartz inclusions. Textural relationships occurring between garnet cores and the inclusions of the M1 phase assemblage remain unclear, even if it is likely that grtI post-dates the growth of the M1 assemblage and predate the successive eclogitic stage. The successive phase assemblage (M2) grew at eclogite-facies conditions and consists of: grtII þ cpxI þ pgI þ czoI þ qtz þ rt þ cal ðM2 Þ: GrtII rims grew in continuity on grtI cores, usually with a euhedral habit. CpxI, pgI, rtII, qtz and cc inclusions were rarely observed within grtII. Omphacitic clinopyroxene (cpxI) paragonite (pgI) and clinozoisite (czoII) are present as millimetric porphyroblasts displaying no shape-preferred orientation. CpxI and pgI porphyroblasts (Fig. 3c) are almost inclusion-free and clearly predate the development of the S1 foliation. In several samples pgI porphyroblasts display evidence of recrystallization at the contacts with the finegrained pgII, individuating the S1 foliation. CzoI porphyroblasts preserve remnants of internal foliation made of qtz and rt, that is sometimes in continuity with S1. Owing to these characters, czoI is considered pre- to early-S1. The S1 foliation shows the same phase assemblage of the porphyroblast population and is made of fine-grained omphacitic clinopyroxene, paragonite, zoisite, quartz and rutile. Locally, garnet rims grew statically on the S1 foliation indicating that garnet growth outlasts the formation of S1. The sin-S1 phase assemblage consists of: grtII þ cpxII þ pgII þ czoII þ qtz + ttn ðM3 Þ: Eclogites are exceptionally well preserved and only a few samples display a non-pervasive re-equilibration at amphibolite-facies conditions. In this case plgI þ amII symplectites grew on cpxI and cpxII, whereas amII þ plgII coronae formed around some garnet porphyroblasts. In a few cases tiny epidote crystals (epII) have been found within amII þ plgII coronae around garnet. The presence of epidote in the phase assemblage, substituting

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for garnet, indicates pressure conditions below the garnet stability field, i.e. within the epidote– amphibolite facies. In two samples a greenschistfacies re-equilibration was observed with thin rims of actinolitic amphibole (amIII) and albite (amIII) growing on cpxI porphyroblasts.

Whole-rock data of eclogites and mafic intrusives The small plutons intruded into the Shanderman Complex are mainly Mg-gabbros, with minor ultramafic cumulates and rare acid differentiates. The intrusives are substantially undeformed and only slightly metamorphosed in greenschist-facies conditions. The magmatic texture is well preserved in most of the samples. The ultramafic cumulates are locally serpentinized, probably due to fluid circulation along various dm- to m-thick steep shear zones affecting the rocks. Supplementary material: GPS co-ordinates of the analysed samples are available in Table 1 at http://www. geolsoc.org.uk/SUP18342. The primary magmatic phase assemblage of the gabbros is: clinopyroxene þ plagioclase þ amphibole þ olivine þ spinel. In some samples epidote seems to be a primary mineral, whereas more often it is recognized as a secondary phase grown on clinopyroxene and amphibole. The ultramafic cumulates are mainly made of clinopyroxene and olivine (i.e. wehrlitic in composition) with minor orthopyroxene. Almost all the collected samples are olivinenormative (i.e. Mg-gabbros) and display a transitional– alkaline character based on major element whole-rock analyses (Fig. 4). The Kalhe – Sahara Gabbro displays a marked alkaline character with respect to the Lachur and Shalerah gabbros. The ultramafic cumulates plot in the gabbronorite field of the De La Roche diagram, i.e. towards the Mg-rich area. Supplementary material: whole-rock chemical data of the mafic intrusives, ultramafic cumulates and metabasites of the Shanderman Complex are available in Table 2 at http://www. geolsoc.org.uk/SUP18342. The high-field-strength (HFS) elements’ abundances and ratios point to a volcanic arc as the most suited tectonic setting for the emplacement of the mafic intrusives. Oceanic character of the arc crust is unlikely in that continental crust has to be preferred owing to the presence of magmatic epidote, which indicates medium pressures of crystallization, compatible with a mid-crustal setting. The rare earth element (REE) chondritenormalized patterns of the gabbros (Fig. 5b) are similar for all the samples, except for the RA10

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Fig. 4. R1– R2 classification diagram (De La Roche et al. 1980) of representative samples of the Shanderman eclogites and mafic intrusives. Eclogites (sample RA32) and most of the gabbros plot within the olivine gabbro field. Serpentinites from Shalerah and Lachur plot in the gabbronorite field.

(Shalerah Gabbro) which displays an evident enrichment in light REE (LREE) (about 80 times chondritic values for La and Ce) similar to ocean island basalt (OIB) or enriched mid-ocean ridge basalt (E-MORB). The REE patterns of the other mafic intrusives and metabasites are of the normal MORB (NMORB) type (Fig. 5b), with LREE less than 10 times chondritic values and the heavy REE (HREE) around or slightly above 10 times chondritic values. Interpretation of the incompatible element diagram is more difficult due to the enhanced dispersion of the values. Incompatible elements, and specially the large-ion lithophile elements (LILE), are known to be very mobile during metamorphism and hydrothermal alteration due to their high solubility in aqueous fluids. As the Shanderman intrusive rocks have been slightly metamorphosed at greenschist-facies conditions, and the circulation of aqueous fluid is demonstrated by the presence of serpentinized ultramafic cumulates, the interpretation of the incompatible element diagrams has to be considered tentative. The most evident differences lie in the LILE abundances with the Shalerah Gabbro

enriched in LILE compared with the Lachur and the Kale Sahara gabbros (Fig. 5a). A more or less marked negative Nb anomaly is present in the pattern, compatible with the interpretation of a volcanic arc based on the HFSE (HFS element) content and ratios. Whole-rock chemical analyses have also been performed on the garnet amphibolites (RA52) and eclogites (RA32). The major element and the trace element data of the eclogites and amphibolites are very similar to those of the intrusive rocks analysed. The Kale Sahara amphibolites display marked positive anomalies of K and Sr; the origin of these LILE anomalies is not well understood and a possible metamorphic origin needs to be taken into account because of the high solubility of these elements in aqueous fluids.

Mineral chemistry Quantitative mineral analyses were obtained with an ARL SEMQ microprobe and a JEOL 8200 Superprobe at the University of Milan, both equipped with WDS spectrometers. Operating

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Fig. 5. (a) Incompatible and (b) REE spider diagrams of the Shanderman mafic intrusives and eclogites. Chondrite values are from Nakamura (1974).

conditions were 15 kV and 15 nA. Natural silicates were used as standards and the resulting data were processed with a ZAF (atomic number (Z), absorption (A) and fluorescence (F)) correction procedure. All the analyses were collected using a focused beam, except for the micas for which a 5 mm-probe diameter was used.

Garnet The chemical formulae of garnet were recalculated on the basis of 12 oxygens; the amount of Fe3þ was determined by charge balance. Garnet is mainly a Grs –Pyr –Alm solid solution, with the Sps component always ,10%. Most

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Fig. 6. Compositional profile through a garnet porphyroblast. Garnet shows a continuous chemical zoning with a core enriched in Grs and Sps. The Pyr and the Alm components increase towards the rim.

garnet porphyroblasts display a continuous zoning with a Grs- and a Pyr-rich core, and an enrichment of Alm and Pyr components towards the rim (Figs 6 and 7). Cores containing inclusions of blueschistfacies assemblage are Alm49 – 62Grs32 – 22Pyr08 – 11 Sps10 – 05 in composition. Towards the rim Alm increases up to 65 mol%, Grs decreases below 20 mol% and the Pyr component reaches values of 20 –22 mol%. Compositional profiles and X-rays maps indicate that all the analysed garnet porphyroblasts are continuously zoned, with no discontinuities at the contact between the

blueschist-facies cores and the outer eclogitic rims. The continuous chemical zoning of garnet is supposed to be due to a re-equilibration during the eclogite-facies metamorphism that annealed previous chemical inhomogeneities or to a continuous garnet growth through the transition from blueschist to eclogitic conditions. Representative garnet microprobe data are listed in Table 1. Supplementary material: a complete list of garnet microprobe data is available in Table 3 at http://www.geolsoc.org.uk/SUP18342.

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Fig. 7. Concentration maps of Ca, Mg, Fe and Mn within garnet. Maps were collected with a dwell time of 100 ms and are 1024  1024, with each pixel 3  3 mm in size.

Clinopyroxene Clinopyroxene analyses were recalculated on the basis of 6 oxygens, the amount of Fe2O3 was determined by stoichiometric charge balance. Two different generations of clinopyroxene are present, both of them with omphacitic composition. CpxI porphyroblasts display a Na content ranging between Jd45 and Jd50, with XMg(Fetot) always close to 0.73. CpxI display a weak zoning with porphyroblast cores somewhat higher in Na content with respect to outer rims, with constant values of Jd48 – 50 for the cores and Jd45 – 49 for the rims (Fig. 8). CpxII, present within the foliated matrix (S1 foliation), is also omphacitic in

composition but with a lower Na content that ranges between Jd40 and Jd46. The XMg(Fetot) values are inhomogeneous, and range between 0.67 and 0.78. Representative electron microprobe analyses of clinopyroxenes from the Shanderman eclogites are given in Table 1. Supplementary material: chemical composition of omphacite from Shanderman eclogites are available in Table 4 at http://www.geolsoc.org.uk/SUP18342.

White mica White mica analyses were recalculated on the basis of 11 oxygens, considering all Fe as Fe2þ.

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Table 1. Representative microprobe analyses of eclogite facies minerals from the M2 and M3 phase assemblages IR24 cpxI

RA32 cpxI

IR24 cpxII

RA32 cpxII

IR24 grt (core)

RA32 grt (rim)

RA32 wmcaI

RA32 pgI

RA32 pgII

IR24 amI

RA32 amI

IR24 ep

IR24 czoI

IR24 czoII

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO MgO CaO Na2O K2O

57.75 0.06 12.05 0.02 2.45 2.31 0.00 6.86 11.23 8.23 0.01

57.72 0.00 11.89 0.03 2.52 2.43 0.04 6.93 11.04 8.21 0.01

56.26 0.04 10.32 0.01 1.51 2.62 0.08 8.47 13.78 6.83 0.01

56.35 0.07 10.96 0.00 0.67 3.35 0.01 8.36 13.44 6.87 0.00

37.91 0.11 20.93 0.06 0.69 28.92 1.25 2.02 8.97 0.01 nd

38.61 0.06 21.08 0.04 0.51 28.49 0.39 4.29 7.38 0.01 nd

51.28 0.16 27.11 0.03 – 2.69 0.01 3.36 0.01 0.78 9.79

47.82 0.05 38.97 0.07 – 0.49 0.00 0.21 0.30 6.85 0.93

47.46 0.05 39.37 0.01 – 0.41 0.03 0.16 0.40 7.04 0.57

45.58 0.46 11.63 0.04 15.64 4.85 0.07 7.92 8.17 3.74 0.12

45.58 0.46 11.63 0.04 15.64 4.85 0.07 7.92 8.17 3.74 0.12

38.32 0.02 23.55 0.02 12.89 – 0.12 0.04 23.08 0.01 0.00

39.47 0.11 28.50 0.08 6.29 – 0.05 0.20 23.11 0.01 –

39.36 0.25 28.01 0.00 9.00 – 0.11 0.32 21.38 0.03 –

Total

100.97

100.82

99.92

100.09

100.88

100.86

95.23

95.69

95.50

98.23

98.23

98.05

97.81

98.46

Si Ti Al Cr Fe3þ Fe2þ Mn Mg Ca Na K

2.02 0.00 0.50 0.00 0.06 0.07 0.00 0.36 0.42 0.56 0.00

2.02 0.00 0.49 0.00 0.07 0.07 0.00 0.36 0.41 0.56 0.00

2.00 0.00 0.43 0.00 0.04 0.08 0.00 0.45 0.52 0.47 0.00

2.00 0.00 0.46 0.00 0.02 0.10 0.00 0.44 0.51 0.47 0.00

3.00 0.01 1.95 0.00 0.04 1.91 0.08 0.24 0.76 0.00 –

3.01 0.00 1.94 0.00 0.03 1.86 0.03 0.50 0.62 0.00 –

3.42 0.01 2.13 0.00 – 0.15 0.00 0.33 0.00 0.10 0.83

3.04 0.00 2.92 0.00 – 0.03 0.00 0.02 0.02 0.84 0.08

3.02 0.00 2.95 0.00 – 0.02 0.00 0.02 0.03 0.87 0.05

6.61 0.05 1.99 0.00 1.71 0.59 0.01 1.71 1.27 1.05 0.02

6.49 0.07 2.19 0.00 1.76 0.54 0.01 1.60 1.25 1.09 0.02

3.04 0.00 2.20 0.00 0.77 – 0.01 0.00 1.96 0.00 –

3.05 0.01 2.59 0.00 0.37 – 0.00 0.02 1.91 0.00 –

3.03 0.01 2.54 0.00 0.52 – 0.01 0.04 1.76 0.01 –

Cations

3.98

3.98

4.00

4.00

8.00

7.99

6.14

6.88

6.91

15.00

15.00

7.98

7.96

7.92

S. ZANCHETTA ET AL.

Sample Stage

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symplectitic amphiboles displaying the higher values up to 0.90 (Fig. 9b). In several samples a thin rim of actinolitic amphibole, amIII, was observed around omphacitic clinopyroxene. A few amphibole analyses are reported in Table 1. Supplementary material: a more complete list of amphibole analyses from Shanderman eclogites is available in Table 6 at http://www.geolsoc.org. uk/SUP18342.

Epidotes

Fig. 8. Clinopyroxene compositions.

White mica, phg, preserved within garnet cores together with ampI, zoI, rt, qtz and cc (the M1 phase assemblage) displays a phengitic composition with Si up to 3.4 pfu (per formula unit) and a constant XMgFe2þ of 0.68 –0.69. Paragonite porphyroblasts, pgI, have a K/Na ratio comprised between 0.09 and 0.11, with XMgFe2þ ranging from 0.43 to 0.51, with the higher values usually in the cores. PgII, the white mica that defines the S1 foliation, displays a similar composition with somewhat lower K/Na ratio (0.05–0.9) and similar XMgFe2þ. The Ca content of paragonite ranges between 0.25 and 0.30 pfu for pgI, and between 0.25 and 0.40 for pgII. Representative white mica analyses from the eclogites are given in Table 1 and are available as supplementary material in Table 5 at http://www.geolsoc.org.uk/ SUP18342.

Amphibole Amphibole analyses were recalculated on the basis of 23 oxygens and 13 cations þ K þ Na þ Ca. AmI, preserved within garnet cores, is barroisitic in composition with Na content ranging between 1.01 and 1.41 apfu (atoms per formula unit), which means a Na amount in the B site between 0.72 and 1.03 pfu (Fig. 9a). In some cases amI inclusions in garnet display a weak zoning, with Na content decreasing towards the rims. XMgFe(tot) values range between 0.40 and 0.43. AmII formed during amphibolite-facies re-equilibration, and it is present together with plagioclase within symplectites on former omphacitic clinopyroxene, cpxI and cpxII, or within amphibole–plagioclase coronae around garnet porphyroblasts. AmII is tschermakitic in composition with high XMg(Fe2þ), usually .0.75, and with

Epidote analyses were recalculated on the basis of 12.5 oxygens considering all Fe as Fe3þ. Epidote (ep) preserved within garnet cores belong to the M1 blueschist-facies assemblage, and show a Fe3þ content between 0.75 and 0.80 pfu with no Mn (Mn , 0.02 pfu). CzoI and czoII show a similar composition, with no Mn (Mn , 0.03 pfu) and Fe3þ ranging between 0.37 and 0.45 pfu, with czoI displaying invariably the lower values, 0.37 – 0.40, and czoII the higher ones, 0.40 –0.45. Representative analyses of different epidote generations are listed in Table 1 and are available as supplementary material in Table 7 at http://www.geolsoc.org. uk/SUP18342.

Other minerals Plagioclase formulae were recalculated on the basis of 8 oxygens pfu. PlgI and plgII occurring with amphibole in symplectites around garnet, display a similar composition with An content around 0.15 –0.18 mol%. The third generation of plagioclase, plgIII, supposed to have grown during the latest stages of exhumation display an albitic composition, with Ab ¼ 97 mol%. Chlorite is present within garnet cores and it is supposed to belong to the M1 blueschist-faciesphase assemblage. Chlorite compositions are represented by clinochlore–daphnite solid solutions, with minor deviations towards the sudoite endmember. The mean XMg values range between 0.50 and 0.61, with the exception of chlorites from the RA32 sample which display lower XMg, within the range of 0.39– 0.47. Titanite, rutile, apatite and rare ilmenite have been found as inclusions in garnet and in the matrix without any significant compositional variations. Chemical formulae of titanite were recalculated considering all Fe as Fe3þ (Oberti et al. 1991). The XAl of titanite, where XAl ¼ Al/(Al þ Ti þ Fe3þ), is low and has constant values of 0.04 –0.06. Supplementary material: analyses of plagioclase, chlorite and titanite from the Shanderman eclogites are available in Table 8 at http://www. geolsoc.org.uk/SUP18342.

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Fig. 9. (a) Na content in the B site v. Si content of amphibole. (b) XMg v. Si content. Amphibole classification on both diagrams is after Leake et al. (1997).

Metamorphic evolution of the Shanderman eclogites Pressure– temperature equilibrium conditions of the different metamorphic phase assemblages in the paragonite-bearing eclogites of the Shanderman Complex were estimated by means of available geothermometers and geobarometers, with additional constraints from available experimental data. The peculiarity of the Shanderman eclogites is the preservation of a prograde P–T path from the blueschist to the eclogite facies. The

M1 phase assemblage, aml þ ep þ phg þ rt þ qtz þ cc + chl, indicates high-pressure– low-temperature conditions roughly constrained to 0.7–1.1 GPa at temperatures of 400–550 8C. The pressure interval is defined on the basis of experimentally determined (Maruyama et al. 1986) and calculated (Evans 1990; Kerrick & Connolly 2001) mafic blueschist-phase relations, which indicate the disappearance of albite with increasing pressure in favour of sodic-amphibole þ clinozoisite/epidote-bearing assemblages at 0.6– 0.8 GPa in a temperature range of 400–500 8C. The absence of sodic-clinopyroxene in the M1

THE SHANDERMAN ECLOGITES (IRAN)

assemblage indicates equilibrium pressure below 1.1 GPa at 400 –550 8C. Constraining the equilibrium temperature for the M1 phase assemblages is more difficult owing to the absence of reliable indicators. Considering garnet cores that preserve M1 inclusions not in equilibrium with them, the maximum temperature for the M1 stage is constrained by the garnet-in reactions, with increasing temperature within the 0.6–1.1 GPa pressure range, that sets an upper temperature limit for the M1 assemblage of around 450 –500 8C (Maruyama et al. 1986; Evans 1990; Okay et al. 1998). A successive pressure increase resolved in the stabilization of eclogitic conditions at which both M2 and M3 phase assemblages have grown. The M2 and M3 assemblages are both characterized by the absence of amphibole and the presence of paragonite, together with garnet, Na-clinopyroxene, clinozoisite, calcite, quartz, rutile (M2) and titanite (M3). Experimentally determined phase relations on H2O-saturated MORB-like systems (Schmidt & Poli 1998; Poli & Schmidt 2004) constrain well the P– T stability field of amphibole-absent and clinozoisite/zoisite-bearing eclogitic-phase assemblages. Amphibole breaks down with increasing pressure at 2.4–2.5 GPa in the 530 –700 8C temperature range, whereas (clino)zoisite is stable up to 3.1 GPa at approximately 700 8C (Poli & Schmidt 1995) for bulk-rock compositions with normative anorthite content (CIPW norm) around 31 wt%. The absence of amphibole in the M2 and M3 eclogitic-phase assemblages of the Shanderman eclogites could possibly represent a lower pressure

71

limit for the pressure peak, whereas the presence of (clino)zoisite is not so indicative due to the dependence of epidote stability on the normative anorthite content of the rock. An increased anCIPW content results in a much larger pressure range for epidote stability (Thompson & Ellis 1994; Wittenberg et al. 2002; Poli & Schmidt 2004) both towards upper and lower pressures. The anCIPW content of the Shanderman eclogites ranges between 33 and 34.5 wt%, so it is slightly higher than the 31 wt% anCIPW of Poli & Schmidt (1995), probably resulting in a higher than 3.1 GPa pressure limit for (clino)zoisite within M2 and M3 phase assemblage of the Shanderman eclogites. Suggested equilibrium pressure for the M2 and M3 phase assemblages consider as indicative the absence of amphibole, even if it may also be due to compositional effect. If this is the case 2.4 –2.5 GPa has to be taken as the maximum pressure, whereas a lower pressure limit is provided by the stability of paragonite (Holland 1979), as discussed below. The upper pressure boundaries for equilibrium conditions of the M2 and M3 phase assemblages can be derived by the presence of paragonite in equilibrium with garnet, clinopyroxene and clinozoisite. The paragonite-out reaction in Figure 10 is derived from Holland (1979): the paragonite breakdown, with a negative dP/dT slope, constrains the M2 and the M3 equilibrium paragenesis at pressure below approximately 2.5 GPa for temperatures between 530 and 650 8C. Following previous considerations and considering as indicative the absence of amphibole, the presence of paragonite brackets the pressure peak recorded by the

Fig. 10. Phase relations in hydrated MORB-like compositions (with anCIPW ¼ 31 wt%) with emphasis on (clino)zoisite-, epidote- and paragonite-bearing equilibria (Poli & Schmidt 2004). Reaction (1) and (2) are from Holland (1979); reaction (3) from Apted & Liou (1983). The P– T equilibrium fields for the M1, M2 and M3 phase assemblages of the Shanderman eclogites are shaded in grey.

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S. ZANCHETTA ET AL.

Shanderman eclogites at 2.3–2.5 GPa. The grossular content of garnet coexisting with zoisite and omphacitic clinopyroxene in the M2 and M3 assemblages is similar to those reported by several experimental data (Poli & Schmidt 1995; Molina & Poli 2000) for pressure in the range of 2–2.5 GPa and bulk compositions approaching that of the Shanderman eclogites, this represents an additional constraint supporting the lower pressure estimate in addition to the absence of amphibole. Temperature estimates for the M2 and M3 phase assemblages have been obtained by the garnet – clinopyroxene Fe2þ –Mg exchange geothermometer. Three different calibration have been applied: Ai (1994), Berman et al. (1995) and Krogh Ravna (2000). Supplementary material: analyses of the garnet –clinopyroxene pairs used for T estimates are available in Table 9 at http:// www.geolsoc.org.uk/SUP18342 together with results in Table 10 at http://www.geolsoc.org.uk/ SUP18342. Equilibrium temperatures obtained with the Ai (1994) and Krogh Ravna (2000) calibrations are lower with respect to those obtained employing the Berman et al. (1995) thermometer (Fig. 11). For pressures ranging between 2.3 and 2.5 GPa, the Ai (1994) and Krogh Ravna (2000) thermometers result in 510 –560 8C for the M2 phase assemblage and 440 –490 8C for the M3 phase assemblage. Temperatures obtained with the Berman et al. (1995) calibration are 50 –60 8C higher than the previous and fit better with the experimentally determined phase relations represented in Figure 10. The temperatures obtained with the Ai (1994) and Krogh Ravna (2000) thermometers indicate equilibrium conditions for the M2 stage and, especially, the M3 stage close or within the lawsonite stability field. As no evidence of lawsonite growth has been observed, results obtained with the Berman et al. (1995) calibration of the garnet – clinopyroxene Fe–Mg exchange

thermometer have to be considered more reliable than estimates made with the Ai (1994) and Krogh Ravna (2000) calibrations. Discrepancies between the three different calibrations arise from the different dP/dT slope of KD curves, with the Berman et al. (1995) calibration displaying a steeper slope compared with the other two. Such a difference results in minor discrepancies at high temperature, where the three calibrations give similar results (Krogh Ravna 2000), but at lower temperature the gap increases up to 60–70 8C. Following previous considerations, reliable temperature estimates for the M2 and M3 eclogitic-phase assemblages are, respectively, 575–615 and 505– 550 8C. Pressure –temperature estimates of the eclogitic assemblages indicate that the M3 stage occurred at lower temperature and lower pressure than the M2. The pressure indication is derived from the jadeite content of cpxII (c. 0.45), belonging to M3, which is somewhat lower than that of cpxI (c. 0.50) in the M2 assemblage. After the pressure peak some eclogite samples display a non-pervasive re-equilibration at epidote– amphibolite-facies conditions. Plagioclase –amphibole symplectites substituted for omphacitic clinopyroxene and amphibole þ plagioclase + epidote formed around garnet porphyroblasts. Epidote coexisting with plagioclase within coronae around garnets provides textural evidence of equilibrium pressures in the range of 0.4 – 0.8 GPa (Apted & Liou 1983). Temperature estimates for the re-equilibration stage have been performed with the amphibole–plagioclase thermometer of Holland & Blundy (1994) on amII and plgII substituting for garnet: estimates range between 550 and 630 8C for that pressure range. The latest stages of exhumation are recorded by the rare presence of actinolitic rims on omphacitic clinopyroxene, sometimes coexisting with albite and

Fig. 11. Temperature estimates for the grt– cpx pairs of the M2 and M3 phase assemblage. Ai94, Ai (1994); B95, Berman et al. (1995); K00: Krogh Ravna (2000).

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73

Fig. 12. (a) Ca/K v. Cl/K ratios of white mica separates from the Shanderman eclogites. (b) Ca/K v. step ages diagram. The Ca/K ratio of the less retrogressed samples, RA32 and RA24, plot in a narrow interval ranging between 1.17 and 1.26. (c) Cl/K ratios v. step ages. (d) Age spectrum diagram obtained by stepwise heating of paragonite separates of the Shanderman eclogites. The least altered sample, RA32, displays three steps, amounting to 76% gas release, ranging between 305 + 2 and 324 + 3 Ma.

by chlorite grown along fractures in garnet porphyroblasts.

Geochronology Three paragonite samples were analysed using 39 Ar – 40Ar stepwise heating. Samples were separated using heavy liquids and purified by handpicking under the microscope. Owing to sample size limitations, heterogeneities in the obtained separates could not be eliminated. These heterogeneities derive to a great extent from intergrown retrogression products, as discussed in the petrographic description. In order to unravel the isotopic signal of mixtures, the classic tool is the use of three-isotope common-denominator correlation diagrams (Villa

2001), because the minerals composing the mixture degas at different temperatures. As paragonite contains Ca in its stoichiometric formula, the most natural choice is use of the Ca/K ratio as abscissa and as ordinate the Cl/K ratio (a monitor of alteration) or step ages (Fig. 12a, b). The Ca/K v. age diagram was successfully used by Tomaschek et al. (2003) to estimate an age for a paragonite –phengite mixture. In Figure 12a, it can be seen that the data points for all three samples define a V-shape: one branch corresponds to variable Ca/K ratios with low Cl/K ratio, while a few points show an anticorrelated increase of Cl/K for decreasing Ca/K ratios. Such a distribution requires at least three different Ar reservoirs. A possible interpretation would identify the high-Cl, low-Ca points as alteration phases of the original Ca-rich

74

S. ZANCHETTA ET AL.

paragonite. The low-Ca, low-Cl points would represent a different alteration or impurity phase different from paragonite. Because the points of all three samples define similar V-shaped trajectories, we are encouraged to believe that while their relative mass balance varies amongst our samples, the identity of the non-paragonite contaminants is the same and reflects the same retrogression/alteration event. In Figure 12b, it can be seen that the step ages for the two less retrogressed samples, RA24 and RA32, correspond to Ca/K ratios in the narrow interval between 1.17 and 1.26. We interpret this ratio as the chemical signature of paragonite and the corresponding ages as those least affected by chemical perturbations. The influence of alteration on ages can be diagnosed using the Cl/K v. age diagram (Fig. 12c). This plot was used by Heuberger et al. (2007) to resolve the effect of secondary modification of a substoichiometric muscovite. For all three samples, the highest step ages correspond to Cl/K ratios between 0.002 and 0.008, a range of values that is typical of white micas. Alteration results in introduction of Cl and removal of Ar, as shown by the negative correlation slope shown by our data points. Extreme Cl/K ratios exceeding 1 correspond to almost zero-age secondary phases. On the left-hand side of the diagram, each sample features one high temperature step with low ratios ,0.0008. These could correspond to minor admixtures/intergrowths of a late-degassing, Cl-free phase such as (secondary young) feldspar. The age spectrum diagram (Fig. 12d) can now be reliably interpreted in terms of multiple generations of variably altered minerals. The least altered sample is RA-32, which features three steps (totalling 76% of the gas release) ranging between 305+2 and 324 + 3 Ma. Given the mutual discordance of these three steps, a conservative age estimate for the paragonite-forming metamorphism is 315 + 9 Ma.

Discussion and conclusions Our data support the interpretation of the Shanderman Complex as a deeply subducted slice of continental crust intruded during exhumation by intermediate –mafic intrusives, in places rich in ultramafic cumulates. Pressure –temperature estimates made on paragonite-bearing mafic eclogites point to peak metamorphic conditions up to 2.3– 2.5 GPa at 575 –615 8C (M2 phase assemblage) indicating a subduction depth of 70 –80 km. In several samples, pre-eclogitic phase assemblages (M1) are well preserved within garnet cores and are thought to record first stages of subduction at

0.7–1.1 GPa and 400–550 8C within the epidote – blueschist facies. Successive re-equilibration occurred within the epidote–amphibolite facies at 0.4–0.8 GPa and 550–630 8C. The consensus of the data set indicates a clockwise P –T–t path for the Shanderman eclogites with a minor cooling event, represented by the transition from stage M2 to M3, before the onset of exhumation. 39 Ar – 40Ar radiometric dating of paragonites belonging to the M2 and M3 phase assemblages gives a minimum age for the eclogite-facies metamorphism of 315 + 9 Ma, indicating that the metamorphic basement of the Shanderman Complex was subducted to crustal depths during the Late Carboniferous. Whole-rock data of mafic intrusives and ultramafic cumulates of the Lachur, Masal and Urma valleys, intruded within the metamorphics of the Shanderman Complex, display a transitional to continental affinity, possibly related to a volcanic arc setting. The intrusives are only slightly metamorphosed, with local serpentinization of ultramafic cumulates occurring along shear zones, and weak and a non-pervasive greenschist metamorphic overprint present on gabbros and diorites. Such data indicate that the intrusion occurred probably after or during the amphibolite-facies re-equilibration recorded by the eclogites and micaschists of the Shanderman basement. The conclusion contrasts with the commonly accepted interpretation of the Shanderman Complex as part of an ophiolitic belt delineating the Palaeothetys suture extending from the Talesh Mountains of NW Iran eastwards to the ophiolites of the Mashad–Binalood Complex (Sengo¨r 1984; Alavi 1996). Our data suggest the metamorphic basement of the Shanderman Complex is a part of a Palaeozoic orogenic belt, the remnants of which are presently exposed within the Shanderman Complex itself and the Gasht Complex (Davies et al. 1972; Clark et al. 1975; Crawford 1977). Several lines of evidence point to relevant orogenic activity occurring during the Carboniferous in the Transcaucasian region along the southern part of the Lesser Caucasus (Sengo¨r 1990). In the Dzirula Massif of the Transcaucasian region, highgrade gneisses, amphibolites and migmatites at the external boundaries of the Chorchana –Utslevi Zone are intruded by ‘Variscan granitoids’, with K/Ar ages mainly clustering in the Late Carboniferous (Adamia et al. 1983). Silurian –Devonian phyllites, gabbros, metavolcanics and sheared serpentinites with Cambrian marble slivers, possibly representing a dismembered ophiolite unit, form the Chorcana –Utslevi Zone, which also shows Late Carboniferous Rb/Sr ages for the metapelites. A similar succession is exposed further east in the Khrami Salient and Loki Massif, where

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Fig. 13. Late Carboniferous– Early Permian reconstruction of Pangea obtained by using paleomagnetic poles and Euler poles of rotation for Gondwana and Laurasia outlined in Angiolini et al. (2007). The Transcaucasian region (T), locus of origin of the Shanderman Complex, is indicated. See the text for discussion.

medium- to high-grade gneisses and greenschist to epidote–amphibolite metamorphic rocks are, respectively, intruded by Lower Carboniferous and Upper Carboniferous granitoids. K/Ar ages of 287 + 15, 319 + 11 and 333 + 11Ma have been obtained on muscovites from the metamorphics of the Loki Massif (Abesadze et al. 1982). Similar ages spanning from 325 to 338 Ma come from the granites intruding the metamorphic basement. Moreover, high-pressure eclogitic rocks, Late Devonian –Carboniferous in age (Philippot et al. 2001; Kazmin 2006; Saintot et al. 2006), also occur to the north of the Great Caucasus range. To the south and east of the Black Sea, evidence of a Late Palaeozoic orogenic event are still widespread as testified by the occurrence of high-grade (amphibolite–granulite facies) metamorphic basement rocks in the Pulur, Devrekani, Uludag˘ and Kazdag˘ massifs of the Sakarya Zone of the Eastern and Central Pontides (Okay et al. 2006; Topuz et al. 2007). Isotopic ages from high-grade gneisses constrain the age of metamorphism of the Pulur and Kazdag˘ massifs to 300–330 Ma (Okay et al. 2006 and references therein). Other evidence of Carboniferous orogenic activity has also been reported recently from the Jandaq metamorphic complex (Bagheri & Stampfli 2008) in central Iran, south of the Great Kavir Fault, where Ar/Ar mica ages from greenschists and amphibolites give a 320 –333 Ma age for the main metamorphic event. In order to gauge the plate tectonic scenario responsible for the Late Palaeozoic age of the high-pressure metamorphism recorded in the

Shanderman Complex, we reconstructed the configuration of Pangea using selected palaeomagnetic data of Late Carboniferous–Early Permian age from stable Africa and Europe, as outlined in Angiolini et al. (2007). This Late Carboniferous – Early Permian Pangea is remarkably similar to Pangea B of Irving (1977) and Morel & Irving (1981), with Africa placed south of Asia and South America south of Europe (Fig. 13), as also confirmed for the Permian by subsequent palaeomagnetic analyses (Muttoni et al. 1996, 2003; Torcq et al. 1997; Irving 2005). Pangea B transformed into a more classic Wegenerian A-type configuration, with Africa immediately south of Europe and South America south of North America, by means of a dextral motion of Laurasia relative to Gondwana that took place during the Early Permian broadly along the Variscan suture (Muttoni et al. 2003, 2004). An increasing body of biogeographical and geological data seems to support Pangea B and its transformation to a Pangea A-type configuration during the Permian (Angiolini et al. 2007; Muttoni et al., submitted). Focusing on arguments more pertinent to this study, Pangea B provides a symmetrical framework for the Appalachian – Variscan fold belt, placing the Appalachians of eastern North America opposite to the northwestern South America fold belt, and the Variscan fold belt of Europe opposite to the northwestern Africa fold belt of the Mauritanides, as originally noticed by Irving (1977) and Morel & Irving (1981). Within Pangea B, Carboniferous high-pressure deformation

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of the Shanderman Complex finds a logical geodynamic context of Variscan convergence of Gondwana relative to Laurasia, which would otherwise be far from obvious assuming instead Variscan convergence in a Pangea A-type configuration where the Paleotethys Ocean bounded to the south the European margin at the longitudes of Transcaucasia. The high-pressure metamorphic event of Transcaucasia was related by different authors to the accretion of a microplate, termed Makera-Pontus/ Transcaucasia, to the European margin of the Scythian Platform. This microplate has been interpreted as a block possibly detached from Gondwana at the beginning of the Palaeozoic (Kazmin 2006), or alternatively as a Europe-related rifted fragment belonging to the southern margin of the Scythian Platform (Saintot et al. 2006). Acknowledging these hypotheses, we propose as an alternative and somewhat easier solution to interpret this and other compressional events of Late Paleozoic age as due to the convergence of Gondwana relative to Laurasia in a Pangea B-type configuration. The Shanderman Complex can thus represent an allocthonous nappe coming from the southern margin of the Transcaucasian region, locus of original Variscan deformation. The nappe was displaced southwards and exhumed during the Eo-Cimmerian orogenesis after being stacked above the northern margin of the Iranian Plate. Its present position was acquired during the late Tertiary formation of the western Alborz belt and especially during the Neogene, when dextral shearing was active along north– south- to NNE– SSW-trending faults, which formed the lateral ramps of the SW-vergent fronts of the main thrust sheets present in the Iranian side of the Talesh Mountains. This event, which is responsible for the sinuous shape of the Talesh Mountains, produced a southwards displacement of the crystalline nappe stack now exposed along the Caspian foothills. This work was funded by the MEBE project (proposal 02– 26, Meso-Cenozoic Evolution of the Alborz Mountain Range, Iran) and by an Italian MURST PRIN Project (2004–2006, leader A. Zanchi) in collaboration with the Geological Survey of Iran. Dr M. R. Ghasemi, Dr A. Saidi and his colleagues of the Geological Survey of Iran are warmly thanked for continuous help and logistics. S. Ando` helped in the separation of paragonites for Ar/Ar dating. Reviews by A. I. Okay and B. Lombardo greatly improved a first version of the manuscript.

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Pennsylvanian –Early Triassic stratigraphy in the Alborz Mountains (Iran) MAURIZIO GAETANI1 *, LUCIA ANGIOLINI1, KATSUMI UENO2, ALDA NICORA1, MICHAEL H. STEPHENSON3, DARIO SCIUNNACH4, ROBERTO RETTORI5, GREGORY D. PRICE6 & JAFAR SABOURI7 1

Dipartimento di Scienze della Terra, Universita` di Milano, Via Mangiagalli 34, 20133 Milano, Italy

2

Department of Earth System Science, Faculty of Science, Fukuoka University, Fukuoka 814-0180, Japan

3

British Geological Survey, Kingsley Dunham Centre Keyworth, Nottingham NG12 5GG, UK 4

Regione Lombardia, Infrastruttura per l’Informazione Territoriale, Via Sassetti 32, 20124 Milano, Italy 5

Dipartimento di Scienze della Terra, Universita` di Perugia, Piazza Universita` 1, 06100 Perugia, Italy

6

School of Earth, Ocean and Environmental Sciences, University of Plymouth, Drake Circus, Plymouth PL4 8AA, UK 7

Geological Survey of Iran, Azhad Square, Tehran, Iran

*Corresponding author (e-mail: [email protected]) Abstract: New fieldwork was carried out in the central and eastern Alborz, addressing the sedimentary succession from the Pennsylvanian to the Early Triassic. A regional synthesis is proposed, based on sedimentary analysis and a wide collection of new palaeontological data. The Moscovian Qezelqaleh Formation, deposited in a mixed coastal marine and alluvial setting, is present in a restricted area of the eastern Alborz, transgressing on the Lower Carboniferous Mobarak and Dozdehband formations. The late Gzhelian–early Sakmarian Dorud Group is instead distributed over most of the studied area, being absent only in a narrow belt to the SE. The Dorud Group is typically tripartite, with a terrigenous unit in the lower part (Toyeh Formation), a carbonate intermediate part (Emarat and Ghosnavi formations, the former particularly rich in fusulinids), and a terrigenous upper unit (Shah Zeid Formation), which however seems to be confined to the central Alborz. A major gap in sedimentation occurred before the deposition of the overlying Ruteh Limestone, a thick package of packstone–wackestone interpreted as a carbonate ramp of Middle Permian age (Wordian– Capitanian). The Ruteh Limestone is absent in the eastern part of the range, and everywhere ends with an emersion surface, that may be karstified or covered by a lateritic soil. The Late Permian transgression was directed southwards in the central Alborz, where marine facies (Nesen Formation) are more common. Time-equivalent alluvial fans with marsh intercalations and lateritic soils (Qeshlaq Formation) are present in the east. Towards the end of the Permian most of the Alborz emerged, the marine facies being restricted to a small area on the Caspian side of the central Alborz. There, the Permo-Triassic boundary interval is somewhat similar to the Abadeh–Shahreza belt in central Iran, and contains oolites, flat microbialites and domal stromatolites, forming the base of the Elikah Formation. The P– T boundary is established on the basis of conodonts, small foraminifera and stable isotope data. The development of the lower and middle part of the Elikah Formation, still Early Triassic in age, contains vermicular bioturbated mudstone/wackestone, and anachronostic-facies-like gastropod oolites and flat pebble conglomerates. Three major factors control the sedimentary evolution. The succession is in phase with global sea-level curve in the Moscovian and from the Middle Permian upwards. It is out of phase around the Carboniferous– Permian boundary, when the Dorud Group was deposited during a global lowstand of sealevel. When the global deglaciation started in the Sakmarian, sedimentation stopped in the Alborz and the area emerged. Therefore, there is a consistent geodynamic control. From the Middle Permian upwards, passive margin conditions control the sedimentary evolution of the From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 79– 128. DOI: 10.1144/SP312.5 0305-8719/09/$15.00 # The Geological Society of London 2009.

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M. GAETANI ET AL. basin, which had its depocentre(s) to the north. Climate also had a significant role, as the Alborz drifted quickly northwards with other central Iran blocks towards the Turan active margin. It passed from a southern latitude through the aridity belt in the Middle Permian, across the equatorial humid belt in the Late Permian and reached the northern arid tropical belt in the Triassic.

A project within the Middle East Basin Evolution (MEBE) Programme was devoted to improve the understanding of the Pennsylvanian –Triassic stratigraphy in the central and eastern Alborz. The field activity was performed during three stages, in September 2003, and in April and September 2004, with fieldwork by M. Gaetani and L. Angiolini. S. Crasquin (Paris) and A. Nicora participated in part of this field activity. All the work was supported as far as the logistics are concerned by the Geological Survey of Iran. J. Sabouri, M. Walik, M. Rahmani, M. R. Majidifard and A. Jalali, geologists of the GSI, shared part of the field experience. Twenty-one stratigraphic sections have been measured for a total thickness of more than 3000 m. Some 1075 samples were collected and processed (Fig. 1). Further fieldwork, scheduled to check some doubtful points, was not performed owing to the subsequent situation. The aim of the present paper is to describe the general stratigraphic features of the time interval considered over a wide area of the central and eastern Alborz, an exercise not attempted in the last 30 years. The palaeontology will be fully developed in papers that will appear separately. Here we illustrate only a few significant fossils. Brachiopods were studied by L. Angiolini, fusulinids by K. Ueno, palynomorphs by M. Stephenson, small foraminifera by R. Rettori, conodonts by A. Nicora and sandstone petrography by D. Sciunnach.

Previous knowledge Central Alborz The Upper Palaeozoic succession of the central Alborz was first studied by Assereto (1963, 1966) and Glaus (1964, 1965). Assereto (1963) introduced the formational terms of Geirud, Mubarak (the spelling was later changed in Mobarak by Bozorgnia 1973), Dorud and Ruteh, whilst Glaus (1964) established the Nesen and Elikah formations. These names are still in use in the official geological maps of Iran. A number of PhD theses from the ETH (Eidgeno¨ssische Technische Hochschule) of Zu¨rich, Switzerland, enlarged the regional knowledge, also illustrating rocks of the interval discussed here (Lorenz 1964; Steiger 1966; Dedual 1967; Cartier 1971). The palaeontology was illustrated by Fantini Sestini (1965a –c, 1966), Fantini Sestini & Glaus

(1966) and Gaetani (1968) as far as brachiopods are concerned. Flu¨gel (1964, 1968, 1994) and Holzer (1976) studied the coral fauna; Khaler (1976) illustrated the fusulinid content. Most of these papers are based on the specimens collected by Assereto, Gaetani and Glaus. The stratigraphic scheme established through these papers was reconsidered by Bozorgnia (1973), who made refinements and corrections on the basis of the foraminifera content of these units, illustrating most of them. A number of papers subsequently addressed more restricted topics, like Stepanov et al. (1969) who discussed the PermoTriassic boundary in several localities of Iran including the Alborz. Heravi (1971) illustrated the Lower Carboniferous conodonts and demonstrated the Mississippian age of the first member of the Dorud Formation sensu Assereto (1963). Recently, Ueno et al. (1997) reported foraminiferal biostratigraphy of the Mobarak Formation from the Shahmirzad area. The micropalaeontology of the Elikah Formation was illustrated in several notes by Bro¨nnimann et al. (1972a, b) and Zaninetti et al. (1972). The P–T boundary interval along the Elikah section was analysed by the Iranian – Japanese Research Group (1981). Two PhD theses were also carried out in the 1970s. Su¨ssli (1976) illustrated the geology of the Emarat anticline and of the lower Haraz Valley, and JennyDeshusses (1983) described the foraminifera content of selected sections in the Permian. Fossils collected were illustrated by Hirsch & Su¨ssli (1973) and Elliott & Su¨ssli (1975). Nazemi (1977) and Lasemi & Mokhtarpour (1994) added information on the southern border of the Alborz, while Ghavidel-Syooki (1995) provided some details from the Dorud Formation in the Hassanakdar area.

Eastern Alborz The pioneering paper by Bozorgnia (1973) was followed by two PhD theses (Jenny 1977; Stampfli 1978) of the University of Geneva, which established the stratigraphic framework for the Eastern Alborz (Jenny & Stampfli 1978). Papers dealing with part of the succession are by Stampfli et al. (1976) on the Elikah Formation and by Jenny et al. (1978) on the Qezelqaleh Formation. Lys et al. (1978) illustrated the micropalaeontology of the Carboniferous–Permian interval, while

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Fig. 1. (A) Main structural and palaeogeographic elements of Iran. MZT, Main Zagros Thrust; black, ophiolites and related rocks. (B) Index map of the central and eastern Alborz with logs and localities studied during the fieldwork.

some palaeobotanical features were studied by Chateauneuf & Stampfli (1978), and Jenny & Jenny-Deshusses (1978) described a new giant ichnogenus in the Ruteh Limestone. Altiner et al. (1980) described a section around the Permian– Triassic boundary. The area of Jam, east to Semnan, was discussed in the PhD thesis of AlaviNaini (1972), partly adopting the stratigraphic subdivision proposed for the Upper Palaeozoic of the Alborz. No other recent research has been published and these papers remain the only information on the eastern Alborz. The general features of the Carboniferous were summarized by Vachard (1996). Some general

papers and reports of the Geological Survey, issued after 1979, are written mostly or exclusively in Persian, like Partoazar (1995) and Shahrabi (1999). The Triassic of Iran, including Alborz, was reviewed by Seyed-Emami (2003), who illustrated the main features of the Elikah Formation. Several recent papers deal with the Carboniferous – Permian stratigraphy of central Iran and its fusulinid content. They include in their discussion interesting correlations to the Alborz (Baghbani 1993, 1997; Leven & Taheri 2003; Leven & Vaziri 2004; Leven et al. 2006; Leven & Gorgij 2006; Boncheva et al. 2007; Davydov & Arefifard 2007).

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The stratigraphic succession

Base of the succession

Pennsylvanian– Triassic rocks crop out rather extensively in the central and eastern Alborz, in a number of geological structures oriented east– west in the central Alborz and SW– NE in the eastern Alborz. These structures belong to the Palaeozoic Central Range and to the Southern Palaeozoic– Mesozoic Zone, which form two separate structural units (Gansser & Huber 1962) dissected by strikeslip and thrust faults and graben within which older sedimentary rocks are preserved (Zanchi et al. 2006). The quality of the outcrops is usually superb to the south of the watershed with the Caspian Sea, where the tectonics is not particularly severe. However, the Caspian side is largely covered by bushes and woods, and is heavily affected by tectonics, making it in some cases difficult to follow the previously described sections. The bases of the sections have been located by GPS on WGS84 co-ordinates. The lithostratigraphic framework established in the literature was a useful aid for describing the succession, but we improved lithostratigraphic characterization of units and their boundaries (Fig. 2). Owing to different ways to transliterate the Persian names, we use the spelling adopted by the official geological maps of the Geological Survey of Iran. When mentioned for the first time in this paper, the original spellings adopted by Jenny & Stampfli (1978) are also reported in brackets, both for formations and for localities.

The base of the studied stratigraphic interval is usually represented by calcareous or marlstone units of Early Carboniferous age, for which the names of Mobarak or Geirud formations have been proposed (Assereto 1963). Bozorgnia (1973) rejected the use of Geirud and in the official geological maps of Alborz only the term Mobarak Formation is adopted, regardless of whether the Pennsylvanian or Lower Permian successions lie on grainstone/packstone limestones or on grey siltstone/marlstones. These different lithologies probably deserve different names within a Mobarak group, yet to be defined. In many sections, both in the central and in the eastern Alborz, the uppermost Carboniferous – Permian rocks lie unconformably on the Lower Carboniferous rocks. In a few places, however, above the calcareous or marly ‘Mobarak’ succession, additional units are present. They are here briefly discussed. Assereto (1963, 1966) described the ‘Member D’ of the Geirud Formation in the area of Abnak, assigning an earliest Permian age (Fantini Sestini 1966) to this grainstone/packstone unit. In fact, it closely resembles part of the Mobarak limestone. Bozorgnia (1973) rejected the Permian age of this unit, giving a Visean age to it. We sampled the Abnak outcrop and small foraminifera confirm its Visean age (P. Brenckle, pers. comm. 2004). Assereto (1963) also described the Member 1 of the Dorud Formation, a succession of dolomitic

Fig. 2. Stratigraphic scheme for the Carboniferous –Early Triassic interval in the Alborz Mountains. Bold lines indicate Fe-enriched surfaces and/or lateritic soils. The glacial intervals of North Gondwana basins are from Isbell et al. (2003), with ages calibrated on the Standard Scale according to Menning et al. (2006).

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limestone with shale, siltstone and fine arenite, also giving to it a Permian age. Bozorgnia (1973), Heravi (1971) and Meissami et al. (1977, 1978) contested this assignment, assigning it a latest Visean-Serpukhovian age. Bozorgnia (1973) proposed the name Dozdehband Formation for this unit. However, he did not provide a detailed description of the type section beyond the fact that it gradually displays an increasing terrigenous content in the area of Dozdehband, north of the Kandevan pass on the Caspian side, along the road to Chalus. Here we refer the lower part of the Toyeh section to the Dozdehband Formation (Fig. 3). In that particular section, an approximately 35 m-thick arenaceous succession is observed above the grey marlstone of the Mobarak Formation. This is, in turn, overlain by a conglomerate layer with a sharp erosional base. The overlying approximately 40 m-thick succession consists of mixed arenites and packstone/grainstones with abraded fusulinids probably referred to Plectostaffella jakhensis Reitlinger of early Bashkirian age (Fig. 3). We consider it more appropriate to equate the Dozdehband Formation and the Mb. 1 of the Qezelqaleh Formation of Jenny et al. (1978) because of their equivalent age; according to Lys et al. (1978) these units are only Serpukhovian and there is no evidence for the Bashkirian. Stampfli (1978) described the Bagherabad Formation in the eastern Alborz. This is a 260 m-thick, mostly shallow-water limestone with terrigenous interbeds of late Visean–early Bashkirian age. The long-needed study of the upper part of Lower Carboniferous was beyond the scope of the present project.

Qezelqaleh Formation In small areas, especially north of Damghan and south of Aliabad in the eastern Alborz, Pennsylvanian continental and shallow-marine rocks are preserved and are referred to the Qezelqaleh Formation (Fig. 4). Name. Jenny et al. (1978) and Stampfli (1978) proposed the name Gheselghaleh (now Qezelqaleh Formation) for rocks belonging to this continental and shallow-marine succession. We exclude the Mb. 1 of Jenny et al. (1978) from the Qezelqaleh Formation because of its age, correlative to the Dozdehband Formation. Lithology. This consists mostly of sandstone and fine conglomerate, with sandy bar cross-laminations, capped by flat laminations, finer arenites and shaley intercalations. Bioclastic packstone forming shelly beds contains fragmentary brachiopods, bryozoans, crinoids and fusulinids, often abraded (details in Fig. 3).

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Sandstone petrography. The sandstones consist of very coarse to fine-grained litharenites, or less commonly sublitharenites, with little detrital feldspar (that is particularly abundant in medium-grained or finer sandstones). Quartz is mostly monocrystalline, with a C/Q (composite vs total quartz) ratio averaging 10%, feldspar is represented by twinned plagioclase, high-temperature alkalifeldspar (anorthoclase, chessboard-albite) and rare microcline. Rock fragments consist of sedimentary and volcanic rock types: oxidized, silty –sandy ‘limeclasts’ sensu Zuffa (1985), interpreted as carbonate extrabasinal grains, and chert are particularly abundant in coarse-grained sandstones at the base of the formation, and are depleted up-section. Terrigenous rock fragments (fine-grained sandstone, wacke and low-grade metaquartzite) occur at definite stratigraphic intervals. Felsic, microlitic and vitric volcanic rock fragments, as well as ‘graphic’ micropegmatite grains of hypabyssal origin, are also widespread, representing most of the lithic fraction in the upper part of the formation. While felsic and vitric rock fragments are derived from the groundmass of intermediate – acidic volcanics, microlithic structures are diagnostic of intermediate –basic lava flows (Dickinson 1970). Their concurrent occurrence is consistent, in principle, with a bimodal volcanic suite, which is common in rift settings (Peccerillo et al. 2003 and references therein). Heavy minerals display a typical, ultrastable ZTR (zircon –tourmaline –rutile) suite. The pseudomatrix is widespread, while intrabasinal rip-up clasts, oolites and bioclasts (echinoids, bryozoans, brachiopods, foraminifers, gastropods and ostracods) occur at certain stratigraphic levels. Pores were filled by quartz, hematite and calcite cements. Thickness and distribution. The unit has a maximum thickness of about 150 m. It rests unconformably on the Dozdehband or Bagherabad formations (Serpukhovian–early Bashkirian), and is unconformably overlain by the Dorud Group (Fig. 3). The type section of the Qezelqaleh Formation is situated in heavily vegetated country close to a military base and thus was not accessible for sampling. However, sampling at the Toyeh section (Fig. 3) north of Damghan is possible. At the Dorud type section, the topmost part of the Mb. 1 of Assereto (1963), very tentatively may belong to the Qezelqaleh Formation. Biochronology. Fusulinids, brachiopods and palynomorphs were collected. Fusulinids. In the Qezelqaleh Formation of the Toyeh section, fusulinids are found in IR 926, IR 941 and IR 942 (Figs 3 and 5), dominated by

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Fig. 3. The Dozdehband and Qezelqaleh formations in the Toyeh section. Co-ordinates (WGS84) for the base of the section N368280 0800 , E548090 4100 .

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Fig. 4. Mid-Pennsylvanian palaeogeographical map. The outcrops in the Neka Valley are shown on the Gorgan Sheet 1:250 000 of the Geological Map of Iran. Identification of the Qezelqaleh Formation in the Abrendan section is tentative.

species belonging to the family Fusulinidae (such as Pseudostaffella and Profusulinella). IR 926 in the basal part of the formation contains several poorly preserved fusulinids, including one form identifiable as Profusulinella sp. Sample IR 941 in the upper part, which is densely packed sandy bioclastic grainstone similar to IR 926, also yields fusulinids such as Eostaffella sp., Ozawainella mosquensis Rauzer-Chernousova, Pseudostaffella (P.) paracompressa Safonova and Taitzehoella cf. prolibrovichi (Rauzer-Chernousova). IR 942 contains an essentially similar to but more diverse fauna than IR 941, comprising Eostaffella mutabilis Rauzer-Chernousova, Eostaffella grozdilovae Maslo & Vachard, Pseudonovella carbonica (Grozdilova & Lebedeva), Millerella? sp., Ozawainella mosquensis, Pseudostaffella (P.) paracompressa, Pseudostaffella (P.) minutissima Putrja, Pseudostaffella (P.) aff. confusa (Lee & Chen), Nankinella sp., Staffella pseudosphaeroidea Dutkevich, Eoschubertella obscura (Lee & Chen), Taitzehoella cf. prolibrovichi, Profusulinella aff. parva (Lee & Chen), Profusulinella convoluta (Lee & Chen) and Fusulinella? sp. (Fig. 5). Palynology. Three samples (IR 932, IR 935 and IR 943) were processed for palynology, but of these only one (IR 943) yielded palynomorphs. The

assemblage is dominated by indeterminate monosaccate pollen and bisaccate pollen, but also contains ?Kingiacolpites subcircularus Tiwari & Moiz, Protohaploxpinus limpidus (Balme & Hennelly) Balme & Playford, P. cf. limpidus, indeterminate taeniate bisaccate pollen, Striatopodocarpites spp., cf. Barakarites spp., Leiosphaeridia spp., Alisporites indarraensis Segroves, indeterminate spores, Limitisporites spp., Potonieisporites novicus Bharadwaj, Verrucosisporites spp., Striomonosaccites spp., Camptotriletes spp., Potonieisporites spp., Florinites cf. flaccidus Mene´ndez & Azcuy, Vesicaspora spp., Cyclogranisporites spp. and Indotriradites spp. Brachiopods. Brachiopods are scattered through the formation and comprise species of Dyctyoclostinae, Tolmatchoffini, Derbyiidae, Neospiriferinae, Syringothyrididae, and the genera Composita and Choristites. Their study is still in progress. Age. The Qezelqaleh fusulinid assemblage is characterized by genera belonging to the family Fusulinidae and lacks schwagerinid species. This demonstrates that the Qezelqaleh Formation is definitely older than the Dorud Group of latest Pennsylvanian–Permian age. Based on the generic and some specific composition of fusulinids including Ozawainella mosquensis, Pseudostaffella

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Fig. 5. Typical fusulinids from the Qezelqaleh Formation. All from IR 942. 1, 2. Eostaffella mutabilis Rauzer-Chernousova, 80. 3. Millerella? sp., 80. 4. Eostaffella grozdilovae Maslo and Vachard, 80. 5. Pseudonovella carbonica (Grozdilova and Lebedeva), 80. 6. Taitzehoella cf. pseudolibrovichi (Rauzer-Chernousova), 30. 7. Ozawainella mosquensis Rzuzer-Chernousova, 20. 8, 9. Pseudostaffella (P.) paracompressa Safonova, 30. 10. Nankinella sp., 20. 11, 12. Eoschubertella obscura (Lee and Chen), 30. 13. Staffella pseudosphaeroidea Dutkevich, 20. 14. Profusulinella aff. parva (Lee and Chen), 20. 15– 17. Pseudostaffella (P.) minutissima Putrja, 30. 18–20. Pseudostaffella (P.) aff. confusa (Lee and Chen), 30. 21, 22. Profusulinella convoluta (Lee and Chen), 20. 23, 24. Fusulinella? sp., 23: 20, 24: enlarged part of 23, showing four-layered spirotheca with vaguely developed diaphanotheca.

(P.) paracompressa, Taitzehoella cf. prolibrovichi, Profusulinella prisca and P. convoluta, the Qezelqaleh Formation is broadly referable to the early Moscovian. In IR 942, moreover, there are several, poorly preserved and poorly oriented specimens that have slightly larger shells than associated Profusulinella species (Fig. 5.23). They show vestiges of four-layered spirotheca consisting of a tectum, diaphanotheca, and upper and lower tectorial (epithecal) layers (Fig. 5.24), which is diagnostic of Fusulinella. These specimens are questionably placed in the genus Fusulinella in this study. If they can be correctly referable to this genus, then at least the upper part of the Qezelqaleh Formation is correlated more precisely with the late early Moscovian (Kashirian) because Fusulinella first occurs in the late Kashirian and Profusulinella

becomes extinct at the Kashirian –Podolskian (early –late Moscovian) boundary (e.g. RauzerChernousova et al. 1951; Isakova 2001). If the age suggested by fusulinids may be assigned also to the sample IR 943 lying 4.8 m above the last fusulinid sample, then the composition of Moscovian palynological assemblages is remarkably similar to those of the Pennsylvanian and Early Permian in the Alborz, suggesting few evolutionary changes through the Mississippian – Early Permian period. This is in contrast to palynological successions in Euramerica and Gondwana, which seem to have been affected by glaciation or far-field effects of glaciation, probably causing rapid change. Up to now, no evidence of Kasimovian stage has been recorded in the Alborz.

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Environment. The Qezelqaleh Formation represents a mixed terrigenous– carbonate unit, formed on a surface gently eroding the substrate. In the lower part, sandstone and microconglomerate prevail, with sandy bar cross-laminations, capped by flat laminations, finer arenite and shaley intercalations. Marine fragmentary fossils are spread, with fusulinids and corals suggesting shallow-water conditions. Upwards the environmental energy decreases, so red–grey siltstone and fine arenite prevail. We consider this unit as deposited mostly under shoreface–sandy coastal bar conditions. The organic content is mostly wood fragments, terrestrial palynomorphs and rarer marine palynomorphs. The mineralogy of sandstones points to a scenario of shallow dissection into older sedimentary successions (mostly cherty limestone and quartzose sandstone), covered by basic – acidic volcanics, that might represent a rift-related bimodal suite. Concurrence of arenaceous rock fragments, quartz grains with multiple overgrowths and microcline (selectively enriched in second-cycle sand: Blatt 1967) strongly suggests recycling of ancient sandstones that locally underwent deep diagenesis–low-grade metamorphism, while high-temperature feldspar might derive from intermediate to acidic volcanics, that are actually documented as rock fragments.

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sections to the SE. The name was proposed by Assereto (1963) to include a basal unit that later appeared to be Carboniferous rather than Permian (Bozorgnia 1973) as proposed by Assereto. Jenny & Stampfli (1978, p. 554) recognized the complex pattern of the Dorud Formation and proposed to upgrade it to group status. Typically, it is broadly tripartite: a lower mostly terrigenous unit, a mostly carbonate lithosome in the middle, in which more than one formation may be identified, and an arenitic unit at the top. This third unit may be absent especially to the east. Thickness may exceed 350 m in the northern part of the central Alborz, but usually it ranges between 150 and 200 m (Fig. 6). Broadly, depocentres were situated to the north, whilst reduced sedimentation or emergent areas were situated to the south. The Dorud Group spans the latest Carboniferous to the early part of the Early Permian. It includes the Toyeh, Emarat, Ghosnavi and Shah Zeid formations.

(1) Toyeh Formation (new name)

The Dorud Group

Name. The name ‘Formation des Gre`s de Dorud’ introduced by Jenny & Stampfli (1978) is invalid because it adopts the same geographical name as for the Group. It is here replaced by the name Toyeh Formation, with type section in the Toyeh Creek (Fig. 7).

This group is almost ubiquitous in the Alborz, only missing or being extremely reduced in a few

Thickness and distribution. The total thickness of the Toyeh Formation ranges from 22 m in the

Fig. 6. SW–NE cross-section for the Dorud Group.

88 M. GAETANI ET AL. Fig. 7. The measured logs in the Toyeh Formation. WGS84 co-ordinates of the base of the logs. Dorud N368000 1800 , E518290 0100 ; Dorud East N368000 2500 , E518290 19.08; Emarat N368120 0000 , E528210 3900 ; Toyeh N368280 13.500 , E548090 41.600 ; Qeshlaq N368530 4600 , E558210 3400 ; Ghosnavi N368550 49.300 , E558270 02.600 .

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Ghosnavi section, 63 m in the Dorud section, more than 72 m in the Emarat 1 section, to about 90 m in the Toyeh type-section (Fig. 7). Lithology. The basal layer is usually a conglomerate unconformably capping the underlying layers on an erosional surface (Fig. 8). Landward, i.e. to the south, it is polymictic, a few metres thick, with well-rounded clasts up to 15 cm in size, consisting of whitish and red quartzose sandstones, subordinate dolostones, and a few dark chert clasts. More basinward, as in the type section, the conglomerates locally bear flat pebbles and brachiopod fragments in the matrix. The basal conglomerate is overlain by sandstones in m-thick beds, subdivided by thin bedded siltstones and shales, often reddish with parallel laminations and rare asymmetric small ripples, locally with clay chips up to 4–5 cm. In some sections to the north and to the east (Emarat 1, Toyeh, Qeshlaq) hybrid calcarenites are irregularly intercalated, in 30–50 cm layers, with wavy internal bedding. The biocalcarenite contains fragmentary bryozoans, crinoids and brachiopods; in the topmost part fusulinids also occur. The microfacies consists of packstones– grainstones, in which bioclasts locally display micritic coatings, with micritic matrix to sparry cement. Subordinate quartz content is observed locally. Thin intercalations of siltstones and mudstones may also occur. Sandstone petrography. Sandstones from the Toyeh Formation are medium–very-fine-grained

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sublitharenites and subarkoses. Quartz is mostly monocrystalline; altered detrital feldspars and vitric volcanic lithics are commonly reduced to pseudomatrix (Dickinson 1970). Large amounts of pseudomatrix might lead to an overestimate of quartz content, and thus to misclassification of the sandstones as quartzarenites. Cherty lithics also occur. Heavy minerals comprise an ultrastable ZTR suite, plus some white mica, monazite, apatite, titanite and rare chromian –ferroan spinel. Limeclasts (Zuffa 1985), locally oxidized and interpreted as extrabasinal carbonate lithoclasts, are particularly abundant in the Aruh section, in beds tentatively referred to this unit. Carbonate intraclasts and bioclasts (echinoids, microproblematica) and locally abundant interstitial ferroan carbonates indicate episodic marine transgressions. Biochronology. Fusulinids and brachiopods were collected. Also, a single sample was productive for conodonts. Fusulinids. Among the studied sections of the Toyeh Formation, fusulinids have been discriminated only from the Emarat 1 section, in which they are found from four levels; IR 376, IR 377, IR 922 and IR 379. Two assemblages are recognized; one is dominated by Ruzhenzevites and Rauserites from IR 376, IR 377 and IR 922, and the other is characterized by Anderssonites from IR 379. The Ruzhenzevites–Rauserites assemblage yields Eostaffella spp., Schubertella sp., Nankinella sp., several Rauserites species such as R. tabasensis

Fig. 8. The base of the Dorud Group eroding the top of the Mobarak Group at Dorud East section.

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Leven, R. infrequentis Leven, R. elongatissimus (Rozovskaya), and R. cf. jakshoensis (Davydov), Ruzhenzevites ferganensis (Dutkevich) and Triticites? sp. (somewhat similar to T.? triangulus Rozovskaya) (Fig. 9). Of them, Ruzhenzevites ferganensis, Rauserites tabasensis and R. infrequentis are dominant. This assemblage is very similar in essential specific composition to that from the upper part of Unit 7 of the Zaladou (Zaladu) Formation (originally referred to as the Zaladu Member of the Sardar Formation) in east Iran (Leven & Taheri 2003), and is also somewhat comparable based on the occurrence of Ruzhenzevites ferganeisis to that from the upper part of the Zaladou Formation in central Iran (Leven & Gorgij 2006). These parts of the Zaladou Formation are correlated to the latest Gzhelian, which means that the relevant part of the Toyeh Formation is also referable to the latest Gzhelian. The occurrence of the conodont Streptognathodus gracilis Stauffer and Plummer in IR 922, indicating a Late Carboniferous age, supports the age assessment by fusulinids. The next younger Anderssonites assemblage, occurring some 7 m above, is composed of Eostaffella sp., Anderssonites cf. subovata (Konovalova) and Dutkevitchia? sp. (Fig. 9). Of them, A. cf. subovata is originally described from the lower Asselian of Timan (Zolotova et al. 1977), although the genus is generally characteristic in the latest Gzhelian and early Asselian (e.g. Bensh 1987; Leven & Gorgij 2006). The Anderssonites-bearing level of the Toyeh Formation is probably assignable to the very latest Gzhelian or earliest Asselian. Brachiopods. Brachiopods from the Toyeh Formation comprise Derbyiidae gen. et sp. indet., Reticulatia uralica (Chernichev), Calliprotonia sp., Costispinifera sp., which, except for Calliprotonia sp., range upward to the Emarat Formation. The genus Calliprotonia occurs in the Pennsylvanian –Lower Permian of North and South America, and Russia (Brunton et al. in Williams et al. 2000). Reticulatia uralica does not appear in beds younger than the Sakmarian in Russia, but it has been reported from the Asselian of Fergana

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(Volgin 1960), the Asselian of northeastern Thailand and the Asselian –Sakmarian of Spitsbergen (Shi & Waterhouse 1996; Perez Huerta et al. 2007). Age. Where it is possible to date it, the age of the Toyeh Formation ranges from the Gzhelian to the earliest Asselian. However, owing to the timetransgressive character of that unit its age could be only Asselian in more landward sections. Environment. The Toyeh Formation was essentially deposited in a fluvial setting, evolving from braided river (the basal conglomerate) to alluvial plain conditions. The red colour testifies to the oxidizing environment, while the absence of typical point-bar sequences should indicate prevailing overbank deposits over laterally stable, anastomosing channel facies. In Emarat and Toyeh, the alluvial plain was cyclically intermingled with shoreface and coastal bar sediments, also containing carbonate bioclasts or even fully marine carbonate intercalations, as in the Toyeh section.

(2) Emarat Formation (new name) Name. Jenny & Stampfli (1978) and Stampfli (1978) introduced several nomenclatural subdivisions for the middle carbonatic part of the Dorud Group. They used the following names: 1. Alestone Formation, for a non-oncoidal facies at the base of the carbonate lithosome in the Aliabad area; 2. ‘Calcaires de Dorud’, for the calcarenite/ oncoidal facies; 3. Ghosnavi Formation, for sheltered low-energy facies, rich in mudstone/wackestone; 4. Kuh-e-Sariambar Formation, for the dolomitized upper part overlying the ‘Calcaires de Dorud’. The name ‘Calcaires de Dorud’ is nomenclaturally invalid and is replaced with the new name Emarat. The section Emarat 1, starting near the SW corner of the military camp fence, is selected as the type section. Co-ordinates of the base N368120 0000 , E528210 3900 .

Fig. 9. Relevant fusulinids of the Dorud Group. 2, 3, 6 from the Toyeh Formation; others from the Emarat Formation. 1, 2. Ruzhenzevites ferganensis (Dutkevich), 1: IR 960 (Toyeh section), 2: IR 922 (Emarat section), both 8. 3. Rauserites infrequentis Leven, IR 376 (Emarat section), 10. 4. Rauserites tabasensis Leven, IR 960 (Toyeh section), 10. 5, 6. Anderssonites cf. subovata (Konovalova), 5: IR 574 (Jauzchal section), 6: IR 379 (Emarat section), both 10. 7. Pseudofusulina? sp., IR 596 (Jauzchal section), 10. 8. Pseudoschwagerina muongthensis (Deprat), IR 599 (Jauzchal section), 10. 9. Sphaeroschwagerina vulgaris (Scherbovich), IR 391 (Emarat section), 10. 10. Praepseudofusulina urumarensis (Scherbovich), IR 978 (Toyeh section), 10. 11. Triticites? sp., IR 377 (Emarat section), 10. 12. Praepseudofusulina kljasmica (Sjomina), IR 586 (Jauzchal section), 10. 13. Praepseudofusulina? sp. (n. sp.?), IR 979 (Toyeh section), 10. 14. Praepseudofusulina incomperta (Scherbovich), IR 390 (Emarat section), 10. 15, 16. Sphaeroschwagerina sphaerica (Scherbovich), both IR 394 (Emarat section), 8. 17. Pseudoschwagerina cf. inaequialis Kahler and Kahler, IR 598 (Jauzchal section), 10. 18. Praepseudofusulina saratovensis (Chernova), IR 691 (Ghosnavi section), 10. 19. Eoparafusulina aff. pusilla (Schellwien), IR 394 (Emarat section), 8. 20. Darvasites? eocontructus Leven and Scherbovich, IR 402 (Emarat section), 10.

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Lithology. The Emarat Formation is characterized by oncolitic cycles, and by fusulinid grainstone/ packstone, replaced laterally and/or upwards by mudstone/wackestone deposited in a low-energy, sheltered environment, which is sometimes dolomitized (Fig. 10). The oncolitic cycles prevail in the lower and middle parts, interbedded with the fusulinid grainstone/packstone. The internal sequence of the oncolitic cycle is fining upwards, with small oncolites in the lowermost cycles and larger, up to 2–3 cm upwards, in the thickest cycles. The oncolites are either self- or matrix-supported and their nucleus may consist of bioclasts including fusulinids (Fig. 11). Their shape is almost exclusively rounded with a number of coatings that may be subsequently partly abraded. The laminae may contain interwoven microbial filaments (Girvanella-like form) and envelop several generations of smaller oncolites. The pure oncolitic facies is usually quartz-free, while the pure fusulinid facies may contain a small amount of terrigenous material. The carbonates form up to five or six major cycles, 10 –15 m thick, subdivided by fine sandstone layers, a few metres thick, or simply by marly intervals. The oncolitic cycles are thinner and more numerous in the Jauzchal section. Brachiopod cocquinas and black shale intercalations are subordinate. According to Jenny & Stampfli (1978) the fusulinid facies should be nearshore and the oncolitic more reefal. We found both facies in all the measured sections, but in the more basinward sections like Emarat and Jauzchal, the oncolites are more abundant and form a thickness of rocks greater than in the landward sections, like Ghosnavi, Toyeh and Dorud. At the base a sharp contact (disconformity?) is observed at Dorud, whereas a gradual onset of bioclastic intercalations occurs at Emarat. Upwards the trend is towards grey dark packstone/wackestone deposition, still cyclic. Bedding may be thin, like in the east (Jauzchal (¼ Dzuchal of Stampfli & Jenny 1978), Qeshlaq, Ghosnavi; Fig. 10) or in thicker beds, often dolomitized (Toyeh, Kuh-e-Sariambar). This uppermost part of the Emarat Formation testifies to the trend towards more sheltered environments, in which the terrigenous input decreases or even vanishes. Sandstone petrography. The unit is mostly devoid of sandstone. We collected arenites only in the type section, where the formation displays very fine-grained, poorly sorted subarkoses with twinned plagioclase grains, carbonate and cherty/vitric? lithics. The heavy mineral suite includes zircon, tourmaline, rutile, white mica and titanite. Interstitial ferroan carbonate is widespread. Thickness and distribution. The total thickness of the Emarat Formation is between 120 m (Jauzchal

section) and 14 m in the Dorud section. To the east it consistently exceeds 100 m, to the west it is usually less than 100 m, as the formation is probably partly replaced by the lower and upper terrigenous units of the Dorud Group. Biochronology and age. Fusulinids are abundant in the Emarat Formation and occur in many levels in the studied sections. The relevant fusulinids from the Emarat Formation are illustrated in Figure 9. In the Emarat Formation, six fusulinid assemblages are recognized, which are respectively characterized by Ruzhenzevites–Rauserites, Anderssonites, Praepseudofusulina, Pseudoschwagerina– Praepseudofusulina, Sphaeroschwagerina-Eoparafusulina and ?Darvasites. The first two assemblages are the same found in the upper part of the underlying Toyeh Formation. This demonstrates that some portions of the basal part of the Emarat Formation represents a facies interfingering with the upper part of the Toyeh Formation, suggesting the diachronous onset of carbonate deposition in the Dorud Group of the central-eastern Alborz Basin. The Ruzhenzevites –Rauserites assemblage, referable to the latest Gzhelian as in the Toyeh Formation, is recognized only in the basal part of the Toyeh section (IR 957, IR 921, IR 959 and IR 960). Its specific composition is nearly the same as that in the Toyeh Formation and is dominated by Rauserites and Ruzhenzevites. The next younger assemblage of a latest Gzhelian or earliest Asselian age is characterized by Anderssonites cf. subovata in association with Eostaffella sp., Nankinella sp., and Rauserites sp. and is observed in the basal part of the Jauzchal section (IR 573 and IR 574). As depicted in Figure 10, a coeval stratigraphic interval would also exist in the lower part of the Toyeh section between IR 960 and IR 966–IR 967, although Anderssonites is not found there. The Asselian is well developed in the Emarat Formation and it is generally represented by oncoidal limestone. The first assemblage of the Asselian is not very conspicuous in taxonomic composition, consisting of Praepseudofusulina kljasmica (Sjomina), P. urmarensis (Scherbovich), P. incomperta (Scherbovich), Nankinella sp. and Schubertella sp. This assemblage, characterized by the occurrence of Praepseudofusulina and the lack of Pseudoschwagerina, is observed in IR 586 and IR 588 in the Jauzchal section and IR 679–IR 684 in the basal part of the Ghosnavi section. In addition, IR 966 and IR 967 in the Toyeh section may also correspond to this interval, although fusulinids found from these levels are rather poor both in occurrence and preservation. The genus Praepseudofusulina, established originally by Ketat & Zolotukhina (1984) with ?Pseudofusulina fastuosa Ketat as the type species, is characteristic of the lower part of the Asselian (Kireeva et al. 1971; Ketat &

PENNSYLVANIAN –EARLY TRIASSIC ALBORZ IRAN Fig. 10. The most relevant measured logs for Emarat and Ghosnavi formations. WGS84 co-ordinates for the base of the logs. Continuation of the previous logs for Dorud, Emarat 1, Toyeh sections. Base of the Jauzchal section N368530 39.600 , E55890 25.200 , Qeshlaq section, the lower part in prosecution of the previous log; the upper part along the creek starting at N368540 0100 , E558210 36.600 ; Ghosnavi section, continuation of the previous log. 93

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Fig. 11. Oncolitic and grainstone microfacies from the Emarat Formation. Scale bar, 2 mm. (A) Composite oncolite including three smaller oncolites, with external cortex partly abraded; IR 394 Emarat 1 section. (B) The oncolite laminae started to grow around a cridoid ossicle, associated with Girvanella tubules, later including also a bryozoan fragment; IR 393. Emarat 1 section. (C) Incipient oncolitic growing on a fusulinid, IR 593, Jauchal section. (D) Fine packstone/grainstone with small litho- and bioclasts, in which sparse displaced oncolites still occur. Upper part of a cycle, sample IR 402, Emarat 1 section. (E) Fusulinids free of microbial envelop may occur in the matrix of oncolites, that however may bind also fusulinids. IR 389, Emarat 1 section. (F) Packstone/grainstone with fine lithoand bioclasts, capping the oncolitic cycle. IR 35, Dorud section.

Zolotukina 1984; Rauzer-Chernousova et al. 1996). In particular, P. kljasmica is also reported from the lower Asselian of the Zaladou Formation of the Anarak section in Central Iran (Leven & Gorgij 2006). Therefore, an early Asselian age for this assemblage is concluded. The basal, some 15– 20 m-thick in parts of the Emarat 1 and Qeshlaq sections, may also be time-equivalent to those represented by this fusulinid assemblage, judging from their stratigraphic positions. The next younger assemblage is characterized by Pseudoschwagerina –Praepseudofusulina, and is recognized in IR 33 and IR 36 in the Dorud section, IR 389 –IR 393 in the Emarat 1 section, IR 969, IR 973 and IR 978 in the Toyeh section, IR 594 –IR 601 in the Jauzchal section, IR 771 –IR 775 in the Qeshlaq section, and IR 685 –IR 692 and also probably IR 693 in the Ghoshnavi section. This assemblage includes Pseudoschwagerina muongthensis (Deprat), P. cf. aequalis Kahler & Kahler, P. cf. truncata Rauzer-Chernousova, P. aff. beedei (Beede & Kniker), P. sp., Sphaeroschwagerina vulgaris (Scherbovich), S. cf. shamovi (Scherbovich), Praepseudofusulina saratovensis (Chernova), P. incomperta, P. urmarensis, P. sp., ?Pseudofusulina sp., Rugosochusenella sp., Boultonia sp., Nankinella sp., Staffella sp., Eostaffella sp. and Schubertella sp. The stratigraphic interval represented by this

assemblage records the first occurrence of the genus Pseudoschwagerina in the Emarat Formation, which is known to appear first in the middle Asselian (e.g. Leven & Scherbovich 1978). This fusulinid information warrants a middle Asselian age for this part of the formation. The next assemblage, typically observed in IR 394 of the Emarat 1 section but also probably represented by IR 53 in the Dorud east section, IR 979 in the Toyeh section, IR 776 in the Qeshlaq section and IR 698 in the Ghosnavi section, is characterized by Sphaeroschwagerina sphaerica (Scherbovich), Eoparafusulina aff. pusilla (Schellwien), Pseudofusulina cf. macra (Shamov), Rugosochusenella cf. paragregaria (Rauzer-Chernousova), Biwaella sp., Eostaffella sp., Schubertella spp., ?Praepseudofusulina sp. (n. sp.?) and Nankinella sp. Of these fusulinids, Sphaeroschwagerina sphaerica is well known as an indicator of the late Asselian –early Sakmarian. Judging from the evidence that the underlying Pseudoschwagerina –Praepseudofusulina assemblage indicates the middle Asselian, the Sphaeroschwagerina –Eoparafusulina assemblage is referable to the late Asselian. In the Emarat 1 section, there is a fusulinid assemblage definitely younger than the late Asselian Sphaeroschwagerina sphaerica assemblage in

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IR 402 and IR 404. It is characterized by ?Darvasites species such as ?D. eocontructa Leven and Scherbovich and ?D. sp., Eostaffella sp., Schubertella sp. and a small unidentifiable schubertellid? (somewhat similar to Grovesella recently established by Davydov & Arefifard 2007). Among the identified fusulinids, ?D. eocontructus is originally reported by Leven & Scherbovich (1980) from the Sakmarian of Darvaz. Although its age assignment is still somewhat controversial, an early Sakmarian age is probably suggested for this youngest assemblage in the Emarat Formation in view of its stratigraphic relationship with the underlying late Asselian assemblage. The brachiopod fauna of the Emarat Formation consists of Derbyiidae gen. et sp. indet., Neochonetes (Neochonetes) sp., Costispinifera sp., Reticulatia uralica (Chernychev), Juresania dorudensis Fantini Sestini, Linoproductus dorotheevi (Fredericks), Cancrinella cancriniformis (Chernychev), Linoproductidae gen. et sp. indet., Acosarina aff. A. juresanensis (Chernychev), Composita sp., Spiriferella sp., Neospirifer sp., Larispirifer fantinisestinii Angiolini & Stephenson and Dielasma sp. Reticulatia uralica seems to be restricted to the basal (early Asselian) part of the formation, L. dorotheevi only occurs in the upper Asselian beds, whereas Neochonetes (N.) sp. and A. aff. A. juresanensis appear at the top of the formation (Sakmarian). However, the ranges of these species are known to be greater outside the region and generally regarded as Asselian –Sakmarian (Chernyshev 1902; Shi & Waterhouse 1996). The brachiopod fauna shows strong affinities with the Asselian– lower Sakmarian faunas of the Urals and of the Russian Platform (Chernyshev 1902; Licharev 1966; Kalashnikov 1986, 1988), with the Sakmarian fauna of the Trogkofel Limestone of the Carnic Alps (Schellwien 1900), with the Asselian –Sakmarian brachiopods of the Yukon Territory, Canada (Shi & Waterhouse 1996) and with the Asselian fauna of northeastern Thailand (Perez Huerta et al. 2007). In general, the Dorud fauna shows a very low endemicity, consisting of widespread boreal and palaeoequatorial genera and species. Very characteristic is the absence of typical Gondwanan elements (Angiolini & Stephenson 2008). The palynomorph assemblage, which is dominated by monosaccate pollen, with very few spores, is most unlike those recorded from the Asselian –Sakmarian Granulatisporites confluens Biozone which is ubiquitous in the Gondwana region. Samples IR 52, IR 53 and IR 830 yielded similar poorly preserved and sparse palynomorphs that are considered as a single assemblage. The assemblage is dominated by monosaccate pollen (mainly Potonieisporites spp.), with very few

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spores and acritarchs. Pollen taxa recorded include Alisporites indarraensis Segroves, ?Barakarites sp., Cannanoropollis bilateralis (Tiwari) Lindstro¨m, ?Complexisporites sp., Corisaccites alutas Venkatachala & Kar, Hamiapollenites fusiformis Marques-Toigo, ?Kingiacolpites subcircularis Tiwari & Moiz, Limitisporites sp., Plicatipollenites malabarensis (Potonie´ & Sah) Foster, Potonieisporites cf. brasiliensis (Nahuys, Alpern & Ybert) Archangelsky & Gamerro, Potonieisporites novicus Bharadwaj, Protohaploxypinus amplus (Balme & Hennelly) Hart, Protohaploxypinus limpidus (Balme & Hennelly) Balme & Playford, Striasulcites tectus Venkatachala & Kar, Striatopodocarpites spp., Sulcatisporites ovatus (Balme & Hennelly) Bharadwaj, Vesicaspora spp. and Vittatina costabilis Wilson. Spore taxa include Apiculiretusispora sp. and Calamospora sp.; acritarch taxa include ?Diexallophasis sp. and Veryhachium spp. Amongst these taxa, perhaps the most distinctive are ?Barakarites sp., ?K. subcircularis and C. alutas, and to a lesser extent Vesicaspora spp. and Striasulcites tectus, which are characteristic of the upper parts of the Lower Gharif Member and Middle Gharif Member in Oman and of the OSPZ3 or OSPZ4 biozones of Stephenson et al. (2003; late Sakmarian–?Kungurian). Thus the assemblage is unlike those typical of the Asselian–early Sakmarian of geographically close areas of the Arabian peninsula and other Gondwana areas in Australia, India, South America and Antarctica, which are characterized by the Granulatisporites confluens Oppel Zone of Foster & Waterhouse (1988). Assemblages of the G. confluens Oppel Zone of Oman and Saudi Arabia (approximately equivalent to the OSPZ2 Biozone of Stephenson et al. 2003) are typically dominated by fern spores of the genera Microbaculispora and Horriditriletes, as well as by colpate pollen such as Cycadopites cymbatus (Balme & Hennelly) Segroves. A few monosaccate pollen taxa such as Plicatipollenites malabarensis and Cannanoropollis spp. are common to the Dorud and Arabian Asselian– Sakmarian assemblages, but overall the quantitative character of the two groups is very different. Age. On the basis of the previous data the Emarat Formation starts in the Gzhelian in the sections situated towards the marine embayment in the eastern Alborz, where the Pennsylvanian rocks are better preserved. Most of the unit is Asselian in age, with the lowermost part of the Asselian not often well documented, whilst the middle Asselian is best represented in every section. Evidence for late Asselian and locally also for earliest Sakmarian is also present. Environment. The onset of prevailing carbonate deposition was always under shallow-water conditions, in a shoreface to carbonate ramp setting,

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episodically flooded by periods of quartz sand deposition that stopped the growth of carbonate sheets. The energy of the environment was higher in the lower part, with many bioclasts, typically coated by microbes, forming oncolites, in fining upwards cycles. The upper calcareous intercalations are mostly dark mudstone/wackestone, suggesting a more sheltered, restricted environment, and minor oxidation of the organic matter.

(3) Ghosnavi Formation Name. Proposed by Jenny & Stampfli (1978) with type section on the cliff to the west of the Ghosnavi road. Lithology. Fairly monotonous succession of thinbedded grey– pale grey mudstone/wackestone. In the Qeshlaq section bedding is locally amalgamated to form m-thick beds. Microfacies varies from mudstone to wackestone and, less commonly, to wackestone/packstone. Exceptional is the presence of fine packstone/grainstone, rarely with fusulinids. Bioclasts consists of thin-shelled bivalves, crinoids, dasycladacean algae, rare Cayeuxia, Tubiphytes and undetermined small tubules. Small foraminifers are very rare. Thickness and distribution. This formation is about 60 m in the type section (but the top is cut by a fault) and at Qeshlaq. It is possibly greater at Jauzchal, but tectonic alteration prevents accurate measurement. Biochronology. Age-diagnostic fossils are scarce. In a packstone intercalation in the Jauzchal section a small, possibly late Asselian or Sakmarian, fusulinid assemblage has been detected (sample IR 609, Fig. 10). Being lateral equivalent of the upper part of the Emarat Formation, the Ghoshnavi Formation is likely to be late Asselian –Sakmarian in age. Environment. The Ghosnavi Formation represents a sheltered, shallow-water part of the carbonate shelf, that characterizes the Lower Permian in the Alborz. Fine bioclastic debris were occasionally derived from the nearby higher energy areas where the Emarat Formation was deposited. Small cyanophycean algae locally built small mounds. Terrigenous input of clay was definitely subordinate. The problem of Alestone and Kuh-e-Sariambar formations. Jenny & Stampfli (1978) introduced the name of Alestone (better spelling Aleston) Formation for a succession of some 65 m thick lying below the oncoidal deposits in the basal part of the Dorud Group. They supposed this unit to be the lateral equivalent of the first unit of the Dorud Group (their Gre`s de Dorud ¼ our Toyeh

Formation). We were unable to resample the type section of the Aleston Formation. However, we do not follow this subdivision because: † the ‘calcaires de Dorud’ (¼ our Emarat Formation) often do not start with the oncoidal facies, as already noted by Jenny & Stampfli (1978); † we dispute the presence of interfingering of the Aleston Formation and Gre`s de Dorud as interpreted by Jenny & Stampfli (1978). The age of the Toyeh Formation is probably largely Gzhelian according to the fusulinids in the Toyeh and Emarat sections. The localities of Aleston and Qezelqaleh lie basinwards in comparison to Toyeh. Therefore, because the aggradation should start from the north, the carbonate facies with fusulinids should be found in the Gzhelian. Instead, the foraminiferal fauna of the Aleston type section (Jenny 1977) suggests an earliest Permian age. We suggest that the Aleston Formation is simply a local extension of the lower part of the Emarat Formation, prior to the spreading of the oncolitic facies. In the upper reaches of the Fazelabad Valley, Jenny & Stampfli (1978) noted the upwards evolution of the carbonate lithosome towards cliffforming dolomitized layers, and proposed the name of Kuh-e-Sariambar Formation. We were not able to restudy that particular section. This appears to be a rather local facies, not found elsewhere in the Alborz Mountains. However, a similar trend towards dolostones was recently noted also in central Iran, in the Anarak region (Leven & Gorgij 2006).

(4) Shah Zeid Formation (new name) Name. It corresponds to lithozone 4 of the Dorud Formation of Assereto (1963). Because it is absent in most of the eastern Alborz, Jenny & Stampfli (1978) did not separate the few arenitic layers occasionally cropping out at the top of the carbonates of the Dorud Group, grouping them into the Qeshlaq Formation. The type section was measured in the Emarat anticline, the lower part to the south of the road to the village of Shah Zeid and its shrine, and the topmost part along that road itself (Fig. 12). Co-ordinates of the base N368120 1700 , E528210 3400 . Lithology. The unit is basically siliciclastic. Its onset is rather gradual, with fine-grained sandstone, siltstone and shale that stop the growth of the carbonate succession. The sediments are often reddish, but also whithish and varicoloured. Typical is the presence of whitish quartzarenite in m-thick packages with low-angle or hummocky cross-lamination. In

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Fig. 12. Measured logs in the Shah Zeid Formation. Dorud, prosecution of the previous log. Emarat 2 section N368120 13.3, E528210 34.2; road to Shah Zeid section N368120 3000 , E528210 7800 ; Emarat quarry section N368120 4700 , E528210 33.3.

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the Emarat area, the topmost white quartzarenites include dark-grey mudstones, 10 m-thick, containing a small brachiopod fauna. Bedding is thin– medium. In the southernmost outcrops (Ruteh –Hassanakdar) of the Jajrud and Karaj valleys, the top of the formation consists of whitish–pink quartzarenites in dm-thick layers, capped by an Fe-enriched sandstone, up to 7 m thick. Sandstone petrography. Sandstones from the Shah Zeid Formation are medium- to very-fine-grained sublitharenites and quartzarenites; only in the lower part of the unit, in the Emarat section, very fine-grained subarkoses occur. Quartzarenites are commonly coarser grained and well sorted, while sublitharenites and subarkoses are only moderately sorted: this suggests some influence of mechanical processes in the sedimentary environment on higher mineralogical stability of sand. Quartz is mostly monocrystalline (the C/Q ratio never exceeds 10%), while feldspar is represented by twinned plagioclase and lithic grains by chert and vitric volcanics, or less commonly by carbonates and phyllites. The heavy mineral suite includes zircon, tourmaline, rutile, white mica, monazite, titanite, allanite and Cr-spinel. Some pseudomatrix and locally abundant syntaxial quartz cement has reduced the primary porosity; interstitial ferroan carbonates are also widespread. Feldspar enrichment in the finer-grained, more distal Emarat section might be explained by grain size effects (Odom et al. 1976), by hydraulic sorting of feldspar in a lower shoreface environment (Ibbeken & Schleyer 1991) or even by less pronounced diagenetic leaching of feldspar in finergrained, less permeable sandstones (E. Garzanti pers. comm.). The basal subarkoses of the Emarat section are arranged in a shallowing-upwards sequence, 70 m thick, capped by well-sorted quartzarenites with abundant syntaxial cement (‘orthoquartzites’ sensu Folk 1974) and sparse iron oxides; the latter suggest deposition in transitional environments fed with detritus from a fairly mature alluvial system. Thickness and distribution. This unit is identified only in the central Alborz; however, it is missing to the south (Bibi Shahrbanu, Mobarakabad). Part of the terrigenous succession at Shahmirzad might be referred to this unit. The thickness is 97.5 m at Dorud and exceeds 160 m at Emarat. Biochronology. The brachiopod faunule collected in the upper part of the unit at Emarat, comprises the species Acosarina aff. juresanensis (Chernyshev)

and Cancrinella cancriniformis (Chernyshev). At Dorud only Neochonetes (Neochonetes) sp. is present in the Shah Zeid Formation. These taxa broadly indicate an Asselian– Sakmarian age. Environment. The package of reddish shales, siltstones and rarer sandstones indicates an emersion under arid conditions and laminar sheet deposition under low energy. Upwards, well-sorted quartzarenites seem to suggest an episodic return to shoreface settings. Towards the top and only in the Emarat area, following an emersion episode, there was a temporary shallow-marine transgression, with partly restricted mudstone/wackestone, containing a brachiopod faunule. The formation ends with shoreface quartzarenites. The ironstone mapped by Assereto (1966) in the southernmost belt of the Jajrud and Karaj valleys suggests diagenetic iron enrichments linked to the accumulation of weathered material, but is not a true laterite. Correlation. To the east, this unit is absent or it is represented by a few metres of quartzarenites. Stampfli (1978) referred these beds to the base of his Qeshlaq Formation. However, the Qeshlaq Formation, as it was originally proposed, seems to be a kind of catch-all to which all the terrigenous sediments post-Dorud and pre-Elikah are referred, spanning the time from the top of the Lower Permian to the lowest Triassic.

Provenance of the Dorud Group sandstones The occurrence of phyllites, white mica and of a relatively rich heavy mineral assemblage suggests that the Toyeh, Emarat and Shah Zeid formations were at least partly fed by a continental crystalline basement, with possible contribution from the recycling of sedimentary successions. Erosion of volcanic successions is also indicated: in particular, detrital chromian spinel occurs in Visean – Guadalupian sediments from the eastern Tethys, where it is interpreted to record the basic endmembers of a bimodal magmatic suite associated to the Neotethyan rifting (Sciunnach & Garzanti 1997). In general, a feeble but definite contribution from an uplifted continental basement, compared to the underlying Qezelqaleh Formation sandstones, might indicate a stage of rift-related tectonic uplift in faraway regions of the drainage basin.

Ruteh Limestone Name. Described as the Ruteh Limestone by Assereto (1963), with type section to the north of

Fig. 13. Measured logs in the Ruteh Limestone. WGS84 co-ordinates for the base of the logs. Ruteh, type section of the formation, N358580 38.300 , E518320 28.900 ; Mangol N368140 58.200 , E528220 05.100 ; Aruh N358390 5100 , E528230 56.400 ; Talar Rud N358550 13.200 , E538060 08.700 ; Shirinabad N368530 4600 , E558090 3000 ; Gheselghaleh N368450 08.600 , E548480 57.700 , the second part starts at N368450 12.700 , E548480 55.9.

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Ruteh (Jaj Valley). With reference to the original definition, the upper boundary is now drawn at the base of the lateritic horizon (Assereto’s upper ironstone: Assereto 1963, p. 537, fig. 11). The Assereto’s level 6 should be split and attributed to the Ruteh, Qeshlaq, Nesen and Elikah formations, respectively (Fig. 13). Lithology. The base (10– 25 m thick) usually consists of dark grey marly bioclastic packstone in thin beds, rich in bryozoans and brachiopods. It is overlain by a rather monotonous succession of grey well-bedded bioclastic packstone and wackestone. Bioclasts are mostly dasycladacean algae, crinoids, brachiopods, bivalves and ostracods. Marly and shaly intercalations may locally occur in the middle and upper part, especially on the Caspian side (Shirinabad section). Chert may replace algae and coral fragments or, especially towards the Caspian, may forms blackish nodules up to 10– 15 cm, mimicking a character of the upper part of the overlying Nesen Formation. Jenny (1978) and Jenny & Stampfli (1978) erroneously attributed the upper part of the Ruteh Limestone in the Qezelqaleh section to the Nesen Formation, because of this character. Widespread are large Zoophycos-like burrows. Su¨ssli (1976) described a 20 m intercalation of fine arenites in the Emarat section. In the upper part beds may amalgamate to form cliffs, some 10 m high. Where observed, the top is invariably characterized by an emersion surface, with karstic features (Mangol) and capped by a lateritic horizon (Ruteh, Bear Gully, Aruh, Shirinabad, Qezelqaleh). Only in the area to the north of the Elikah– Nesen belt, basaltic lava flows and tuffaceous layers, up to 150 m thick but quickly thinning out eastwards, are present. The spilitized lava flows have a medium-grained plagioclase lathwork, partially replaced by authigenic calcite. They lie within the upper part of the Ruteh Limestone and not at its top as stated by Glaus (1965). Palaeomagnetic data from this basaltic flow were obtained by Besse et al. (1998). Thickness and distribution. The Ruteh Limestone is spread over most of the area, being absent only in isolated spots to the south (Mobarakabad, Shahmirzad) and to the east (Abrendan, Qeshlaq, Ghosnavi) (Figs 14 and 15). The thickness is rather homogeneous, between 150 and 250 m (Fig. 13). Towards the Caspian it may reach greater thickness (600 m in the Emarat anticline according to Su¨ssli 1976, or more than 300 m in the Talar Rud section). Biochronology Brachiopods. Usually only the basal part is rich in macrofossils, with brachiopods and bryozoans

(Fantini Sestini 1965a) being rarer upwards. Other brachiopod assemblages in the middle– upper part of the Ruteh Limestone have been collected in the Shirinabad section in the eastern Alborz. The brachiopod assemblage from the base of the Ruteh Limestone in the Dorud, Ruteh and Kuban Pass sections comprises species of Neochonetes (Nongtaia), Marginifera, Spinomarginifera, Kotlaia, Derbya, Meekella, Orthothetina, Cleiothyridina, Squamularia, Spiriferellina and Dielasma, an association of wide-ranging and palaeoequatorial genera. Brachiopods from higher levels have been collected along the Shirinabad section, where two distinct assemblages occur. The lower one is from the middle part of the Ruteh Limestone, and it is characterized by species of the genera Ogbinia, Haydenella, Marginifera, Tyloplecta, Linoproductus, Magniplicatina, Vediproductus, Chonostegoides Urushtenoidea, Orthothetina, Perigeyerella, Schuchertella and Dielasma, which are mainly palaeoequatorial taxa. The second assemblage comes from the upper part of the formation and besides Vediproductus, ranging up from the lower assemblage, comprises the genera Ombonia, Entacanthadus, Notothyris and Labaia. Notwithstanding the obvious faunal change, it is difficult to give a more precise age other than Wordian –Capitanian to the three brachiopod assemblages. A Roadian age for the base of the formation may be excluded owing to the occurrence of Spinomarginifera. In the middle part of the formation the first occurrence of progenitor species of typical Lopingian genera such as Haydenella and Tyloplecta, together with typical Guadalupian taxa (Vediproductus, Chonostegoides, Urushtenoidea) may indicate a Capitanian age. Foraminifera. Smaller foraminifera are fairly common, as well as occasionally fusulinids, even if they never form dense accumulations as in the Emarat Formation. Several assemblages may be recognized within the smaller foraminifera. A first assemblage is characterized by the dominance of Biseriamminidae mainly represented by the genera Globivalvulina and Paraglobivalvulina associated with small fusulinids, bryozoans, dasycladacean algae – mostly Permocalculus – ostracods and small gastropods. A second assemblage is characterized by dominating Palaeotextulariidae, such as Climacammina, Cribrogenerina and Palaeotextularia, associated with small Hemigordiopsidae. Lagenidae are present mostly in the upper part of the section, as well as dasycladacean algae with Permocalculus and fairly spread Mizzia (Fig. 16). In contrast to smaller foraminifers, fusulinids occur only sporadically in the studied sections (Fig. 16). In the Talar Rud section, the upper part

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Fig. 14. Stratigraphic scheme for the Middle and the Upper Permian of the Central and Eastern Alborz along NW– SE and SW– NE transects. The asterisks show the position of the volcanics.

(IR 1199–IR 1210) bears Pseudofusulina padangensis (Lange) together with Yangchienia haydeni Thompson in its uppermost part. According to Kotlyar et al. (1989) and Leven (1998), these species characterize the upper Gnishik and lower Arpa horizons of the Middle Permian in Transcaucasia, their biostratigraphy being essentially applicable to the Permian of the Alborz Mountains. The relevant Transcaucasian horizons can be correlated to the upper Murgabian (¼ Wordian) and lower Midian (¼ Late Wordian –Capitanian) in the Tethyan standard stratigraphic subdivision

(Kotlyar et al. 1989; Leven 1998). A similar fusulinid assemblage dominated by Pseudofusulina padangensis is also found in the uppermost part of the Qezelqaleh section (IR 556), giving a late Murgabian–early Midian age (Wordian–Capitanian) for this level. In other sections for the Ruteh Limestone, fusulinids are extremely poor and are only represented by small forms, such as Schubertella, Minojapanella and Codonofusiella, and/or rather long-ranging genera like as Nankinella and Staffella. In the Ruteh type-section, IR 68, IR 69 and IR 72 yield

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Fig. 15. Palaeogeographic scheme for the Middle Permian in the Alborz . Localities recorded in the literature are also included. The asterisks show the position of the basalt flows in the middle–upper part of the Ruteh Limestone.

Minojapanella (M.) sp., Schubertella sp. A and Schubertella sp. B. In IR 104 of the uppermost part of the section, Codonofusiella sp., together with the smaller foraminifer Dagmarita sp., is found. The age assessment of these levels is not very easy, but they are probably younger than the interval represented by the Pseudofusulina padangensis assemblage in the upper part of the Talar Rud section, and are potentially referable to some parts of the Khachik Horizon (Midian ¼ Capitanian) but according to Leven (1998) the uppermost Chanakhchy Beds of the Khachik Horizon are early Wuchiapingian. Similar Schubertella–Minojapanella assemblage is also found in IR 630 of the Shirinabad section. In the Mangol and Aruh sections, only Nankinella, fragments of schwagerinids and small unidentified forms (probably boultoniids) are detected, so that their exact age assignment is difficult, although it could be probable that they are broadly referable to the upper part of the Middle Permian. In the middle and upper part of the unit, colonial rugose corals, mostly in life position, are patchy in distribution. Flu¨gel (1964) identified some 20 species, in an assemblage characterized by Asiatic forms with the genera Ipciphyllum and Polythecalis (Schroeder pers. comm. 2007). Age. The age of the Ruteh Limestone spans the Middle Permian, with both Wordian and Capitanian being represented, defined on the base of foraminifera

and brachiopods. However, we did not find any conclusive evidence for a Late Permian age for the upper part of the Ruteh Limestone. Consequently, the gap at the base of the Ruteh Limestone roughly spans from the late Sakmarian to the Roadian, for a time of about 15 –20 Ma (Gradstein et al. 2004; Menning et al. 2006). Environment. We interpret the Ruteh Limestone as deposited on a wide carbonate ramp near the fairweather wave-base or just below. This environment was stable through most of the unit. The very diffuse bioclastic packstone microfacies suggests repeated input from nearby higher energy environments. The presence of sparse colonial rugose corals in life position in the middle and upper parts of the unit demonstrates that these parts never became a very deep environment, even if the small increasing of the Lagenidae in the upper part should suggest a more open environment.

Nesen Formation Name. Introduced by Glaus (1964), was later amended by Stepanov et al. (1969) and Bozorgnia (1973), who excluded the first three levels of Glaus (1964) from the Nesen Formation. We follow this interpretation, and our Bear Gully section (Fig. 17) should be very proximal to the original section measured by Glaus. The positioning of the original type section is equivocal, because the co-ordinate

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Fig. 16. Continued.

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reference given by Glaus (1964) does not coincide with the described type locality.

1994) (Fig. 18). The Nesen Formation is laterally replaced to the east by the Qeshlaq Formation.

Lithology. Two members are identified in the Nesen Formation (Fig. 17). The lower member (up to 70 m thick) consists of dark grey, splintery shale interbedded with nodular, thin- to mediumbedded marly limestone. At their base rare, very thin volcanic flows (spilites with remnants of lathwork structures and altered plagioclase phenocrysts) and siltstone are interbedded to the shale. Upwards, the calcareous fraction gradually prevails. Its microfacies consists of mudstone or wackestone with some intercalation of bioclastic packstone, with brachiopods, dasycladacean algae (Permocalculus sp.) and small foraminifera. The upper member records a decrease in shaley intercalations, and a concurrent increase of the micritic limestone (wackestone up to mudstone), in 15–30 cm-thick beds. The dominating microfacies consists of bioclastic wackestone with fine dasycladacean algae fragments and rare small foraminifera, especially Lagenidae. Matrix is often recrystallized. Marly intercalations are still present in the lower– middle part of this member, to eventually disappear in the topmost part. Silica gradually increases with sparse sprinklings, growing upwards to chert nodules, forming locally (Gallobe) complete silicified beds up to 20 cm thick. Chert decreases in the topmost layers (Ruteh, Elikah and Bear Gully sections).

Biochronology. The Nesen Formation is rich in brachiopods and palynomorphs (Figs 17, 19 –21). However, foraminifers and conodonts are rare, as well as echinoids, bivalves, ostracods and algae. Brachiopods, collected bed by bed, comprise 41 species belonging to 27 genera (Haydenella, Ogbinia, Spinomarginifera, Tyloplecta, Araxilevis, Dictyoclostinae gen. ind., Vediproductus, Anidanthus, Tschernychewia, Dorashamia, Leptodus, Orthothetina, Paraorthotetina, Perigeyerella, Ombonia, Schuchertella, Enteletes, Uncinunellina, Janiceps, Transcaucasathyris, Juxathyris, Araxathyris, Squamularia, Permophricodothyris, Dielasma, Whitspakia and Gefonia). Of these, 37 species have been analysed by Angiolini & Carabelli (2005) using the Unitary Associations method of Guex (1991) and Savary & Guex (1991), a deterministic mathematical model to construct biozones. A single run has been processed with the PAST Program (Hammer et al. 2001) based on a data set containing 37 species from five sections (Elikah, Bear Gully, Mangol Quarry, Mangol Restaurant and Abrendan) and yields six Unitary Associations (UAs), with good superpositional control and lateral reproducibility and a minimum number of contradictions, all related to low geofrequency and to indeterminate relationships between the taxa (Figs 17 and 19). Endemic taxa were not used in the analysis. Based on the quantitative biostratigraphic analysis of the brachiopod assemblages, four concurrent range biozones have been detected, which are generally separated by barren intervals.

Thickness and distribution. In the type area, the Nesen Formation reaches a thickness of 130 m. The most complete successions are preserved to the NW, towards the Caspian. In a north–south crosssection in the central Alborz, the Nesen Formation thins out rapidly (Fig. 14). In the southern part of the range the succession is usually less than 20 m thick, and is represented only by the upper cherty limestone (Ruteh, Mobarakabad). To the east in the Abrendan section we measured the Nesen Formation to be nearly 100 m in thickness but in Shirinabad, it becomes only few metres. Jenny & Stampfli (1978) referred to the formation dark limestones, tens of metres thick, cropping out between Viru and Jauzchal in the Ramian Valley. Southwards, the Nesen Formation is missing in the Aruh, Talar Rud and Shahmirzad sections, as well as at Gaduk (Steiger 1966) and Bibi Shahrbanu (Lasemi & Mokhtarpour

(1) Tyloplecta persica Biozone. Named from the most characteristic species, it corresponds to UA1 and it is characterized by the co-occurrence of three species of the genus Tyloplecta. This biozone has been strictly defined in the Mangol Quarry section and in the Abrendan section. (2) Araxilevis intermedius Biozone. Named from the most characteristic species, it corresponds to the union of UA2 and UA3, and it is characterized by the co-occurrence of Araxilevis intermedius and Spinomarginifera spinosocostata, as well as by

Fig. 16. (Continued) Microfacies and most relevant microfossils of the Ruteh Limestone. 1–3. Pseudofusulina padangensis (Lange), 1, 3: IR 1209, 2: IR 1200, all Talar Rud section, 10. 4. Yangchienia haydeni Thompson, IR 1210, Talar Rud section, 20. 5. Nankinella sp., IR 1205, Talar Rud section, 12. 6. Codonofusiella sp., IR 104, Ruteh section, 40. 7. Schubertella sp. A., IR 72, Ruteh section, 40. 8. Minojapanella (M.) sp., IR 69, Ruteh section, 40. 9, 10. Schubertella sp. B, IR 69, Ruteh section, 40. 11. Algal packstone with Permocalculus sp. and Gymnocodium sp., sample IR 74, Ruteh section, 9. 12. Climacammina sp., IR 99, Ruteh section, 25. 13. Nankinella sp., IR 102; Ruteh section,  10. 14. Permocalculus sp. and Mizzia sp., IR 94, Ruteh section; 25. 15. Mizzia sp., IR 86, Ruteh section, 20. 16. Cribrogenerina sp., IR 86, Ruteh section, 25.

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Fig. 17. Measured logs and correlation chart of five logs of the Nesen Formation (Elikah section N368130 2700 , E518200 4800 ; Bear Gully section N368150 3000 , E518250 5300 ; Mangol Quarry section N368140 5700 , E528210 4300 , Mangol Restaurant section N368150 0500 , E528220 1000 ; Abrendan section N368210 2400 , E548150 5700 ) showing the distribution of brachiopod biozones 1– 5 obtained with the Unitary Association method (see Fig. 19), the first occurrence of Triassic foraminifera, and the few conodont occurrences.

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Fig. 18. Upper Permian palaeogeography of the Alborz Mountains. The dashed line separates to the north the sections where the Nesen Formation is complete with both members. Asterisks show where feeble volcanic activity has been detected.

Fig. 19. Results of the Unitary Association method applied to five logs of the Nesen Formation drawn in Figure 17 and to 37 brachiopod taxa, showing the composition of the Unitary Associations (UA), the reproducibility matrix and the biostratigraphic graph (output of Past software, Hammer et al. 2001). N indicates the number of species; D the diversity of each Unitary Association (UA).

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the appearance of five other species of the genus Spinomarginifera. This biozone has a significant stratigraphic range and it is strictly defined in the Bear Gully, Mangol Quarry and Abrendan sections (Fig. 17). (3) Permophricodothyris ovata Biozone. Named from the most characteristic species, it corresponds to the union of UA4 and UA5, and it comprises Permophricodothyris ovata and P. iranica. This biozone records the maximum biodiversity and rate of faunal turnover. This biozone has a lateral reproducibility of 4, occurring in the Elikah, Bear Gully, Mangol Quarry and Mangol Restaurant sections. (4) Enteletes lateroplicatus Biozone. Named from the most characteristic species, it corresponds to UA6, and it is characterized by Leptodus nobilis and Enteletes lateroplicatus. This biozone is strictly defined in the Elikah section, Bear Gully section and Mangol Restaurant section. The top of this biozone records the maximum disappearance rate (100%) for the brachiopod faunas of the Nesen. Above this, only few scattered specimens of Ombonia sp. have been found in the Bear Gully section. Comparing this biozonation with the brachiopod biozones previously identified at Dorasham (Ruzhentsev & Sarycheva 1965) and Kuh-e-Ali Bashi (Stepanov et al. 1969), the Tyloplecta persica Biozone of Central Alborz may be slightly older, while the Araxilevis intermedius Biozone fits with the Araxilevis Biozone of Transcaucasia. The Permophricodothyris ovata Biozone does not easily correlate to the Oldhamina Biozone of Ruzhentsev & Sarycheva (1965) and may be younger, whereas the Enteletes lateroplicatus Biozone correlates with the Haydenella Biozone and ‘Comelicania’ (¼ Gruntallina) beds of the Djulfa area. As in Djulfa, Araxilevis appears before Haydenella, and Permophricodothyris is restricted to the lower biozones, but the appearance of species of Spinomarginifera, Leptodus, Araxathyris and Transcaucasothyris is slightly delayed in north Iran while Tyloplecta appears earlier. Palynology. Productive samples were obtained at the Bear Gully and Mangol Quarry sections. Most of the productive samples originate from the lower part of the formation (lowest two brachiopod biozones), but they display a very different preservation (the Mangol Quarry assemblages are quite well preserved) and there seems to be slight differences in taxonomic composition. From samples IR 861 and IR 864 in the Bear Gully section, the following taxa were recorded: indeterminate bisaccate and monosaccate pollen

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and spores, Alisporites indarraensis, Indotriradites spp., common Densoisporites/Lundbladispora spp., Vesicaspora spp., Distriatites insolitus Bharadwaj and Salujah, and possible fragments of Deusilites sp. and of the distinctive phytoclast ‘Maculatasporites lignin type’ (of commercial reports). Correlative beds from the Mangol Quarry section (IR 312, IR 315, IR 320 and IR 330) yield, besides Distriatites insolitus, Triquitrites proratus Balme, and ?Florinites balmei Stephenson and Filatoff (Fig. 20). The presence of very common Indotriradites/ Lundbladispora spp. suggests similarities with the highest part of the Changshingian Chhidru Formation of the Salt Range at Wargal (the ‘White Sandstone Unit’), which contains common spores of a similar type (Balme 1970). Also, Triquitrites has been recorded from the upper part of the Chhidru Formation in the Salt Range of Pakistan (Balme 1970), and the upper Baralaba Coal Measures, Narabeen Group and Rewan Formation of Australia (Helby 1973; Foster 1979). Foraminifera. Foraminifera are less abundant in comparison to the Ruteh Limestone. A very rich bed with the fusulinid Reichelina pulchra Miklukho– Maklay, Nanlingella meridionalis Rui & Sheng, and N.? simplex (Sheng & Chang) (Fig. 21) has been detected in the Bear Gully section (IR 880), corresponding to the base of the Permophricodothyris ovata biozone. Of these fusulinids, N. meridionalis is reported from the lower part of the Changhsingian in South China (Rui & Sheng 1979). Other two fusulinids are also well known in upper Wuchiapingian and Changhsingian strata in the Tethyan Realm. The lower part of the upper member (cherty limestone) in the Elikah section, up to bed IR 192, contains Nankinella cf. discoidea (Lee), Dagmarita chanakchiensis Reitlinger, Paradagmarita spp., Paraglobivalvulina spp., Reichelina sp., Hemigordiellina regularis (Lipina), Hemigordius sp., Neodiscus sp., Multidiscus sp., Graecodiscus sp., Pseudomidiella sp. with rare lagenides and can be referred to the Paradagmarita Zone ¨ zgu¨l (2001). Above the brachiopod of Altiner & O extinction horizon, i.e. above the Enteletes lateroplicatus Biozone, in the beds between IR 193 and IR 200 there is a significant decrease of abundance and a shift in the foraminiferal assemblage with an increase of lagenides (i.e. Frondina permica Sellier de Cevrieux & Dessauvagie and Rectostipulina quadrata Jenny-Deshusses) associated with small Hemigordiopsidae and rare Paraglobivalvulina sp., Nankinella cf. discoidea, Pseudoammodiscus sp. The increase in the abundance of the lagenide foraminifers and the concurrent decrease of other foraminifer taxa has been already observed

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Fig. 20. Continued.

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Nesen Formation in their Elikah Valley section (¼ our Elikah section). The absence of preserved brachiopods in the topmost part of the unit and the general decrease of the faunal content might testify to a change in environment, with flourishing of Radiolaria and unfavourable conditions for conodonts and brachiopods. The Nesen Formation appears to be very rich in ostracodes, mostly being new taxa. They will be studied by S. Crasquin-Soleau (Paris).

Fig. 21. Fusulinids from sample IR 880 (Bear Gully section) of the Nesen Formation. 1– 3. Nanlingella? simplex (Sheng and Chang), 40. 4, 5. Nanlingella meridionalis Rui and Sheng, 40. 6, 7. Reichelina pulchra Miklukho-Maklay, 50.

in the Permo–Triassic (P –T ) boundary beds of the Bulla and Tesero sections of the Dolomites and in the Taurids (Groves et al. 2007). In the Ruteh section the uppermost beds contain Pachyphloia ovata Lange, Pachyphloia schwageri Sellier de Civrieux & Dessauvagie, Langella sp., Robuloides lens Reichel and bioclasts of the alga Permocalculus sp. Surprisingly enough, the conodont content in these facies is almost barren. Up to now only a single element of Hindeodus julfensis (Sweet) was found at the base of the cherty limestones. Also the Iranian–Japanese Research Group (1981) found only Gondolella orientalis Barskov & Koroleva in the lower part of the upper member of the

Age. The Nesen Formation is referable to the Late Permian, with evidence both of Wuchiapingian and Changhsingian stages. The lower member is Wuchiapingian. This assignment is firstly supported by brachiopods. The Araxilevis intermedius Biozone is correlative of the Araxilevis Biozone of the Djulfa section, considered of Dzhulfian (¼ Wuchiapingian) age, due to its association with the ammonoid Araxoceras (Leven 1981). The palynological content of the lower member does not contradict this assignment, even if age indications are less distinct. Distriatites insolitus occurs only in OSPZ5 or younger sediments in the Arabian peninsula (Stephenson et al. 2003), and ‘Maculatasporites lignin type’ occurs only in OSPZ4 or younger sediments. Triquitrites proratus has been recorded from the upper part of the Chhidru Formation in the Salt Range of Pakistan (Balme 1970), which can be dated reliably by conodonts as Changhsingian (Wardlaw & Pogue 1995). The presence of Distriatites insolitus and ?Florinites balmei (index taxa for the OSPZ5 and OSPZ6 biozones, respectively; see Stephenson et al. 2003 and unpublished information), and Columinisporites sp., Pyramidosporites racemosus and specimens similar to Playfordiaspora cancellosa indicate that the unit is not older than OSPZ6 (Wordian –Capitanian; see Stephenson et al. 2003), supporting the age suggested by the brachiopod assemblages. The absence of Aratrisporites indicates that the age of the section is not likely to extend into the Triassic (see Ouyang & Utting 1990). The age of the lower part of the upper member is still open. The Paradagmarita zone is referred to the late Wuchiapingian–Changhsingian

Fig. 20. (Continued) Palynomorphs from the Dorud and Nesen formations. The locations of specimens are given first by England Finder reference, then by slide reference number. Slides and sample records are held in the Palaeontological Museum of the Department of Earth Sciences, University of Milano, Italy. 1 –12 from the Dorud Formation; 13– 18 from the Nesen Formation. 1. ?Barakarites sp., T63, IR 52/7, 200. 2. ?Corisaccites alutas, C59/1, IR 830/6, 272. 3. ?Kingiacolpites subcircularis, S45, IR 830/7, 625. 4. Potonieisporites novicus, F45, IR 52/1, 225. 5. Striasulcites tectus, F69/3, IR 52/3,550. 6. Potonieisporites novicus, S66/2, IR 53 g,250. 7. ?Diexallophasis sp., H58, IR 52/4, 290; probably reworked. 8. Corisaccites alutas, F70/1, IR 52/6, 315. 9. Vittatina costabilis, K63/4, IR 52/6, 545. 10. ?Barakarites sp., J34, IR 52/6, 175. 11. Vesicaspora sp., L48, IR 52/6, 450. 12. ?Complexisporites sp., R42/4, IR 53e, 500. 13. Lundbladispora sp., T46/3, IR 315, 450. 14. Triquitrites proratus, W41/1, IR 312, 450. 15. Triquitrites cf. proratus, M33, IR 312, 450. 16. Distriatites insolitus, C39, IR 312, 225. 17. ?Florinites balmei, L66/2, IR 330, 450. 18. Pyramidosporites racemosus, O63/3, IR 312, 450.

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¨ zgu¨l 2001). The fusulinid assemblage (Altiner & O (IR 880 in the Bear Gully section) suggests an early Changhsingian age. Also, the occurrence of the conodont Hindeodus julfensis in sample IR 157 below the Permophricodothyris ovata Biozone supports this age assignment, as its distribution ranges from the Wuchiapingian to the lower Changhsingian (Wardlaw & Mei 1998). The middle and upper parts of the upper member of the Nesen Formation are more confidently referred to the Changshingian. Even if the brachiopod ranges are variable, as said above, the Permophricodothyris ovata Biozone may be younger than the Oldhamina Biozone of Ruzhentsev & Sarycheva (1965) which contains the ammonoid Araxoceras of middle –late Dzhulfian (¼ late Wuchiapingian) age, whereas the Enteletes lateroplicatus Biozone correlates with the Haydenella Biozone and ‘Comelicania’ (¼ Gruntallina) beds of the Djulfa area which are Changhsingian due to the occurrence of Phisonites and Gondolella subcarinata (Leven 1998; Shen et al. 2004). This is also supported by the faunal turnover at the base of the Permophricodothyris ovata Biozone and the similarity of both biozones with the lower Changhsingian assemblage of South China (Xu & Grant 1994). Finally, the genus Ombonia is a characteristic component of the low-diversity brachiopod faunas of latest Permian age, such as those from the Tesero Member, Werfen in the Dolomites, Italy (Broglio Loriga et al. 1988; Beretta et al. 1999), the Gerennavar Limestone in the Bukk Mountains, Hungary (Posenato et al. 2005), and the lower Kathwai Member, Mianwali of the Salt Range, Pakistan (Grant 1970). The topmost beds, still Permian in age, are less defined, because of reduced fossil content. Environment. The environment of the Nesen Formation is interpreted as a carbonate ramp, not very deep, on which the distal part of a terrigenous apron, fed by emerging land, was initially expanded. The transgression towards the south and to the east caused backstepping of the clastic facies and of the source areas. The sedimentation on the ramp became mostly carbonate, with chert derived by the dissolution of radiolarian tests during the diagenesis.

Qeshlaq Formation Name. Established by Jenny & Stampfli (1978) under the spelling Gheshlagh, with type section to the SW of the Qeshlaq village (Fig. 22). Lithology. The unit consists mostly of arenite and siltstone, with red and red –brown dominating colours, sometime blackish in the silty and shaley intercalations. Conglomerates are extremely rare.

Arenites form layers usually less than 1 m thick, with low-angle cross-laminations and, rarely, high-angle point bar festoons. Parallel laminations in the arenites are subordinate, as well as small ripple-marks with anastomosing ridges. Occasionally dark shales with plant fragments are interbedded (Chateauneuf & Stampfli 1978). Pedogenesis may occur at several horizons, with nodular reddened surfaces and caliche nodules. In the middle and upper parts of the Qeshlaq Formation in the Qeshlaq and Ghosnavi sections several lateritic horizons are observed (Fig. 23). In the topmost part of the unit the bedding planes locally show iso-oriented disarticulated shells of bivalves, the arenites are very fine and contain fragments of marine fossils. To the east (Qeshlaq and Ghosnavi) the Qeshlaq Formation represents a single depositional sequence that displays excellent correlations between the two measured sections that lie 7 km apart one from the other. The sequence starts with moderately sorted subcalclithites, displaying asymmetrical ripplemarks and clay chips (IR 710–IR 711 at Ghosnavi; IR 733 at Qeshlaq) passing upwards to well-sorted, fine-grained quartzarenites with syntaxial quartzose cement (IR 712 at Ghosnavi; IR 735 at Qeshlaq). These sandstones, deposited during a stage of aggradation of alluvial facies (HST, highstand systems tract), are overlain by lateritic paleosols with abundant nodular concretions (pisoids: IR 713), corresponding to the LST (lowstand systems tract). The following transgression is documented initially by moderately sorted, medium-grained sandstones with embayed quartz and abundant silcrete/ferricrete intraclasts, pointing to the ravinement of soil profiles developed during the previous LST (IR 714–IR 716 at Ghosnavi; IR 737–IR 739 at Qeshlaq). Eventually, the unit is capped by thin-bedded ferruginous arenites and siltstones (IR 720– IR 723 at Ghosnavi; IR 741–IR 742 at Qeshlaq) with interbedded bioclastic cocquinas (storm deposits?) containing abundant ostracodes, bivalves and gastropods, and thus documenting the encroachment of shallow-marine, probably paralic settings. Sandstone petrography. The terrigenous beds consist of very fine- to medium-grained quartzarenites and sublitharenites, mostly moderately sorted, with rare altered feldspars. Quartz is almost exclusively monocrystalline, with multiple syntaxial overgrowths particularly conspicuous in the Qezelqaleh section, whereas lithic grains include prevailing carbonate and subordinate volcanic rock fragments: the latter, displaying vitric – felsitic textures and containing glass shards and spherulites, are particularly abundant in the Qezelqaleh section. Heavy minerals are limited to an ultrastable ZTR suite, with very few detrital mica

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Fig. 22. Measured logs in the Qeshlaq Formation and across the P –T boundary. WGS84 co-ordinates for the base of the logs. Ruteh section N358580 5300 , E518320 3100 ; Gheselghaleh section N368450 17.500 , E548490 11.200 ; Abrendan section N368210 21.300 , E548150 56.700 ; Qeshlaq section 368540 0400 , E558210 0000 ; Ghosnavi N368550 57.800 ; E558260 32.200 .

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Fig. 23. (A) Outcrop of a lateritic horizon, Ghosnavi section, IR 718, scale bar 10 cm. (B) Lateritic pisolithe, sample IR 1026, Abrendan section, 5. (C) The lateritic soil material has been reworked and sedimented in the fine quartzarenitic alluvial plain. Sample IR 858, Bear Gully section, 10. Microphotographs courtesy of L. Trombino, Milano.

flakes (consistent with the absence of metamorphic and plutonic rock fragments). Widespread pseudomatrix (Dickinson 1970) consists of green clays – including chamosite – grown at the expense of labile grains, and of silcrete –ferricrete intraclasts from soils. Compatible cements (quartzose in quartz sandstones, carbonate in bioclastic ferruginous arenites) are also widespread. Widespread pseudomatrix derives from the deformation of green clayey non-carbonate intraclasts (Zuffa 1980; Garzanti 1991), consisting of nearly stoichiometric chamosite (Table 1), a ferroan chlorite whose formation is favoured by leaching of Fe-enriched soil profiles by sea water concurrent with reduced clastic input (Odin 1988). In the Abrendan section, a basal conglomerate in continental facies passes upwards to fine- and very-fine-grained quartzarenites with syntaxial cements and sublitharenites containing subordinate carbonate and volcanic lithics. Greenish clayey pseudomatrix is particularly common towards the top of the unit and is observed also in fine-grained sandstones at the base of the overlying Nesen Formation. A similar ferruginous pseudomatrix

also characterizes the upper, sandier part of the Qeshlaq Formation in the Qezelqaleh section. Thickness and distribution. The Qeshlaq Formation is typical for the SE part of the east Alborz, where it locally reaches a maximum thickness of about 80 m in the type section and at Abrendan, whilst it is reduced to 34 m in the neighbouring Ghosnavi, being cut by a fault at its base. It reaches some 40 m in Qezelqaleh, to be reduced to merely 22 m at Ruteh (Fig. 22). In the southern part of the central Alborz, the Qeshlaq Formation may be largely missing altogether and only a few sandstone beds and residual lateritic soils (Aruh section) are preserved below the Elikah Formation. In the Shahmirzad section, all of the interval between the Mobarak Group and the Elikah Formation consists of terrigenous rocks, and it is difficult to separate the part pertaining to the Dorud Group from that of the Qeshlaq. The part with lateritic horizons is some 20 m thick. Biochronology. The fossil content of the Qeshlaq Formation is poor and consists mostly of pollens

Table 1. Scanning electron microscopy-energy dispersive spectrometer (SEM-EDS) normalized analyses on chamosite pseudomatrix in ferruginous arenites at the top of the Qeshlaq Formation, type section (courtesy of A. Rizzi, CNR-IDPA, Milan)* IR 741

IR 742

TiO2 SiO2 Al2O3 FeO MnO MgO CaO Na2O K2O [H2O] Tot

0.09 27.25 19.88 32.49 0.06 5.40 0.63 0.69 0.07 13.44 86.56

0.03 26.45 20.28 32.79 0.00 4.65 0.61 0.83 0.21 14.15 85.85

0.00 25.51 18.64 31.95 0.00 4.59 0.77 0.57 0.21 17.77 82.23

0.00 29.95 19.83 26.95 0.16 7.12 0.95 0.84 0.03 14.16 85.84

0.20 26.71 19.43 33.19 0.12 4.65 0.75 0.68 0.08 14.18 85.82

0.00 25.91 18.91 34.78 0.12 5.45 0.31 0.76 0.07 13.67 86.33

0.00 25.67 19.70 33.62 0.00 5.00 0.46 0.71 0.22 14.62 85.38

0.00 24.77 19.20 34.74 0.00 5.02 0.43 0.68 0.08 15.09 84.91

[Tot]

100.00

100.00

100.00

100.00

100.00

100.00

100.00

100.00

*Live time is 50 s, metallic cobalt for standard; figures in square brackets are recalculated theoretically.

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and some plant remains. Samples collected in the two more typical sections (Qeshlaq and Ghosnavi) were barren. We get poorly preserved palynomorphs, wood fragments and sheet cellular materials from the Ruteh section (sample IR 1088). Palynomorphs are brown– dark brown in colour, indicating relatively high thermal maturation. The following were recorded: indeterminate bisaccate and taeniate bisaccate pollen and spores, Alisporites indarraensis Segroves, Leiosphaeridia spp., Cyclogranisporites spp., Punctatosporites spp. and a possible specimen of Pyramidosporites sp. The presence of taeniate bisaccate pollen, Alisporites indarraensis and ?Pyramidosporites sp. suggests a broad Permian age for the assemblage, but no precise assessment can be given. However, a sample (IR 614) collected above the Ghosnavi Formation, along the road of the Jauzchal section, and tentatively assigned to the Qeshlaq Formation yielded a well-preserved and rich palynomorph assemblage containing common Distriatites/Hamiapollenites spp. (similar to Hamiapollenites dettmannae Segroves), indeterminate monosaccate and bisaccate pollen, Alisporites cf. nuthallensis and A. indarraensis, as well as Vittatina costabilis Wilson, Punctatisporites spp., Plicatipollenites malabarensis (Potonie´ & Sah) Foster, ?Lundbladispora spp., Protohaploypinus cf. limpidus, Potonieisporites spp., Lundbladispora spp., Protohaploypinus limpidus (Balme & Hennelly) Balme & Playford, Leiosphaeridia spp., ?Playfordiaspora sp., ?Distriatites sp., Nuskoisporites sp., Laevigatosporites cf. callosus Balme, Vesicaspora sp., Indotriradites cf. apiculatus Stephenson and Osterloff, Distriatites insolitus Bharadwaj & Salujah, Striomonosaccaites sp., Leiosphaeridia spp., ?Schweitzerisporites spp. and ?Weylandites sp. The assemblage is very similar to those of the Khuff Formation of Arabia, indicating a probable Mid-Permian age. Age. Direct evidence for the age of the Qeshlaq Formation is scanty. Chateauneuf & Stampfli (1978) and Stampfli (1978) considered it to be of Late Permian age, by correlation to the marine Nesen Formation. However, the sample (IR 614) from the Jauzchal section suggests an older age. We consider the Qeshlaq Formation mostly as Late Permian, interpreting it as lateral equivalent of the well-constrained Nesen Formation. Environment. The environment of the Qeshlaq Formation is transitional –continental. In the Abrendan section it is represented by a continental alluvial plain and a near shoreface, with a couple of lateritic episodes. In the Qeshlaq–Ghosnavi area, however, sedimentation is continental throughout the unit, with alluvial plain facies, including also isolated

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marsh deposits (Chateauneuf & Stampfli 1978) and deep soil profiles, up to laterites, that were episodically transgressed by brief episodes of ferruginous, shelly accumulations (tempestites) at the top. Intrabasinal grains in sandstones suggest the co-existence of lateritic soils on emerged areas and of tolerant benthic communities on a shallowmarine shelf, that rapidly encroached onto continental areas during the final stages of deposition of the unit. These repeated transgressions reworked the soils formed on the continental alluvial plain. Mineralogy of sandstones from the Qeshlaq Formation points to a scenario of erosion of older sedimentary successions (mostly quartzose sandstones and carbonates), locally mantled by intermediate –acidic volcanics.

The relationships between the Nesen and Qeshlaq formations In the present paper we consider the Nesen and Qeshlaq formations as lateral equivalents, belonging to the same depositional sequence (Fig. 2). We recognize that further fieldwork and data will add more substantial evidence, but this seems to us the most logical interpretation. Physical evidence is found in the Ruteh and Abrendan sections, where the terrigenous sediments are interbedded with a marine bioclastic interval and are overlain by the cherty limestone of the upper part of the Nesen Formation. To the north and to the west, the shales characterizing the lower part of the Nesen Formation are considered the lateral and distal equivalent of the terrigenous spells of the Qeshlaq Formation. We disagree with the interpretation of the Qeshlaq Formation as a lateral equivalent of all the Middle and Upper Permian succession (Jenny & Stampfli 1978). It would be difficult to explain how, during the sedimentation of the Ruteh Limestone, the alluvial plain continued to aggrade and carry terrigenous sediments that are not preserved in the basin. Therefore, a sedimentation gap, possibly linked to an aridity phase, should also have occurred on the emerged lands. This does not exclude that very local episodes of continental sedimentation could be preserved, as the case of Jauzchal, where the fine arenites and siltstones are associated with caliche nodules. The reappraisal of the alluvial system would also be testified to by the lateric horizons that require sufficient humidity to evolve. Local reduction of the accumulation rate enabled a major concentration of silica, with large cherts nodules and whole silicified beds. Towards the top of the succession, the worldwide crisis of the radiolarians community near the Permian– Triassic boundary (Kidder & Worsley 2004)

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caused the reduction and then the disappearance of the silica content in the succession.

The Permian – Triassic boundary interval The upper limits of the Nesen and Qeshlaq formations approach the Permian– Triassic boundary. Owing to the significance of this boundary it is discussed separately. For most of the Alborz Mountains, the marine Triassic Elikah Formation transgresses onto the continental alluvial Permian substrates, suggesting the onsetting of a new subsidence pattern for the area. However, in the central Alborz (Ruteh, Ruteh Valley, Elikah, Dasht-e-Nadir and Bear Gully sections: Figs 22 and 24) the boundary between the two units is more gradual. The nodular limestone (wackestone/mudstone) of the Nesen Formation gradually reduces its cherty content and it is overlain, often with an inconspicuous ravinement surface, by the chert-free limestone of the Elikah Formation, often oolitic. In the Ruteh section, 1.2 m of grey wackestone/mudstone, with sparse oolites, yielded at their top the conodont Hindeodus parvus Kozur & Pjatakova and Isarcicella turgida Perri, index for the base of the Triassic. Hindeodus parvus was found also in the Elikah type section (Iranian– Japanese Research Group 1981). The limestones are overlain by recrystallized microbialites with domal structures, for about 2 m. Recrystallized laminar microbialites continue for

the following 18 m. In the nearby Ruteh Valley section, spectacular giant domal stromatolites of the basal Elikah Formation are exposed (Fig. 25). To the north, along the belt Elikah– Nesen, the domal stromatolites were not observed, but the planar microbialites of the basal Elikah are more than 20 m thick. Both facies are typical for the anachronistic carbonates that characterize the base of the Triassic (Baud et al. 2007). We processed some 20 samples but no conodonts were obtained in the microbialites. In the Elikah type section, the Iranian –Japanese Research Group (1981) obtained Isarcicella isarcica (Huckriede) a few metres above the base of the Elikah Formation. The topmost Permian ages are less documented, because of reduced fossil content. Small and abundant Permian Hemigordiopsidae are present in the basal beds of the Elikah Formation. These are capped by 50 cm of abiotic oolitic limestones in the Elikah section and then the first occurrence of the disaster taxon ‘Cornuspira’ mahajeri Bro¨nnimann, Zaninetti & Bozorgnia indicates an Induan age. This is the evidence that the base of the Elikah falls in the latest Permian and corresponds to a transgressive tract in the sequence. The d18O and d13C isotope curves along the Permo-Triassic succession of the Elikah section (Fig. 26) confirm that there is no significant gap at the Permo-Triassic boundary in the central Alborz succession. Positive d13C values above þ3‰ have been recorded for the Wuchiapingian part

Fig. 24. Palaeogeography of the central and eastern Alborz around the P– T boundary time.

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Fig. 25. Domal stromatolites at the base of the Elikah Formation. (A) vertical cut of a domal stromatolite, Ruteh section, proxy to the sample IR 1095. (B) – (D) Three-dimensional domal stromatolites in bottom, lateral and upper views; Ruteh Gorge section N358590 1300 , E518320 0600 . The hammer is 32 cm long and the white bar is 18 cm long.

of the Nesen Formation. d13C shows major shifts between þ3.84 and þ1 in the Changhsingian, and drops to very low positive values and then to negative values (20.03) across the boundary. In the Induan, values range from 20.04 to 20.93 and raise again to positive values (þ0.03 to þ2.13) in the second lithozone of the Elikah Formation. This trend is very similar to that recorded along the Tesero section in the Dolomites (Farabegoli et al. 2007), in the Abadeh section of Central Iran (Korte et al. 2004), in three sections in northern and central Iran by Horacek et al. (2007), and in the Kuh-e-yagma section in the central Zagros Mountains by Wang et al. (2007).

Elikah Formation Name. Glaus (1964) established this unit with two members, the lower consisting of various lithologies, with limestone, dolostone, marl and rare arenite, and the upper consisting of thick layered dolostone. The type section was designated to the

right-hand side of the Elikah creek. Our Elikah section corresponds to the type section of the formation. In the official maps of the GSI, the name is spelt either Elika or Elikah. We use the spelling Elikah, under the authority of the review made by Seyed-Emami (2003). Lithology. This sequence is rather uniform over the whole range (Seyed-Emami 2003). It consists of several lithological units, lithozones or members at their own rank. The lowermost one, present only where there is no major gap and unconformity with the underlying Nesen Formation, consists of thick-bedded grey limestone (mudstone/wackestone), locally partly dolomitized, up to some 20 m thick. Domal or planar stromatolites and biolithites are widespread, which has been discussed in the previous section (Fig. 25). Where, however, at the top of the underlying unit (Nesen Formation or Qeshlaq Formation) there is an emersion surface, usually the lower boundary is made of yellow –grey dolostone (a

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Fig. 26 d13C and d18O curves along the Elikah river section in the upper part of the Nesen Formation and the lower part of the Elikah Formation. The isotopic curve matches the stable isotope curve along the P –T boundary interval of the Abadeh section curve (Korte et al. 2004).

PENNSYLVANIAN –EARLY TRIASSIC ALBORZ IRAN

few metres thick) overlain by the second unit, which is ubiquitous (Fig. 14). The latter consists of thinbedded nodular grey mudstones/wackestone, locally densely bioturbated (vermicular limestone), rich in bivalves and at minor extent in microgastropods. Subordinate are microgastropod ooidal packstone or grainstone, also containing clay chips and partly indurated lithoclasts. This facies usually forms beds from 10 to 30–40 cm thick, interspersed in the vermicular limestone succession. Also very subordinate are flat-pebble conglomerates that may form 20– 40 cm-thick beds with micritic pebbles up to 15 cm in length. There are also layers that do not appear to be of conglomeratic nature on the outcrop, but in thin section consist of micritc fragments of 1–3 mm long with sharp edges and sparry cement. Very subordinate are pinkish marl and reddish fine quartzarenite. The total thickness is around 100 –150 m. The third unit consists of thin-bedded grey recrystallized limestone or dolostone, some 50–100 m thick. The fourth unit, up to 300 m thick, is made of cliffforming, thick-bedded to massive light-grey coarse dolostone. Laminar stromatolites may be locally observed. The uppermost unit consists of light-grey bedded limestone (packstone/wackestone), up to 50 m thick. The Elikah Formation is truncated by an erosional surface, and the next formation often overlies it with an angular unconformity. It may be immediately overlain by sedimentary units of the Shemshak Group (Fu¨rsich et al. 2009) or by lava flows and/or by a tuffaceous horizon. South of Sari, volcanics may be several hundreds of metres thick. Biochronology. The very basal metre, which is an oolitic grainstone, contains Permian Hemigodiospsidae in the Ruteh and Elikah sections, overlain by mudstone with sparse oolites bearing the conodont Hindeodus parvus, the index for the base of the Triassic. When the stromatolitic and biolithitic structures start to develop it appears the foraminifer ‘Cornuspira’ mahajeri, which forms a typical oligotypic assemblage for the very base of the Triassic. This species spans all of the basal unit. In the top layer of this basal unit, the conodont Isarcicella turgida Perri (IR 251) was found in the Elikah section. In the same position, the Iranian –Japanese Research Group (1981) also obtained several conodont species. The taxonomy of some of them might be revised, but most of the species suggest an earliest Induan age. In the second unit, bivalves are locally abundant, dominated by the genera Claraia and Eumorphotis. Also, endobiontic bivalves, such as Myophoria and Unionites, are present. The diffuse occurrence of Claraia confirms that the lower units of Elikah were deposited during the Induan. The ‘gastropod

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oolite’ facies is also typical for the Lower Triassic of the low-latitude Tethys (Fraiser et al. 2005). In the overlying beds the heavy dolomitization prevents any age assignment. Only the topmost calcareous unit contains rare foraminifera, whose age is disputed. In the Aruh section, Zaninetti et al. (1972) discovered Triassic foraminifera considered of Norian age. Later Zaninetti & Bro¨nniman (1974) reconsidered this fauna and attributed it to the Ladinian. Therefore, the age of the Elikah, except for the first very few decimetres, which could be still latest Permian, is Triassic. The thickness of the Lower Triassic part, which may amount to up to half of the total thickness, should be emphasized. No confirmed ages younger than the Middle Triassic are available. Environment. The Elikah Formation was deposited on a wide flat plain, submerged by a very shallow sea for most of time, with lateral gradients fairly homogeneous. Evidence for this: (1) the domal and laminar stromatolitic beds at the very base and through the thick dolostone member (Upper Member of Glaus 1964 in the type section); (2) the flat-pebble edge-wise conglomerate linked to desiccation crusts stripped away during storms (Pruss et al. 2005); and (3) the gastropod ooids beds requiring high wave energy for their formation. The terrigenous input was feeble or nonexistent, concentrated in the second member, with clay, silt and very occasional fine sand. Therefore, it would suggest both a very feeble land relief and an absence of significant rivers reaching that part of the coast line. A modern counterpart could be the Arabian side of the Gulf. However, the absence of extended evaporitic deposits, notwithstanding the subtropical latitude of that part of the Iranian plate, should be noted.

Depositional and geodynamic interpretation The Pennsylvanian– Permian sequences of central and eastern Alborz indicate that the depocentres were situated toward the NW of the range, whilst the marginal, transitional or emergent areas were to the south and east. Towards the depocentres the total thickness of the mostly marine successions may reach 1200–1500 m, whereas to the south it may be reduced to few tens of metres. The major sequences are bounded by unconformities and depositional hiatuses, represented by emersion surfaces. They are thus unconformity-bounded sequences (Van Wagoner et al. 1988) and might be interpreted as second-order sequences with significant intervening time gaps. Only the contact between the Nesen and Elikah formations is less sharp, marked by a disconformity of only limited

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extent with a local ravinement surface. The onset of the Elikah sequence occurred, however, under different depositional and palaeogeographic conditions. Discussion on the sequence interpretation will be developed in a separate paper. Three major factors control the Pennsylvanian– Early Triassic evolution: global sea-level change, regional geodynamics and climate.

Global sea-level change The Carboniferous and the earliest Permian times were characterized by the Gondwana glaciations and related lowstands of sea level, followed by the late Sakmarian deglaciation, generally periods of highstands of sea level that persisted until the end of the Middle Permian. Subsequently, the trend was to lower sea-level stands, with a recovery during the Early Triassic (Crowell 1995; Wopfner 1999; Kidder & Worsley 2004; Scheffler et al. 2006). Comparing the Alborz record with the glacio-eustatic record for the Pennsylvanian and Early Permian (Fig. 2), we observe that the increase of terrigenous input in the Serpukhovian and the internal erosion surfaces in the Dozdehband Formation may be linked to the high-frequency cycles occurring during the Glacial II interval (Isbell et al. 2003, fig. 4; Jones & Fielding 2004). The Moscovian Qezelqaleh Formation may be linked to the glacio-eustatic highstand, corresponding to the interval between Glacial II and Glacial III. Particularly, it could be linked to the lower Moscovian peak in coastal onlap (Gradstein et al. 2004). In contrast, the uppermost Carboniferous– lowermost Permian Dorud sequence and the 15– 20 Ma gap in the Early Permian –earliest Middle Permian are out of phase with the global trend. The aggradation of the Dorud Group occurred when the glaciation reached its climax during the Glacial III interval (Montan´ez et al. 2007). Inside the Dorud Group, the carbonate units (Emarat & Ghosnavi) may represent the coastal onlap of the middle Asselian, i.e. a minor aggradation inside a regressive trend (Gradstein et al. 2004). The massive deposition of oncoidal limestones in the Asselian part of the Emarat Formation is peculiar. Shi & Chen (2006) related this facies in southernmost China to a high-energy, shallow-water carbonate platform environment in tropical settings, arguing that Lower Permian oncoidal limestones are the answer to prolonged lowstand eustatic condition during a global sea-level drop caused by the onset of the major phase of the Gondwana glaciation. Shapiro & West (1999) emphasized the very frequent occurrence of stromatolites and oncolites at low latitudes around the Carboniferous –Permian

boundary, as well recently observed in Turkish Taurids (Kobayashi & Altiner 2008). When the deglaciation inundated the continental margins, the Alborz was initially the site for erosion and/or terrigenous sedimentation, eventually to experience a long gap in sedimentation. The deglaciation succession (Stephenson et al. 2007) is totally missing. With the Middle Permian transgressional event represented by the Ruteh Limestone, the Alborz area again returned to the global sea-level pattern, and remained in phase for the Late Permian and Early Triassic. We refer the generalized emersion of the top of the Ruteh Limestone as linked with the global low standing at the base of the Upper Permian (Jin et al. 1998, 2001), where the boundary between the Middle and Upper Absaroka Megasequences also is drawn (Haq et al. 1988; Gradstein et al. 2004).

Regional geodynamics Most palaeogeographic reconstructions locate the Alborz and central Iran close to the Gondwana margin, assigning them to the terranes of the Peri-Gondwanan Fringe and later to form the Carboniferous and earliest Permian Megalhasa or Cimmerian Continent (Sengo¨r 1979; Berberian & King 1981; Dercourt et al. 1993; Stampfli & Borel 2002). According to these models, Pennsylvanian rifting developed between Arabia and Sanandaj–Sirjan (Iran) (Al-Belushi et al. 1996; Jassim & Goff 2006) and culminated in middle of the Early Permian spreading, giving birth to the Neotethys (Garzanti 1999; Angiolini et al. 2003). According to the interpretation developed in this paper, the Alborz acted in this time as the northern margin of the central Iran microplate. It did not share the evolution of the Arabian margin, which is in tune with the glacio-eustatic signal during this period (Sharland et al. 2004; Insalaco et al. 2006). A more complex pattern controlled the evolution of the microplate. Our data suggest that the Alborz evolution was predominantly controlled by geodynamics when rifting with Arabia started to be particularly active. Vertical movements linked to the rifting may explain the late Moscovian –Kasimovian gap. Later prosecution of the rifting may have caused tilting of the Iran microplate, allowing sedimentary deposition notwithstanding the sea-level drop around the Carboniferous–Permian boundary. The absence of the deglaciation sequence is another hint at a tectonically complex history. In the evolution of a simple passive margin, in which the Dorud Group should represent the tectonic subsidence phase, the absence of the deglaciation sequence and the consequent sedimentation gap is problematical. They may be related to the

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complexity of tectonic movements that lead to the fragmentation of the Gondwanan margin. In particular, we note this time lag is coeval with the megashear active during the shift from Pangaea B to Pangaea A (Muttoni et al. 2003). Palaeomagnetic and palaeobiogeographic data (Angiolini et al. 2007; Angiolini & Stephenson 2008; Muttoni et al. 2009) constrain central Iran to a more equatorial and western position than usually indicated in global palaeogeographic reconstructions, and consequently closer to the megashear zone. In the Middle Permian a wide outer-marine shelf occupied the margin of the Alborz (but its SE part) as well central Iran, sharing an evolution common to other Cimmerian blocks, such as Transcaucasia and Taurids, and to the Gondwanan margin. With the opening of the Neotethys, according to many authors, the Cimmerian blocks moved northwards and the subduction of Paleotethyan oceanic crust was activated along the Eurasian margin, persisting during most of the Permo-Triassic (e.g. Alavi 1991; Ruttner 1993; Dercourt et al. 1993, 2000; Besse et al. 1998; Stampfli & Borel 2002; Zanchi et al. 2009). The main opponent of this proposition is Sengo¨r (1979, 1990), who suggested that Paleotethyan subduction was also initially directed southwards underneath the Cimmerian terranes. This interpretation has recently been reconsidered by Jassim & Goff (2006) and Ruban et al. (2007), who supposed a southwards subduction active until the Triassic. In our interpretation, after the Middle Permian the Alborz evolved as the northern passive margin of the Iran microcontinent, while areas south to the Alborz subsidence was low or even absent, therefore areas remained emergent (Fig. 27). We infer that some locally thick sections of the Ruteh Limestone may have been triggered by normal faults. This setting may explain: † the unusually large thickness of the Ruteh Limestone towards the NW (an undecompacted accumulation from 25 m Ma21 to the south up to 80 –90 m Ma21 in the north); † the presence of volcanic products in the middle – upper part of the Ruteh Limestones, at the base of the Nesen Formation, and as clasts in the Qeshlaq Formation sandstones. Felsitic and vitric rock fragments are derived from the groundmass of intermediate –acidic volcanics; microlitic structures are, however, diagnostic of intermediate –basic lava flows (Dickinson 1970). Their concurrent occurrence is consistent, in principle, with a bimodal volcanic suite, which is common in rift settings. This evolution temporarily halted with the emersion and karstification on the Ruteh Limestone, even where major subsidence was observed

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Fig. 27. Cartoon suggesting the possible geodynamic setting of the Alborz during the Pennsylvanian– Early Triassic times.

(Emarat anticline, Mangold section), because of the sea-level fall around the Middle–Late Permian boundary. It resumed in a similar setting in the Late Permian, as evidenced by protracted emersion to the south and SE, progressive onlap of the upper part of the Nesen Formation in the Central Alborz towards southern directions and preservation of the most complete Nesen succession to the north. The accumulation rate did not exceed 20 m Ma21 also to the north, even though the subsidence rate was higher there. This pattern changed with the general transgression of the Elikah Formation over the whole area during the earliest Triassic, with higher rates of deposition in comparison to the Middle and Upper Permian (Brunet et al. 2007). The topographic relief over the emerged area was apparently never very high, as indicated by the persistently fine grain size of the terrigenous sediments. We

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suppose that this change may be at least partly related to the approach of the Alborz to the oblique subduction zone of the Turan arc (Natal’in & Sengo¨r 2005). The distance between the active margin of the Turan arc to the north and the Alborz margin might not have exceeded 500– 700 km, according to the palaeomagnetic data of Feinberg et al. (1996) and Muttoni et al. (2009). The Iranian lithospheric microplate consisted of a remnant of the oceanic Palaeotethys, the continental crust of Alborz – central Iran and the oceanic part of the Neotethys so that deformation should have been concentrated in the transition areas between the different kinds of lithosphere (Turcotte & Schubert 2002), thus leaving the bulk of the area free of significant tectonism.

Climate The third major element of control of the late Paleozoic section is linked to the climate. With the northward drift, the Iran microplate moved from a tropical latitude in the southern hemisphere to a tropical latitude in the northern hemisphere (Besse et al. 1998; Muttoni et al. 2009). The passage through the tropical aridity belt is possibly the reason why there is no significant terrigenous transport from the emerged area to the basin during the Middle Permian. The Ruteh Limestone is significantly arenite-free. Entering in the humid equatorial zone during the Late Permian, the humid climate enhanced the laterization of soils on the emerged areas, the alluvial and deltaic system resumed, and the fine fraction of clay and silt were transferred to the basin, forming the lower part of the Nesen Formation. Carbonate tidal flats of the Elikah developed with continued drift to the north, when the central Iran microplate returned to a arid tropical latitude.

Correlations The Pennsylvanian –Early Triassic of the central and eastern Alborz displays close affinities with the central Iranian blocks of Tabas and Anarak (Yazd block) sharing similar facies, timing of transgression–regression and unconformities (SeyedEmami 2003; Leven et al. 2006). Leven & Gorgij (2006) recognized three lithostratigraphic units (groups) related to three major transgressive– regressive cycles in Pennsylvanian–Permian strata of central and east Iran. They broadly correspond to the late Serpukhovian– Moscovian, Gzhelian – Sakmarian and Kungurian–Changshingian (Bolorian– Dorashamian in their terminology), and are referred to as the Sardar Group, the Anarak Group and the Shirgesht Group, respectively. Among them, the Sardar Group is subdivided into the

lower, Ghaleh Formation, and the upper, Absheni Formation; the latter of which is represented predominantly by sandy shales. According to Leven et al. (2006), the basal beds of the Absheni Formation contain a fusulinid assemblage similar to that of the Kashirian (Lower Moscovian) of the Russian Platform. Our Toyeh fusulinid assemblage is considered to be broadly coeval to this Absheni assemblage. Therefore, the Qezelqaleh and Absheni formations are thought to belong to the same depositional cycle. The Absheni Formation is unconformably overlain by the Zaladou Formation of the Anarak Group in central and east Iran, in which fusulinid assemblages characterized by Gzhelian Rauserites and Ruzhenzevites and Asselian Pseudoschwagerina are found (Leven & Taheri 2003; Leven & Gorgij 2006). These Zaladou assemblages are essentially similar to those from the Dorud Group in the Alborz Mountains. In the Shirgest area (eastern Iran) Leven & Vaziri (2004) described the Bagh-e-Vang Member, later upgraded to the rank of formation as the base of the Shirgest Group (Leven & Gorgij 2006). The age of the Bag-e-Vang Formation is thought to be latest Early Permian (Leven & Vaziri 2004), whilst sediments of this age are not yet reported in the Alborz. The monotonous succession of limestone and dolostone of the Jamal Formation in east Iran may be correlated with the Ruteh Limestone and perhaps the lower part of the Nesen Formation in the Alborz. However, laterites have never been reported in central Iran at this level. Affinities with the sedimentary belt of Abadeh – Shahreza–Hambast are less obvious, because the succession of this belt is more complete and with more diversified fauna (Iranian –Japanese Research Group 1981; Baghbani 1993; Kobayashi & Ishii 2003). However, the timing of sedimentation and gaps, and the position of the unconformities, are the same, indicating that they all shared the same geodynamics, being parts of the same lithospheric microplate. On the basis of these affinities, there is no evidence for the Alborz to belong only to the Hun Superterrane (Ruban et al. 2007). Recently, Davydov & Arefifard (2007) illustrated a different succession in the Kalmard area (West Tabas), located within the Posht-e-Badam block, one of the four blocks forming the central Iran (Alavi 1991). The late Sakmarian –Artinskian Khan Formation has time of deposition that corresponds to a major unconformity between the Sardar and Jamal formations in the Shirgesht and Ozbak Kuh areas; the Kalmard area emerged during the Middle and Upper Permian, to be unconformably overlain by the Triassic Sorkh Shale. Therefore, the Khan Formation was deposited in

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the gap time between the Dorud Group and the Ruteh Limestone in the Alborz. Its fusulinid fauna is of Peri-Gondwana affinity, demonstrating that the mosaic of the central Iran blocks does not have a single palaeogeographic origin. The mosaic appears further complicated in a very recent paper by Bagheri & Stampfli (2008).

Conclusions After 30 years since the last attempt to synthesize the stratigraphy of the Alborz (Jenny & Stampfli 1978), a new survey of the Pennsylvanian –Early Triassic has been carried out, based largely on original data. Five major depositional sequences have been identified; all separated by unconformities with emersion surfaces. The lowermost sequence, corresponding to the Qezelqaleh Formation, consists of quartzarenite– litharenite alternating with shale and bioclastic arenitic limestone, preserving fusulinid and brachiopod faunas. Its age is Moscovian and it is restricted to a limited area in the eastern Alborz. The underlying rocks are Serpukhovian or, locally, of earliest Bashkirian age. However, on the larger part of the Alborz, the second sequence, formed by the Dorud Group, directly overlies the Lower Carboniferous successions. The Dorud Group, up to 600 m thick, is typically tripartite, with arenitic –shaly units at the bottom and top, separated by prevailing carbonatic units in the middle. New formational terms are used for the arenitic units, the Toyeh and Shah Zeid formations, respectively. For the carbonate lithosome, the Ghosnavi Formation introduced by Jenny & Stampfli (1978) is accepted, while the new name, Emarat Formation, is used to replace the invalid term, Calcaires de Dorud. The Emarat Formation consists of packstone –grainstone cycles, with larger oncoids arranged in fining-upwards parasequences, and fusulinid packstone. The general trends are to evolve to a lower-energy sheltered environment of deposition, with prevailing mudstone/wackestone. The reappraisal of the terrigenous sedimentation occurred with the Shah Zeid Formation, with subarkose at the base, rapidly overlain by quartzarenite and sublitharenite. Only basinwards was a single carbonate intercalation observed, with a brachiopod faunule. A major gap of 15–20 Ma occurs between the Dorud Group and the Ruteh Limestone. A rather uniform blanket of packstone/wackestone is widespread on the Alborz, missing only in the most eastern and southern areas, while it is much thicker basinwards. Its age is Middle Permian. The Ruteh Limestone was involved in the emersion that affected the entire area towards the end of the

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Middle Permian. Karst features and lateritic soils are well developed. A new transgression occurred with the Late Permian. Terrigenous sedimentation on the alluvial plains to the east and SE, and marine sedimentation to the NW, formed the couple Qeshlaq– Nesen formations. Deepening was more significant basinwards, to the NW, where cherty limestone and chert layers also developed mostly during the Changshingian. A detailed brachiopod biozonation is proposed. An emersion concerned the Alborz once more towards the end of the Permian, with the exception of a limited area in the NW that had very minor hiatuses in sedimentation. In that area, the Permian –Triassic boundary interval is represented by the succession, in ascending order: sparse oolitic beds, domal stromatolites and planar microbialites, forming a package up to 20 m thick, referred to the first lithozone of the Elikah Formation. The presence of H. parvus and the typical shift in the dC13 isotopic curve allow recognition of the base of the Triassic. Most of the second and third units of Elikah Formation are represented by nodular vermicular limestone, wackestone/mudstone, well-bedded dolomitic limestone, fine siltstone and shale, often reddish, and very minor arenite. The age of the first three units is Early Triassic. A very thick package of coarse light dolostone, with occasional stromatolite layers, forms the bulk of the upper part of the Elikah Formation. The depositional setting was controlled by three main factors: sea-level changes; geodynamic evolution; and climate. The Alborz succession was in phase with global sea-level fluctuation, except for the interval from latest Carboniferous to Early Permian. During the Glacial III interval (Isbell et al. 2003), sedimentation occurred in the Alborz notwithstanding sea-level drop, while widespread ice sheets covered much of Gondwanaland. In the Sakmarian, Alborz progressively emerged while general deglaciation occurred, moving sea to high-stand conditions. Consequently the deglaciation succession is totally absent. The geodynamic scenario suggests that the Alborz, together with central Iran and the Abadeh –Shahreza –Hambast strip (but not the Kalmard area in central Iran), were already distinct from the Arabian passive margin of Gondwana in the Pennsylvanian, because they do not share any sedimentation and palaeobiological elements with that area. They should lie much more towards the equator and westwards than previously considered (Angiolini et al. 2007; Muttoni et al. 2009). In Late Permian–Early Triassic times, the Alborz approached the Turan Plate margin, with deeper sediments northwards, an emerging area to the

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south during the Late Permian and a rapid increase of the sedimentation rate in the Early Triassic. The climate had a very important control. The Alborz and most of central Iran moved from a subtropical position in the southern hemisphere during the Pennsylvanian– earliest Permian, through a tropical arid phase in the late Early Permian to approach the equatorial latitudes during the Middle Permian and especially in the Late Permian, controlling the lateritic soil formation. With the continuous fast northwards drift, the Alborz returned to tropical arid conditions in the northern hemisphere during the Middle Triassic, with carbonate deposition on the wide peritidal platform of the Elikah Formation. The MEBE (Middle East Basin Evolution) Programme funded the field work and part of the laboratories expenses. M. Ghassemi and A. Saidi of the GSI, Tehran, made easier the project and its development. Warm thanks are due to geologists and drivers of the Geological Survey of Iran, Tehran, who helped so much in the fieldwork. Part of the log drawings were carried out in Paris under the control of S. Crasquin. L. Trombino, Milano, made micrographies of the laterites and suggestions on their interpretation. C. Doglioni, Roma, suggested improvements on the geodynamic interpretation. S. Schro¨eder, Ko¨ln, made coral identifications. P. Brenckle, Maine, confirmed the Visean age of the microfauna of the Abnak locality. A. Zanchi, G. Muttoni, F. Berra and G. Gosso, Milano, discussed various aspects of the paper. F. Benigni, Milano, made part of the sandstone quantitative analysis. Reviews by M. Berberian, New Jersey, and C. Henderson, Calgary, greatly improved the original draft of the paper.

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PENNSYLVANIAN –EARLY TRIASSIC ALBORZ IRAN S ENGO¨ R , A. M. C. 1990. A new model for the late Palaeozoic-Mesozoic tectonic evolution of Iran and Implication for Oman. In: R OBERTSON , A. H. F., S EARLE , M. P. & R IES , C. (eds) The Geology and Tectonics of the Oman Region. Geological Society, London, Special Publications, 49, 797–831. S EYED -E MAMI , K. 2003. Triassic in Iran. Facies, 48, 95– 106. S HAHRABI , M. 1999. Triassic of Iran. Geological Survey of Iran, Tehran (in Persian). S HAPIRO , R. S. & W EST , R. 1999. Late Paleozoic stromatolites: new insights from the Lower Permian of Kansas. Lethaia, 32, 131– 139. S HARLAND , P. R., A RCHER , R., C ASEY , D. M., D AVIES , S. H., S IMMONS , M. D. & S UCLIFF , O. E. 2004. Arabian plate sequence stratigraphy – revisions to SP2. GeoArabia, 9, (1), 199– 214. S HEN , S. Z., G RUNT , T. A & J IN , Y. G. 2004. A comparative study of Comelicaniidae Merla, 1930 (Brachiopoda: Athyridida) from the Lopingian (Late Permian) of South China and Transcaucasia in Azerbaijan and Iran. Journal of Paleontology, 78, 884–899. S HI , G. R. & C HEN , Z. Q. 2006. Lower Permian oncolites from South China: Implications for equatorial sea-level responses to Late Palaeozoic Gondwanan glaciation. Journal of Asian Earth Sciences, 26, 424–436. S HI , G. R. & W ATERHOUSE , J. B. 1996. Lower Permian brachiopods and molluscs from the upper Jungle Creek Formation, northern Yukon Territory, Canada. Geological Survey of Canada Bulletin, 424, 1–241. S TAMPFLI , G. 1978. Etude ge´ologique ge´ne´rale de l’Elbourz oriental au S de Gonbad-e-Qabus, Iran N-E. PhD thesis, Universite´ de Gene`ve. S TAMPFLI , G. & B OREL , G. D. 2002. A plate tectonic model for the Palaeozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth Planetary Science Letters, 196, 17–33. S TAMPFLI , G., Z ANINETTI , L., B RO¨ NNIMAN , P., J ENNY D ESHUSSES , C. & S TAMPFLI -V UILLE , B. 1976. Trias de l’Elburz oriental, Iran. Stratigraphie, se´dimentologie, micropale´ontologie. Rivista Italiana di Paleontologia e Stratigrafia, 82, 467– 500. S TEIGER , R. 1966. Die Geologie des West-Firuzkuh-Area (Zentralelburz/Iran). Mitteilungen Geologische Institut ETH, n.s., 68. S TEPANOV , D. L., G OLSHANI , F. & S TOCKLIN , J. 1969. Upper Permian and Permian– Triassic Boundary in North Iran. Reports, Geological Survey of Iran, 12. S TEPHENSON , M. H., A NGIOLINI , L. & L ENG , M. J. 2007. The early Permian fossil record of Gondwana and its relationship to deglaciation: a review. In: W ILLIAMS , M., H AYWOOD , A. M., G REGORY , F. J. & S CHMIDT , D. N. (eds) Deep-time Perpectives on Climate Change: Marrying the Signal From Computer Model and Biological Proxies. Micropalaeontological Society, Special Publications, 2, 103–122. S TEPHENSON , M. H., O STERLOFF , P. L. & F ILATOFF , J. 2003. Palynological biozonation of the Permian of Oman and Saudi Arabia: progress and challenges. GeoArabia, 8, 467–496.

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Lithostratigraphy of the Upper Triassic –Middle Jurassic Shemshak Group of Northern Iran ¨ RSICH1*, MARKUS WILMSEN1,2, KAZEM SEYED-EMAMI3 & FRANZ THEODOR FU MAHMOUD REZA MAJIDIFARD4 1

GeoZentrum Nordbayern der Universita¨t Erlangen-Nu¨rnberg, Fachgruppe Pala¨oUmwelt, Loewenichstrasse 28, D-91054 Erlangen, Germany (e-mail: [email protected])

2

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany 3

School of Mining Engineering, University College of Engineering, University of Tehran, P.O. Box 11365-4563, Tehran, Iran 4

Geological Survey of Iran, P.O. Box 131851-1494, Tehran, Iran

*Corresponding author (e-mail: [email protected]) Abstract: The Upper Triassic –lower Middle Jurassic Shemshak Group is a siliciclastic unit, up to 4000 m in thickness, which is widespread across the Iran Plate of northern and central Iran. The group is sandwiched between two major unconformities: the contact with the underlying platform carbonates of the Elikah and Shotori formations is characterized by karstification and bauxite– laterite deposits; the top represents a sharp change from siliciclastic rocks to rocks of a Middle– Upper Jurassic carbonate platform– basin system. In the Alborz Mountains, the group consists of a Triassic and a Jurassic unit, separated by an unconformity, which is in part angular in the northern part of the mountain range and less conspicuous towards the south. Published lithostratigraphic schemes are based on insufficient biostratigraphic and lithological information. Here we present a new lithostratigraphic scheme for the central and eastern Alborz Mountains modified and enlarged from an unpublished report produced in 1976. Two major facies belts, a northern and a southern belt running more or less parallel to the strike of the mountain chain, can be distinguished. In the north, the Triassic part of the group is composed of the comparatively deep-marine Ekrasar Formation with the Galanderud Member (new name) at the base followed by the Laleband Formation, which represents prodelta–delta front environments. Up-section, the latter is replaced by the fluvial–lacustrine, coal-bearing Kalariz Formation. The equivalent Triassic lithostratigraphic unit in the south is the Shahmirzad Formation, redefined here, with the Parvar Member at the base. The formation represents fluvial, coastal plain and shallow- to marginal-marine environments. In the north, the Jurassic part of the group consists exclusively of the Javaherdeh Formation, coarse conglomerates of alluvial fan –braided river origin, which towards the south grades into the Alasht Formation, rocks of fluvial–lacustrine origin with coal. Further south, the Alasht Formation represents intertonguing marginal-marine– flood-plain environments and is followed by the Shirindasht Formation, sandstones and siltstones, indicative of the storm-dominated shelf, and the Fillzamin Formation (new), which is characterized by comparatively deep-marine shales. In the south, the group ends with the Dansirit Formation of deltaic– coastal-plain origin. This lithostratigraphic scheme reflects the tectono-sedimentary evolution of the Shemshak Foreland Basin of the Alborz Mountains where, during the Late Triassic, a relict marine basin in the north became gradually infilled, whereas in the south non-sedimentation and subaerial erosion prevailed and sediments record largely non-marine– marginal-marine conditions. During the early Lias, the basin was filled with erosional debris of the rising Cimmerian Mountain Chain, deposited largely in non-marine environments. During the early Middle Jurassic, in contrast, rapid subsidence in the south resulted in the deepening and subsequent infilling of a marine basin.

The Late Triassic– Middle Jurassic Shemshak Group is one of the most extensive lithostratigraphic units of the Iran Plate. The latter is part of the Cimmerian terranes, sandwiched between the Turan Plate in the north as part of Eurasia and the

Zagros fold belt in the south as part of Gondwana (Fig. 1). The Iran Plate, composed of the central-east Iranian microcontinent, the Alborz Mountains and northwestern Iran, became detached from the northeastern margin of Gondwana during

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 129–160. DOI: 10.1144/SP312.6 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Outcrops of the Shemshak Group in the Alborz Mountains and locations of sections and points of investigations. 1, Jajarm area; 2, Tazareh section; 3, Shahmirzad section; 4, Parvar area; 5, Sharif Abad section; 6, Djam area; 7, Rian area; 8, Damavand area; 9, Shemshak type area; 10, Paland area; 11, Galanderud area; 12, Ekrasar area; 13, Javaherdeh area.

the Early Permian (Stampfli & Borel 2002) and moved northwards during the Triassic, thereby gradually closing the Palaeotethys Ocean (Fig. 2). It collided with Eurasia towards the end of the period (e.g. Stampfli & Borel 2002), although the precise age of collision is still under debate (e.g. Sengo¨r et al. 1988; Sengo¨r 1990; Alavi et al. 1997; Saidi et al. 1997; Seyed-Emami 2003; Golonka 2004). A result of this so-called Early Cimmerian orogeny was the uplift of the Cimmerian mountain chain situated along the suture zone of the Palaeotethys. Denudation of this mountain chain produced large quantities of sediment that collected in an extensive foreland basin situated to the south. It is this rock unit that has been termed the Shemshak Formation and Shemshak Group, respectively, by various authors. In most areas it rests unconformably on carbonate platform sediments of the Elikah Formation, but locally it also lies on older beds (e.g. Permian or even older rocks). The group is followed by another carbonate unit, the Dalichai –Lar basin-carbonate platform complex, from which it is separated by another unconformity, reflecting the so-called MidCimmerian tectonic event. Characteristic features of the Shemshak Group are highly variable thicknesses reaching up to 4000 m, a nearly exclusively siliciclastic nature, widespread coal beds and environments ranging from proximal alluvial fans to deep marine.

The Shemshak Formation was defined by Assereto (1966), based on the type area around Shemshak, north of Tehran. The type section is situated in the upper Ruteh Valley, the co-ordinates of the base of the section being N358590 4200 , E518320 0200 . Earlier on, a number of terms were used by various authors to describe this lithostratigraphic unit, such as Lower Lias coal-beds (Goeppert 1862), Schichten mit Pflanzenresten, z.T. Kohlenflo¨zchen (Fischer 1914, 1915) and Jurassic plant-bearing series (Tipper 1921) (for a complete list of synonymous terms see Assereto 1966). Owing to its enormous thickness in many areas, it has been suggested to elevate the formation to Group rank (e.g. Aghanabati 1998; Seyed-Emami 2003). Assereto (1966) and Assereto et al. (1968) distinguished four ‘lithozones’ within the formation (lower sandstone, lower carbonaceous series, upper sandstone and upper carbonaceous series: Fig. 3). Subsequently, the Shemshak Formation/Group was subdivided into various members/formations, and different lithostratigraphic schemes have been proposed (Fig. 3) (Bragin et al. 1976; Nabavi & Seyed-Emami 1977; Nabavi 1980; Repin 1987; Aghanabati 1998). An elaborate lithostratigraphic scheme, supplemented with biostratigraphic data, was put forward by Bragin et al. (1976), who distinguished several ‘suites’ and correlated a number of key sections. Nabavi & Seyed-Emami (1977) and Nabavi (1980) divided the Shemshak Formation

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into five members (Parvar (ferruginous member), Jaban (volcanic member), Tazareh, Shahmirzad and Diktash), provided stratotype information, but only very rudimentary information on the various type sections. Repin (1987, fig. 30) subdivided the Shemshak Group into three formations: Tazareh, Ganu and Javaherdeh, which are composed of members based on the ‘suites’ of Bragin et al. (1976). The schemes of Bragin et al. (1976), Nabavi (1980) and Repin (1987) are all unpublished reports. Aghanabati (1998) summarized the existing information on the Shemshak Group. He elevated, without any discussion, the five members of Nabavi (1980) to Formation rank, and accepted the proposed subdivisions and descriptions of Bragin et al. (1976) and Repin (1987) for the Alborz area without, however, providing much additional information. Thus, all available schemes are either based on insufficient lithological, lithostratigraphic and biostratigraphic information, and/or are unpublished. Therefore, the purpose of this paper is to propose a new, comprehensive lithostratigraphic scheme for one of the key areas of the Shemshak Group, the Alborz Mountains, and to correlate key sections. Furthermore, the relationship of this scheme to other occurrences of the Shemshak Group, in particular of east-central Iran, is discussed. The scheme is then used to briefly trace the sedimentary and geodynamic evolution of the foreland basin.

Sections

Fig. 2. Palaeogeographic sketches illustrating the movement of the Cimmerian Continent Collage towards Eurasia since the Permian (modified after Stampfli & Borel 2002).

Extensive fieldwork in the Alborz area in the years 2003– 2006 focused on detailed logging of sections, and recording sedimentological, ichnological, palaeoecological and, wherever possible, biostratigraphic information. The field area comprised the eastern and central Alborz and both, northern and southern, areas of the mountain range. Previous studies (e.g. Bragin et al. 1976; Nabavi 1980; Repin 1987; Vollmer 1987; Fu¨rsich et al. 2005) have shown that the Shemshak Group is arranged in facies belts that run roughly parallel to the present-day Alborz Mountains, i.e. in an east – west direction. In contrast, relatively rapid facies changes occur in the north–south direction. Conventionally, three facies zones (southern, central and northern) are distinguished, whereby in the central zone the Triassic part of the group is strongly reduced and the group rests partly on rocks such as Devonian, Carboniferous or Permian in age (e.g. Allenbach 1966; Assereto 1966; Nabavi 1980; Vollmer 1987). The section localities are given in Fig. 1. Complete sections through the Shemshak Group were measured at Golbini, Tazareh, Shahmirzad and west of Damavand. Shorter segments were measured, particularly

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Fig. 3. Existing lithostratigraphic schemes of the Shemshak Group and their relationship to the scheme proposed here.

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in the northern Alborz Range, where outcrop conditions are less favourable and complete sections were difficult to obtain. Furthermore, critical intervals, such as the lower and upper boundaries of the group or the contacts between formations, were studied at additional localities (see the Appendix).

Biostratigraphic framework For a long time only a very rudimentary biostratigraphic framework existed for the Shemshak Group. This was partly owing to the non-marine nature of large parts of the group, partly also to lack of detailed investigations. Plant macrofossils and, to a certain degree, palynomorphs have extensively been used to date the non-marine segments of the group (Assereto et al. 1968; Kilpper 1975; Bragin et al. 1976; Sadovnikov 1976; Corsin & Stampfli 1977; Fakhr 1977; Schweitzer 1978; Achilles et al. 1984; Schweitzer et al. 1987, 2000; Schweitzer & Kirchner 1995, 1996, 1998, 2003; Vaez-Javadi & Ghavidel Syooki 2002; Vaez-Javadi 2006). Bragin et al. (1976) recognized three floral complexes characteristic of the Norian –Rhaetian (I), Hettangian– Pliensbachian (II) and Middle Jurassic (III), and used them to correlate their lithostratigraphic units. Dinoflagellate cysts have been used to date the basal part of the Shemshak Group at Galanderud as Norian –Rhaetian (Ghasemi-Nejad et al. 2004) by recognizing the Rhaetogonyaulax wigginsii and Rhaetogonyaulax rhaetica biozones. In the course of the present study, numerous ammonites have been collected from the marine Jurassic parts of the succession, in particular the upper part of the Shemshak Group, and have been documented in several papers (Seyed-Emami et al. 2005, 2006, 2008; Seyed-Emami & Hosseinzadeh 2006). Earlier studies on the ammonite faunas have been carried out by Nabavi & Seyed-Emami (1977), Seyed-Emami & Nabavi (1985) and Seyed-Emami (1987). As the ammonite fauna is identical with that of northwestern Europe, the zonal scheme from that area can be applied to the Shemshak Group. Based on ammonites, sections of the marine upper parts of the Shemshak Group can be correlated across the Alborz (e.g. Seyed-Emami et al. 2005, 2006). Earliest forms (Vermiceras sp., Arnioceras mendax rariplicata Fucini (¼Paltechioceras cf. oosteri (Dumortier)) document the Late Sinemurian (Nabavi & Seyed-Emami 1977), while youngest forms (e.g. Stemmatoceras sp. ?) indicate a Bajocian age (Seyed-Emami et al. 2008). The time span Toarcian –Aalenian is well represented by ammonites, and strata of this age can be correlated in detail. Unfortunately, this is not the case in the marine

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parts of the Upper Triassic strata. Ammonites of (Middle) Norian age have been found in the basal part of the section at Golbini (a new genus still awaiting description) and Rezaabad (Seyed-Emami & Wilmsen, 2007), and ammonites of Carnian – Norian age such as Arcestes (?Proarcestes) sp., ?Arcestes (Ptycharcestes) sp., Thisbites cf. agricolae, Stikinoceras cf. kerri, Griesbachites cf. pseudomedleyanus, Griesbachites cf. himalayanus, Juvavites cf. multicostatus and Dimorphites sp. have been collected from the basal part of the Shemshak Group at Ekrasar and Galanderud in the northern Alborz (Seyed-Emami et al. 2009). Again, this material has not yet been studied in detail. The scarcity of ammonites in the more marginalmarine parts of the Triassic succession was partly compensated for by bivalves such as Inoperna (Triasoperna) schafhaeutli (Stur) from the Shahmirzad Formation at Shahmirzad and Trigonia (Modestella) sp. from the Laleband Formation at Galanderud, both indicative of the Norian–Rhaetian. The occurrence of various species of Halobia such as H. superba and H. rugosa (e.g. Seyed-Emami 2003) in the fully marine Ekrasar Formation at Galanderud supports a Carnian–Norian age of the formation (see also Repin 1987; Vollmer 1987). Similarly, the bivalve assemblage consisting of Inoperna (Triasoperna) schafhaeutli (Stur), Chlamys cf. favrii (Stoppani), Palaeocardita iranica Hautmann, Costatoria weishanensis Guo and Mesosaccella subzelima (Krumbeck), which co-occur with the new ammonoid taxon mentioned above, corroborate the Norian–Rhaetian age of the basal Shemshak Group at Jajarm (Golbini). Liassic bivalves have been described by Fantini Sestini (1966). Figure 4 denotes the chronostratigraphic levels for which biostratigraphic data are available and the fossil groups on which the data are based.

Lithostratigraphy of the Shemshak Group As already hinted at by Assereto (1966, p. 1166), and later taken up by Repin (1987, p. 201), Aghanabati (1998, p. 19) and Seyed-Emami (2003), the Shemshak Formation of the Alborz Mountains is herewith formally elevated to Group rank. This way the group corresponds to the age-equivalent Shemshak Group of east-central Iran, which comprises the Upper Triassic Nayband Formation, the Lower Jurassic Ab-e-Haji and Badamu formations, and the Middle Jurassic Hojedk Formation (e.g. Aghanabati 1998; Wilmsen et al. 2003). A visit to the upper Ruteh Valley revealed that the section designated by Assereto (1966) as type section is poorly exposed, not complete and not as fossiliferous as several other sections. We therefore propose the well-exposed and expanded Tazareh

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section as hypostratotype (reference section; see figures in the subsection on ‘Kalariz Formation (Rhaetian)’ later in this paper). The subdivision of the Shemshak Group presented below (see also the figures in the next subsection) is based on our fieldwork supplemented by unpublished information gained by Russian geologists during their extensive field studies in the 1970s (Bragin et al. 1976; Repin 1987). Owing to limited time, not all outcrops surveyed by the Russian scientists could be revisited so that information on type sections is, in some cases, based on their unpublished data. Where this is the case, appropriate reference is made.

Ekrasar Formation (Upper Carnian – Norian) The Ekrasar Formation corresponds to the Ekrasar Suite of Bragin et al. (1976). Information on the formation is given in Cartier (1971), Repin (1987), Vollmer (1987), Dabiri (2001), Seyed-Emami (2003) and Ghasemi-Nejad et al. (2004). Vahdati Daneshmand (1982) introduced the name Paland Formation, a synonym, for the succession at Paland. Type section. Slope east of Ekrasar village, designated by Bragin et al. (1976) (Fig. 5). Additional sections. A section through most of the formation was measured at Galanderud (N368250 5800 , E518530 5100 ) (Fig. 6), and the upper part of the formation was recorded at Paland (N368010 0300 , E528530 4400 ), where the transition to the overlying Laleband Formation is well exposed.

Fig. 4. Chronostratigraphic framework of the Shemshak Group and supporting biostratigraphic evidence. 1, Seyed-Emami et al. (2009). 2, Vollmer (1987), Repin (1987), Seyed-Emami (2003). 3, Seyed-Emami (2003), Seyed-Emami & Wilmsen (2007). 4, Norian– Rhaetian bivalves in the Shahmirzad and Jajarm sections (this paper). 5, Bragin et al. (1976), Fakhr (1977), Schweitzer (1978), Achilles et al. (1984), Schweitzer et al. (1987, 2000), Schweitzer & Kirchner (1995, 1996, 1998, 2003), Vaez-Javadi & Ghavidel Syooki (2002), Vaez-Javadi (2006). 6, Ghasemi-Nejad et al. (2004). 7, Fantini Sestini (1966). 8, Nabavi & Seyed-Emami (1977); Seyed-Emami & Nabavi (1985), Seyed-Emami (1987), Seyed-Emami et al. (2005, 2006, 2008), Seyed-Emami & Hosseinzadeh (2006). 9, Majidifard (2004). 10, Bragin et al. (1976), Repin (1987). Absolute ages after Gradstein et al. (2004). Abbreviations: Ca, Carnian; No, Norian; Rh, Rhaetian; He, Hettangian; Si, Sinemurian; Plb, Pliensbachian; To, Toarcian; Aa, Aalenian; Baj, Bajocian; Bat, Bathonian.

Thickness. 1002 m (at type locality), thinning to zero in the southern Alborz Range. At Galanderud, the formation is at least 550 m thick. At Paland, the formation appears to have thinned to approximately 200 m, and according to Vahdati Daneshmand (1982) to 300 m. Lithology. The formation consists predominantly of monotonous, dark-grey argillaceous siltstone and siltstone, locally also of blackish shaly argillaceous siltstone with much pyrite (Fig. 7B). Occasionally, thin siltstone and fine-grained ripple- or parallel-bedded sandstone beds are intercalated, arranged in thickening-upwards parasequences. In the basal part of the formation, well-bedded, silty micritic limestones are locally developed in the northern Alborz (e.g. at Galanderud, Ekrasar: Fig. 8). They are usually reddish, the colour stemming apparently from bauxitic clays reworked from the subaerially exposed Elikah platform further south. From the Doab region (on the road to Sari) Vahdati Daneshmand (1982) recorded an 80 mthick succession of mudstones, gypsum, dolomites and lateritic beds, which overlie the Elikah

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Fig. 5. Type section of the Ekrasar Formation. The highly schematic section is based on information from Bragin et al. (1976). Key of symbols applies for all figures.

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Fig. 6. Section through the Ekrasar Formation, stratotype of the Galanderud Member at Galanderud. For the key of symbols see Figure 5.

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Formation. This evaporitic unit is followed by approximately 200 m of basic –andesitic volcanics and finally by siliciclastic rocks of typical Shemshak character (see also Dabiri 2001). The siliciclastic unit is about 200 m thick and contains dinoflagellates, among them Rhaetogonyaulax rhaetica (Sarjent), which indicates the Upper Norian. Clearly, the evaporitic succession merits recognition as a separate member, but as we have not seen the succession, and as its relationship to the Elikah Formation is not clear to us, further information is necessary before a new formal lithostratigraphic unit can be erected. Boundaries. The boundary with the underlying Elikah Formation is conformable, but distinct. It is characterized by a sharp change from the whitish structureless dolostones of the Elikah Formation to soft, red-brown argillaceous silt, which may represent altered volcanic ash or bauxitic silty clay (Figs 8 and 9D –F). The boundary with the overlying Laleband Formation is gradational. It is drawn at the change from siltstone-dominated- to sandstone-dominated-facies. Fossil content. The basal limestones contain a low-diversity fauna of bivalves (e.g. Cassianella, Halobia), and ammonoids (e.g. Arcestes (?Proarcestes), ?Arcestes (Ptycharcestes), Thisbites, Stikinoceras, Griesbachites, and Dimorphites; Bragin et al. 1976; Seyed-Emami et al. 2009), indicating a Carnian –Norian age. Halobia is also found, partly abundantly, in the lower– middle parts of the silt –argillaceous silt unit, associated with rare specimens of the nautiloid Grypoceras sp., whereas the upper part is largely unfossiliferous. Age. Based on the ammonoids and Halobia, a late Carnian –Norian age can be assigned to the formation. A detailed study of the ammonoids surely will produce a more precise age for the base of the Ekrasar Formation. Based on dinoflagellates a Norian –Rhaetian age has been suggested by Ghasemi-Nejad et al. (2004). Subdivisions. In contrast to the predominantly finegrained siliciclastic nature of most of the Ekrasar Formation, the basal part is, in some areas (e.g. Galanderud, Ekrasar), developed as impure carbonates, which form a discrete unit. For this reason these carbonates are recognized here as a formal member within the formation. Galanderud Member, new Type section. Road-cut south of Galanderud village (N368250 5800 , E518530 5100 : Fig. 6). Thickness. 87 m at the type section, approximately 30 m at Ekrasar, 20 m at Paland.

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Lithology. The member is very heterolithic, consisting of argillaceous silt, silty marlstone and red-brown silty–fine-sandy bio-wackestones. Limestones are the characteristic facies type and separate the member from other parts of the Ekrasar Formation (Fig. 7A). The member may also contain intercalations of volcaniclastic sediments and basic volcanic rocks, and commonly exhibits at its base red clays and silt as well as beds of iron pisoliths. Fossil content. Fossils are relatively common. Apart from ammonoids (see above), bivalves, gastropods, brachiopods, the trace fossil Zoophycos and wood pieces occur. Marine fossils occur even in the red beds. Palaeoenvironmental interpretation of the Ekrasar Formation. The Ekrasar Formation is restricted to the northern part of the Alborz Mountains, diminishing in thickness towards the south. It represents fully marine, comparatively deep-water environments, for the most part well below storm-wave base. The open marine character is shown also by the main faunal components, ammonoids and the bivalve Halobia. Low-energy conditions are also supported by the widespread occurrence of Zoophycos in the Galanderud Member at Galanderud. Although the red, bauxitic nature of the lowermost sediments seems to point to subaerial weathering, fossils in these sediments clearly demonstrate their marine character. Rather, the bauxitic sediments were introduced into the basin from subaerially exposed parts of the Elikah platform further south. Within the Ekrasar Formation, the distinct facies change from the carbonate-dominated, somewhat condensed sediments of the Galanderud Member to the monotonous, exclusively fine-grained siliciclastic sediments of the remainder of the formation most probably reflects a change in the sediment source from south to north. Subsidence at least kept pace with sedimentation. The thick deposits most probably represent a distal signal of the initial collision of the Iran Plate with the margin of Eurasia and of the concomitant nappe formation north of the Shemshak Basin. The lack of any faunal elements in the formation at Paland raises doubts on its marine character. However, the black, pyrite-rich, partly shaly, argillaceous silts (Fig. 7B) point to anoxic bottom conditions, which may explain the lack of any benthic organisms. Early diagenetic dissolution of calcareous shells in these environments may, in addition, have prevented the preservation of nektic elements such as ammonoids. Basal parts of the Shemshak Formation in the southern Alborz Mountains are also partly

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Fig. 7.

Continued.

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shallowed towards the south where it had a marginal-marine–non-marine character, or whether several smaller basins existed, some of which had no or only a restricted connection with the open sea.

Laleband Formation (Upper Norian–Rhaetian) The Laleband Formation corresponds to the Laleband Suite of Bragin et al. (1976). The formation is only developed in the northern Alborz Range. According to Repin (1987), the formation is also developed in the central and southern Alborz Mountains, in some areas reaching thicknesses of more than 1000 m. However, we have not seen the characteristic facies of the formation in these areas. What Repin recognized as Laleband in the central and southern Alborz Mountains was apparently included by us in the Shahmirzad Formation. Type section. 15 km SE of Laleband village (Karmozd coal mine, north of Damghan) (Fig. 10). Thickness. 435 m at type locality (Bragin et al. 1976). Lithology. Regular alternations of bedded sandstones and argillaceous siltstones (Fig. 7C). The sandstones are mainly fine- to medium-grained, commonly sharp-based and graded. In some cases the sandstones have erosional bases. Sedimentary structures include planar lamination, ripple bedding, in some cases also large-scale trough cross-bedding. Plant debris is common. Fig. 8. Sketch of the basal Ekrasar Formation at Ekrasar (Galanderud Member). For the key of symbols see Figure 5.

characterized by volcanic, bauxite and limestone intercalations, and by thick successions of darkgrey–blackish argillaceous silts. The associated rare fauna, however, suggests a brackish-water – freshwater (lake) character of these strata, which are described below as the Parvar Member of the Shahmirzad Formation. At present it is not clear whether at that time interval the Shemshak Basin was a single, large sedimentary basin, which

Boundaries. The boundary to the underlying Ekrasar Formation is gradational. It is drawn at the first thick sandstone intercalation. The boundary with the overlying Kalariz Formation is sharp and defined, with the first occurrence of coal or channel sandstone of fluvial origin. Fossil content and age. Fossils are very rare and consist predominantly of plant fragments and plant debris (e.g. Equisetites, Podozamites). Rare marine trace fossils (Teichichnus, Thalassinoides) have been recorded from the Galanderud section. At the same locality thin shell beds with bivalves (e.g. Trigonia (Modestella) sp.) occur and are

Fig. 7. (Continued) (A) Galanderud Member (Carnian– ?Lower Norian) of the Ekrasar Formation (Carnian–Norian) south of Galanderud. (B) Black shales of the Ekrasar Formation (Carnian– Norian) at Paland. Marker pen is 140 mm long. (C) Alternation of (turbiditic) sandstones and dark shales of the Norian (– ?Lower Rhaetian) Laleband Formation at Paland. (D) Lacustrine carbonaceous shales and bedded sandstones with tree stems in growth position of the Rhaetian Kalariz Formation, Tazareh section. (E) Roots of calamitids at the base of a sandstone bed of the Rhaetian Kalariz Formation, Tazareh section. Marker pen is 140 mm long. (F) Cliff-forming conglomerates of the Lower Jurassic Javaherdeh Formation overlying soft and plant-covered sandstones and shales of the coal-bearing Kalariz Formation (Rhaetian) at Galanderud. (G) Angular unconformity betwen the soft Kalariz Formation (Rhaetian) and the Lower Jurassic conglomerates of the Javaherdeh Formation SW of Javaherdeh. (H) Polymict conglomerate of the Javaherdeh Formation SW of Javaherdeh.

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Fig. 9.

Continued.

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evidence of a Late Triassic (Norian –possibly Rhaetian) age of the formation. Palaeoenvironmental interpretation. The Laleband Formation reflects the gradual infilling of the marine –lacustrine basin(s) represented by the Ekrasar Formation. The sandstone–argillaceous siltstone alternations record deposition by density currents and can be interpreted as prodelta turbidites, which in the case of the Galanderud section shallow upwards into shoreface deposits. The distinct change to a fluvial regime above most probably corresponds to a tectonically forced sequence boundary, caused by uplift. At Paland, the formation contains neither marine fossils nor trace fossils. The complete lack of trace fossils in turbiditic sandstones is unusual considering their high preservation potential in this kind of environment. One explanation might be the continuing anoxic nature of the bottom waters, another one that, at Paland, we are dealing not with a marine basin fill, but with the fill of a timeequivalent lacustrine basin.

Shahmirzad Formation (Norian – Rhaetian) The Shahmirzad Formation corresponds to parts of the Kalariz Suite of Bragin et al. (1976), partly to the Tazareh Member of Nabavi (1980), and, in the southern facies belt, to the Ekrasar Suite of Bragin et al. (1976). It does not correspond to the Shahmirzad Member of Nabavi & Seyed-Emami (1977) and Nabavi (1980), nor to the Shahmirzad Formation of Aghanabati (1998) (¼ more or less Shirindasht Formation as defined here). Type section. Mountain slope east of Shahmirzad (N358470 1500 , E538200 0700 : Fig. 11). Additional sections. Tazareh, Mobarak Abad Valley, Golbini, Parvar (see the Appendix). Fig. 10. Type section of the Laleband Formation. The highly schematic section is based on information from Bragin et al. (1976). For the key of symbols see Figure 5.

Thickness. 464 m (at type locality). In the Mobarak Abad Valley, the formation is 419 m thick, at

Fig. 9. (Continued) (A) The Shemshak Group west of Emamzadeh Hashem Pass in the Mobarak Valley. Overlying light-coloured Permian– Triassic carbonates (lower left), the Shemshak Group encompasses approximately 2000 m at this locality, consisting of the Shahmirzad, Alasht, Shirindasht, Fillzamin and Dansirit formations. (B) Soft-weathering Upper Triassic Shahmirzad Formation (left) of the lower Shemshak Group at Golbini (Jajarm area) unconformably overlying the Lower– Middle Triassic Elikah Formation (right). The contact is a bauxitic palaeokarst. (C) Palaeokarst topography filled with red bauxitic clays of the Parvar Member (?Carnian) at Golbini (Jajarm area). (D) Type section of the Parvar Member (?Carnian) south of Parvar, NE of Shahmirzad, overlying cliff-forming carbonates of the Lower– Middle Triassic Elikah Formation. (E) Parvar Member (?Carnian) at the Tazareh section consisting of reddish volcanic and volcaniclastic rocks overlying the Lower–Middle Triassic Elikah Formation (light-coloured, lower left). (F) Gradational contact of the Elikah Formation (left) and ?Carnian tuffitic breccia and shale of the lower Ekrasar Formation at Paland. (G) Elikah Formation (left) overlain by the Upper Triassic Shahmirzad Formation at Shahmirzad type section. The lower (dark) part corresponds to the Parvar Member (?Carnian). (H) Pahoehoe lava; Parvar Member (?Carnian) of the Upper Triassic Shahmirzad Formation at Shahmirzad type section.

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Fig. 11. Section of the Shemshak Group at Shahmirzad (left) and Kuhe Bashm-e-Dehsufian (right) with type locality of the Shahmirzad Formation. For the key of symbols see Figure 5.

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Fig. 11. Continued.

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Golbini 223 m and at Parvar 232 m. The greatest thickness is reached at Tazareh with 1250 m. Lithology. The Shahmirzad Formation is developed in most parts of the southern Alborz Mountains, and laterally interfingers with the Ekrasar and Kalariz formations. It is very heterolithic, with grain sizes ranging from clay to mediumgrained sand. Siliciclastic sediments predominate, but at the base volcanic and volcaniclastic rocks, bauxite and, rarely, carbonates occur (see Parvar Member later). At the type section, the lowermost 136 m are volcanic rocks (Fig. 9G) followed by argillaceous silt and silt, between which several m-thick, sharp-based, graded, large-scale trough cross-bedded sandstone packages are intercalated. Very common are several cm- to dm-thick ripple-laminated fine-grained sandstone intercalations. Hummocky cross-stratification and parallel lamination also occurs. In parts of the section the sediments are organized as coarseningupwards parasequences. Plant fragments are common, whereas wood pieces and wood logs are rare. At two levels carbonaceous argillaceous silt occurs. At Jajarm (Fig. 9B), the sandstones are mainly arkosic and some of the sandstone packages show lateral accretion. The succession is completely siliciclastic except for 1 m-thick carbonate unit with marine body and trace fossils. At Parvar and Tazareh, a thick intercalation of blackish argillaceous silt with scattered concretions is a characteristic feature in the lower part of the formation. At the latter locality, rootlet horizons occur repeatedly in the upper half of the formation. The Mobarak Abad Valley section, finally, has a more sandy character whereby the sandstones range from highly mature quartzarenites to immature, strongly lithic arenites. Thickening- and coarsening-upwards parasequences are well developed here. Boundaries. As the underlying Elikah Formation has a completely different lithology (dolostones or carbonate mudstones), the lower boundary is quite sharp and represents an unconformity. In the Mobarak Abad Valley and at Golbini (Fig. 9C) the relief of the Elikah Formation is highly irregular due to karstification. Both lower and upper boundaries are distinctly diachronous. Where the Shahmirzad Formation is followed by the Kalariz Formation, the boundary is drawn with the onset of coals (e.g. at Tazareh) and/or the change from marine to non-marine conditions (e.g. Mobarak Abad Valley). Where the formation is overlain by the Alasht Formation, the boundary is characterized by the change from silt –fine- and medium-grained sandstones to thick packages of coarse-grained sandstones (e.g. at Golbini). At the type section the upper boundary is drawn with the

base of the first of a number of more than 10 m-thick coarse-grained sandstone packages, which are attributed to the Alasht Formation. Fossil content. Trace fossils occur at several horizons and are represented by Ophiomorpha, Diplocraterion, Rhizocorallium, Lockeia, Taenidium, Parahaentzschelinia, Curvolithus and Gyrochorte (Fu¨rsich et al. 2006). The occurrence of the bivalve Inoperna (Triasoperna) in the upper part of the type section indicates a Late Triassic (Norian –Rhaetian) age. The microbial limestones at Tazareh contain ostracods and gastropods. The overlying volcaniclastic rocks and silty calcareous shales yield scattered euryhaline bivalves such as Bakevellia. The thin marine intercalation at Golbini contains an undescribed ammonoid and the bivalves Inoperna (Triasoperna) schafhaeutli (Stur), Chlamys cf. favrii (Stoppani), Palaeocardita iranica Hautmann, Costatoria weishanensis Guo and Mesosaccella subzelima (Krumbeck) (M. Hautmann, pers. comm. 2006). Age. Based on bivalve evidence, the Shahmirzad Formation spans the Norian– Rhaetian. Subdivisions. The bauxitic– lateritic (ferruginous) basal units of the Shemshak Formation in the southern Alborz Mountains have been accommodated by Nabavi & Seyed-Emami (1977) and Nabavi (1980) in the so-called Parvar Member. Simultaneously, the Jaban Member was erected for the volcanic deposits at the base of the Shemshak Group. As bauxitic sediments and volcanic rocks are closely associated and commonly intercalated it is not practical to have two different members, but to unite them in a single member, i.e. the Parvar Member. Parvar Member Type section. Slope east of the road to Parvar (N358590 2900 , E538290 4700 : Figs 9D and 12). Thickness. Highly variable, from 0 to more than 100 m. At the type section the member is 40 m thick, at Tazareh 80 m, at Golbini 0 –8 m, at Mobarak Abad Valley 7 m and at Shahmirzad 136 m. Lithology. The member consists either of volcanic (basalts, olivine-basalts) and volcaniclastic rocks, of lateritic –bauxitic sediments, or more rarely of both. Bauxite deposits prevail, for example, at Golbini (Fig. 9B, C), whereas at Shahmirzad (Fig. 9G), Mobarak Abad Valley, Parvar and Tazareh (Fig. 9E) volcanic rocks are exclusively present. The volcanic rocks are partly olivegreen –blackish basaltic lava, which often exhibits

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an amygdaloidal texture. Red-brown –beige tuffs and lapilli tuffs are also widespread. At Tazareh, intercalations of light-grey weathering microbial limestones, partly dolomitized, and volcaniclastic sandstone with plant debris are common. At Parvar, the tuffs locally contain nodules of jaspis and some intercalations of tuffitic micrite and marlstones, with plant debris and a low-diversity mollusc fauna. The bauxite deposits usually fill a pronounced relief of the Elikah Formation (e.g. Fig. 9C) and therefore strongly vary in thickness over short distances. The red-brown bauxite is commonly associated with iron crusts. Fossil content. Low-diversity bivalve assemblages occur in the carbonate rocks of the member. They are dominated by Bakevellia, a taxon indicating a marine influence, although most probably only brackish conditions. The small microbial reefs at Tazareh are associated with ostracods and small gastropods.

Fig. 12. Type section of the Parvar Member of the Shahmirzad Formation. For the key of symbols see Figure 5.

Palaeoenvironmental interpretation of the Shahmirzad Formation. The Shahmirzad Formation represents environments frequently fluctuating between marginal-marine (nearshore shallow shelf) and fluvial. Evidence for non-marine conditions are rootlet horizons, rare Calamites stems in growth position and fining-upwards sandstone units with erosional bases (fills of fluvial channels). The tabular shape of these sandstone bodies suggests that they represent meandering river deposits developed on a coastal plain. Grey argillaceous silts, partly with ferruginous concretions, and commonly grading into siltstones and finally sandstones can be interpreted as the fills of small lakes that developed on the flood plain. The thick package of black argillaceous siltstones at Tazareh and Parvar most probably represent deep lakes, which gradually developed from marginal-marine environments that prevailed at the base of the formation. Fossil evidence points to subaqueous deposition of most of the volcaniclastic and tuffitic material. Most probably, no fully marine conditions existed in the basal part of the Shahmirzad Formation because the microbial reefs and the lowdiversity molluscan fauna, in particular Bakevellia, point to brackish conditions. Thus, it appears that initially marginal-marine conditions existed in the Shemshak Basin at Tazareh and Parvar, but that marine connections were severed so that the marine –brackish water body developed into a freshwater lake, which, through time, became filled with sediment. At Mobarak Abad Valley and Shahmirzad, marginal-marine conditions evidenced by trace fossils frequently alternated with fluvial conditions shown by root horizons and fining-upwards sandstone bodies. At Golbini,

Fig. 13. Section of the Shemshak Group at Tazareh coal mine area with type sections of the Kalariz, Shirindasht and Dansirit formations, and reference section of the Alasht Formation. This is the most expanded section of the Shemshak Group and is defined as a reference section. For the key of symbols see Figure 5.

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a non-marine coastal plain existed right from the beginning, and marine influence was restricted to a short-lived transgression, which is documented by a fully marine molluscan fauna.

Kalariz Formation (Rhaetian) The Kalariz Formation (‘suite’) was coined by Bragin et al. (1976), although, as redefined here, it attains only a third of the thickness given by these authors. The Tazareh Formation of Nabavi (1980) partly corresponds to the Kalariz Formation in the present definition. The Kalariz Formation extends across much of the Alborz Range, but is absent in the Shahmirzad area where it is replaced by the marginal-marine– coastal-plain sediments of the Shahmirzad Formation. The main coal deposits of the Shemshak Group in south and central Alborz come from this formation. Type section. At Tazareh coal mine, 15 km east of Kalariz village (N368240 3700 , E548290 1000 : Fig. 13). Additional section. Mobarak Abad Valley (N358 480 3900 , E518580 2200 : Fig. 9A). Thickness. 255 m at type locality, 131 m at Mobarak Valley. Lithology. Grey argillaceous –silty fine-grained sandstones with thin intercalations of ripple-bedded fine-grained sandstones, argillaceous silt with ferruginous concretions, carbonaceous silt and coal alternate with fine-grained, large-scale trough cross-bedded sandstones. Root horizons (Fig. 7D, E) and plant remains are common throughout the succession. Coarsening-upwards parasequences starting with partly interlayered argillaceous silt and ending with trough crossbedded sandstones are commonly developed. In the middle part of these cycles thin, parallellaminated or hummocky cross-stratified finegrained sandstones are often found. Boundaries. At Tazareh, the lower boundary of the formation is defined at the top of the last thick sandstone unit of the Shahmirzad Formation, at Mobarak Abad Valley at the change from marine sandstones of the Shahmirzad Formation to fluvial–lacustrine silts and sandstones with coal. A transgressive lag marks the boundary to the overlying Alasht Formation at Mobarak Valley, whereas at Tazareh the upper boundary is formed by thick, coarse-grained sandstone packages of the Alasht Formation. Fossil content. Plants and plant fragments are highly abundant, especially in the fine-grained

sediments (e.g. Neocalamites, Baieria, Nilssonia, Nilssoniopteris Pterophyllum, Thainguyenopteris, Podozamites, Otozamites). Neocalamites stems occasionally occur in growth position (Fig. 7D, E). Root horizons are widespread. Apart from ubiquitous plant fragments and more complete plant fossils at some horizons, no fossils have been found, except rare trace fossils (Planolites, Skolithos). Age. Based on the rich flora, Bragin et al. (1976) assigned a Late Triassic age to the formation, here interpreted to correspond to the Rhaetian, as the formation is underlain by Norian strata. Palaeoenvironmental interpretation. The Kalariz Formation is a non-marine unit. Several m-thick fining-upwards sandstone bodies represent deposits of small fluvial channels that grade into finergrained levee and flood-plain sediments. The latter clearly dominate and suggest that we are dealing with a coastal plain that was dotted with lakes and swamps, and crossed by small rivers. Coarsening-upwards siltstone–sandstone successions are interpreted as marginal lacustrine deposits, which commonly are followed by coal or carbonaceous silts, forming in swampy areas of the lake margin and on the flood plain. Dark-grey, finegrained siliciclastic sediments represent deeper lake sediments.

Alasht Formation (Hettangian– Toarcian) The Alasht Formation (Alasht Suite of Bragin et al. 1976) corresponds to parts of the Tazareh Member of Nabavi (1980). In contrast to Bragin et al. (1976), the formation is thought to extend, as a narrow zone, into the Toarcian in the northernmost part of its outcrop belt. Type section. Left bank of Alasht River at Karmozd (Bragin et al. 1976) (Fig. 14). Additional sections (reference sections). Tazareh (N368250 500 , E548290 2500 : Fig. 13), Golbini (N378030 , E568320 ), Shahmirzad (N358470 , E538200 : Fig. 11) and Mobararak Abad Valley (N358480 3900 , E518580 2200 : Fig. 9A). Thickness. 735 m at type locality, 1693 m at Tazareh, 570 m at Golbini, 492 m at Shahmirzad and 309 m at Mobarak Abad Valley. The formation thus decreases in thickness towards the west. Lithology. The Alasht Formation is characterized by alternating packages of fine-grained and medium- to coarse-grained siliciclastics, separated by finer-grained interbeds. Sandstones are arkosic, fine- to coarse-grained, in some cases pebbly or conglomeratic, exhibit large-scale trough

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cross-bedding and have thicknesses of the order of 10 –30 m (Fig. 15B). The pebbles consist of quartz or are intraformational in origin (reworked concretions, rip-up clasts of clay and silt). The sandstones, which invariably exhibit a sharp, erosional base, generally fine upwards, and occasionally show lateral accretion. Tree trunks occasionally occur at the base of the sandstone bodies. At Shahmirzad and Mobarak Abad Valley, some of the fine-grained sandstones are very well sorted, exhibit tabular bedding or very-low-angle lamination, and are associated with marine trace fossils and shells. The finergrained units consist of alternations of thin, finegrained sandstones, siltstones and argillaceous silt with ripple lamination, interlayered bedding, rootlet horizons, and wood and plant fragments. Intercalations of carbonaceous silt and coal seams are common (Fig. 15C). Occasionally, coarseningupwards sequences are also developed. Pedogenic features are widespread, and include mottled siltstones and ferricretes. In the uppermost few tens of metres of the formation there is a distinct change in the colour of sandstones from grey to greenish-grey. Some of these sandstones are bioturbated, and exhibit remains of marine bivalves and trace fossils. In the two westernmost sections (Shahmirzad and Mobarak Abad Valley) the formation differs in that it partly exhibits a marginal-marine character, evidenced by trace fossils and levels with marine fauna. Coal beds are absent, but hummocky cross-stratification and layers of well-rounded sandstone pebbles associated with quartz granules, ammonites, belemnites, and bivalves are common. Repin (1987, p. 223) subdivided the formation (as defined by him; see Fig. 3) into a lower sandy– silty member (Assiab) and an upper, more fine-grained member with coal seams (Pashkalat). According to Repin, the central facies zone disappears towards the east so that the southern and the northern facies belt interfinger from the Jajarm area eastwards.

Fig. 14. Type section of the Alasht Formation. The highly schematic section is based on information from Bragin et al. (1976). For the key of symbols see Figure 5.

Boundaries. The base of the formation is sharp, characterized by the onset of thick packages of coarse-grained sandstones (up to 70 m at the type locality) (Fig. 15A). The top is gradational, and the boundary to the overlying Shirindasht Formation is drawn at the top of the last thick fining-upwards sandstone body. The last coal seams and rootlet horizons usually occur several metres below the boundary. At Kuhe Bashm-eDehsufian and Mobarak Abad Valley, the upper boundary is characterized by a change to finergrained sediment and a distinct increase in the degree to which the sediments are bioturbated. Laterally, the Alasht Formation interfingers towards the south with the Shirindasht Formation and to the north with the Javaherdeh Formation.

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Fig. 15. Continued.

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Fossil content. Plant fossils are ubiquitous (e.g. Nilssonia, Nilssoniopteris, Otozamites, Ginkgo, Podozamites, Thaumatopteris, Swedenborgia, Czekanowskia). In the eastern sections invertebrate body fossils are absent, except in the top part of the formation where bivalve shells such as Protocardia occur at several horizons. Trace fossils are rare throughout. Diplocraterion parallelum has been found near the top of the formation at Tazareh, and the bivalve resting trace Lockeia and a questionable Thalassinoides have been encountered in the middle part of the formation at Tazareh. Towards the west, fossils become far more common, but tend to be concentrated in levels. Apart from rare corals, belemnites, ammonites (e.g. Paltechioceras cf. oosteri, Oregonites cf. imlayi), gastropods, serpulids and bivalves occur at Shahmirzad and Mobarak Abad Valley. Trace fossils are common. In the well-sorted, low-angle cross-bedded sandstones at Shahmirzad, Ophiomorpha, Diplocraterion, Parahaentzschelinia, Gyrochorte, and Taenidium occur (Fu¨rsich et al. 2006). Additional ichnotaxa are Ophiomorpha, Rhizocorallium irregulare, Jamesonichnites and Thalassinoides.

of which were heavily vegetated and often formed swamps. Towards the top of the formation a coastal plain with rivers and small lakes gradually was transgressed by the sea, whereby several short-lived marine ingressions, evidenced by marine bivalves (Protocardia) and trace fossils (Diplocraterion), established marginal-marine to ?estuarine conditions. In the west the marine influence is much stronger, and both body and trace fossils show that shallow- to marginal-marine conditions frequently alternated with non-marine conditions. In the upper half of the formation the former exclusively prevailed. The well-sorted fine-grained sandstone packages with parallel–low-angle lamination are interpreted as beach deposits (Fu¨rsich et al. 2006), an interpretation that is supported by the characteristic trace fossil suite. The repeatedly occurring thin layers of sandstone pebbles and associated shells, including ammonites, are best interpreted as lags produced by transgressive pulses. Lower delta-front sandstones form a 100 m-thick package at the top of the formation at Kuhe Bashm-e-Dehsufian, and illustrate the great facies variation across short distances.

Age. Based on the Triassic character of the plant association of the underlying Kalariz Formation and on the Toarcian ammonites occurring at the base of the overlying Shirindasht Formation, the Alasht Formation is assigned an Early Liassic age (Hettangian/Sinemurian–Pliensbachian, most probably also Toarcian). This is confirmed by Late Sinemurian– Pliensbachian ammonites recorded at Shahmirzad (Seyed-Emami & Hosseinzadeh 2006; Seyed-Emami et al. 2008).

Shirindasht Formation (Lower Jurassic –Aalenian)

Palaeoenvironmental interpretation. The Alasht Formation represents mainly fluvial environments. Fining-upwards sandstones with lateral accretion point to meandering river systems, rootlets and soil horizons (mottled ferruginous levels and iron crusts) to flood-plain environments. On the flood plain, small lakes developed, the marginal zones

The Shirindasht Formation, as defined here, corresponds to the Anan suite and the lower part of the Shirindasht suite of Bragin et al. (1976), and to the upper part of the Tazareh Formation and the Shahmirzad Formation of Nabavi (1980). Type section. Left slope of Shirindasht gorge, Tazareh area (N368250 1700 , E548280 1000 : Fig. 13). Additional sections. Sharif-Abad (N358420 5900 , E538260 0100 ), Kuhe Bashm-e-Dehsufian (N358510 0300 , E538260 5900 ; Fig. 11), Golbini (N378030 , E568 310 ), Tappehe Chefte Qazi (N368540 5700 , E568160 0200 ) and Mobarak Abad Valley (N358490 0400 , E518570 4700 ).

Fig. 15. (Continued) (A) Sharp contact of the soft-weathering, fine-grained ?upper Norian –Rhaetian Kalariz Formation (right) and the cliff-forming coarse and thick-bedded sandstones of the Lower Jurassic Alasht Formation in the Tazareh section. (B) Sharp-based, approximately 20 m-thick fluvial channel sandstone of the Lower Jurassic Alasht Formation in the Golbini section. (C) Lower Jurassic coal seam of workable thickness in the Alasht Formation of the Tazareh section. (D) Bedded sandstone and shale of the Toarcian– Lower Aalenian Shirindasht Formation, Tazareh section. (E) Alternation of hummocky cross-bedded sandstones (sandy tempestites) and siltstones (background sediments) of the Toarcian–Lower Aalenian Shirindasht Formation at Golbini (Jajarm area). (F) Soft-weathering, slightly marly Fillzamin Formation (Middle Aalenian–Lower Bajocian) at Sharif-Abad (east of Diktash) overlain (in the background) by sandstones of the Lower Bajocian Dansirit Formation. (G) Dark mudstones of the Fillzamin Formation (Middle Aalenian– Lower Bajocian) overlain by deltaic sandstones of the Lower Bajocian Dansirit Formation at Fillzamin, Ira valley. (H) Package of deltaic sandstones of the Lower Bajocian Dansirit Formation with workable coal seams at Fillzamin, Ira Valley, sandwiched between Aalenian–Lower Bajocian mudstones of the Fillzamin Formation (below) and marls of the Upper Bajocian–Bathonian Dalichai Formation (above, covered by scree of the cliff-forming limestones of the Callovian– Upper Jurassic Lar Formation).

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Thickness. 551 m at the type locality, 365 m at Golbini, 111 m at Sharif-Abad, 238 m at de Sufian Bashm and 440 m at the Mobarak Abad Valley. Lithology. At the type locality, the Shirindasht Formation is characterized by small-scale alternations of heavily bioturbated dark-grey–olive-grey silt– fine-sandy silt and fine-grained sandstones (Fig. 15D). The sandstones exhibit a sharp base, planar lamination, ripple bedding or hummocky cross-stratification. Layers with intraformational pebbles and shell concentrations rarely occur. Shells also occur scattered throughout the section, as do wood fragments and comminuted plant debris. The rocks are organized in small coarsening-upwards parasequences, topped by sandstones with hummocky cross-stratification or with large-scale trough cross-stratification. The basal part of the formation is sandstone-dominated, and there is a gradual decrease in grain size towards the top, which is strongly bioturbated (Fu¨rsich et al. 2005). At Shahmirzad and Sharif-Abad, the formation is characterized by highly bioturbated argillaceous silt with scattered, often highly fossiliferous limestone concretions, which are occasionally reworked into pebble layers, and some intercalations of hummocky cross-stratified or parallel-laminated sandstones. At Golbini, the lithologies and facies architecture strongly resemble those of Tazareh, with highly bioturbated fine-sandy silt being commonly punctuated by hummocky cross-stratified fine-grained sandstone beds (Fig. 15E), and several layers with intraformational sandstone pebbles and shell concentrations (bivalves, belemnites, ammonites). At the Mobarak Abad Valley, predominantly highly bioturbated fine-grained sandstones and sandy siltstones with scattered bivalves and ammonites, and several pebble layers of reworked sandstones, dominate. Several large-scale trough cross-bedded sandstone packages, up to 15 m thick, are intercalated between the bioturbated sediments, which fine up-section into argillaceous silt with scattered concretions. Part of this succession exhibits well-developed parasequences. For the last 80 m, the grain size increases again and ripplelaminated, parallel-laminated and large-scale trough cross-bedded fine- to medium-grained sandstones are present. Wood fragments and plant debris are common throughout the formation. Boundaries. At Jajarm, Tazareh and Sharif-Abad, the lower boundary is placed at top of the last thick, cross-bedded fluvial sandstone package and a corresponding change to finer-grained, generally thin-bedded sediments with a high proportion of silt. At Kuhe Bashm-e-Dehsufian and the Mobarak Abad Valley, the lower boundary is drawn at the

distinct increase in bioturbation. The boundary to the overlying Fillzamin Formation is in all sections gradual and characterized by a change in grain size from sand to silt combined with a nearly complete lack of primary sedimentary structures. Fossil content. Plant debris and wood fragments are common throughout the formation. Trace fossils are abundant and include Rhizocorallium irregulare, Zoophycos, Thalassinoides, Planolites, Taenidium, Curvolithus and Chondrites. Benthic macroinvertebrates include bivalves (e.g. Entolium, Pleuromya, Pinna, Myophorella, Pseudomytiloides, Parainoceramus, Grammatodon), terebratulid brachiopods, gastropods, serpulids and crinoid ossicles (Fu¨rsich et al. 2005). Ammonites (e.g. Grammoceras, Podagrosites, Pseudogrammoceras, Dumortieria, Pleydellia, Leioceras; Seyed-Emami & Nabavi 1985; Seyed-Emami 1987; Repin 1987; Seyed-Emami et al. 2005, 2006) and belemnites are common at some horizons. Age. Based on the available ammonite evidence, the formation ranges from the Toarcian to the early Aalenian. Palaeoenvironmental interpretation. The Shirindasht Formation represents storm-dominated shelf deposits, as is shown by the abundance of parallel and hummocky cross-stratified sandstones. The decrease in grain size and in primary sedimentary structures, coupled with an increase in bioturbation up-section, indicates a general deepening of the basin (Fu¨rsich et al. 2005), except at the Mobarak Abad Valley. There, the coarsening-upwards cross-bedded sandstone package at the top of the formation indicates shallowing, and probably represents a delta-front environment. Shelly pebble layers represent transgressive lags caused by highfrequency fluctuations of relative sea level.

Fillzamin Formation (Aalenian– Lower Bajocian?), new The Fillzamin Formation, erected herein, corresponds to the upper part of the Shirindasht Suite of Bragin et al. (1976), and to the lower part of the Diktash Formation of Nabavi & Seyed-Emami (1977) and Nabavi (1980). It is introduced because the existing schemes overlooked the distinct facies changes occurring at the base and top of the formation, and because of the characteristic, very uniform, facies of the formation that is unique within the Jurassic part of the Shemshak Group. Moreover, the Diktash member sensu Nabavi (1980) incorrectly also includes the lower marls of the Upper Bajocian–Callovian Dalichai Formation.

LITHOSTRATIGRAPHY OF THE SHEMSHAK GROUP

Type section. Fillzamin near Ira valley, southcentral Alborz (N358500 2700 , E518520 5500 : Fig. 16). Additional sections. Kuhe Bashm-e-Dehsufian (N358510 , E538260 : Fig. 11), Sharif-Abad (N358 420 5900 , E538260 0100 : Fig. 15F), Tazareh (N368260 1500 , E548280 1100 : Fig. 13) and Golbini (N378030 5900 , E568 310 5700 ). Thickness. 417 m at the type locality, 560 m at Kuhe Bashm-e-Dehsufian, 141 m at Sharif Abad, 678 m at Tazareh and 643 m at Golbini. Lithology. At Ira Valley, Mobarak Abad Valley and Kuhe Bashm-e-Dehsufian, the formation is a monotonous succession of indurated, dark-grey– slightly greenish, highly bioturbated, partly finesandy, argillaceous silt with occasional limestone concretions in the lower part. At Sharif-Abad (Fig. 15F), the sediment is marly and ferruginous concretions are common in the lower part. At Parvar, scattered comminuted plant debris is found at several levels. At Tazareh and Golbini, the rocks are slightly sandier, except for a shaley middle part, 20–30 m thick, at Tazareh. At this locality, the upper 160 m of the formation contain numerous intercalations of thin, fine-grained sandstone beds that exhibit a sharp base and parallel-lamination, and right at the top also hummocky cross-stratification (Fig. 13). Several m-thick, trough cross-bedded, fine-grained sandstones occur at some levels. At Golbini, the sediment is argillaceous –fine-sandy silt with common ferruginous concretions and intercalations of thin, hummocky cross-stratified fine-grained sandstone beds. These beds occur occasionally in the lower third of the section, are absent from the middle part and become very abundant in the upper half. Generally, the rocks are composed of coarsening-upwards parasequences. The lower part of such cycles is usually highly bioturbated, whereas the top sandstone is characterized by hummocky cross-stratification. A distinct coarsening of the succession, partly represented by small coarsening-upwards cycles, hummocky cross-stratification, parallel or ripple lamination can be seen in all sections of the Fillzamin Formation, although the thickness of this facies may be only a few metres. Boundaries. The lower boundary of the formation is gradational, and characterized by a decrease in grain size combined with an increase in the degree of bioturbation. The contact to the overlying Dansirit Formation is relatively sharp at Kuhe Bashm-e-Dehsufian and Tazareh, at the remaining section it is somewhat gradational and has been drawn at the base of the thick, cross-bedded sandstone unit of the Dansirit Formation (Fig. 15G).

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Fossil content. The macrofauna is largely confined to concretions and consists of in situ concentrations of inoceramids (Parainoceramus, Pseudomytiloides), the crustacean Glyphea, bivalves such as Entolium and Pleuromya, and rare gastropods. Belemnites and ammonites (e.g. Ludwigia, Brasilia, Graphoceras, Planammatoceras, Euaptetoceras) occur at several levels (Repin 1987; Seyed-Emami 1987; Seyed-Emami et al. 2005, 2006). Age. The ammonites indicate an Aalenian –(Early) Bajocian age of the formation. Palaeoenvironmental interpretation. At the type section, Parvar, Tazareh and at Kuhe Bashme-Dehsufian, the Fillzamin Formation represents deep shelf environments well below storm-wave base except for the topmost part. The rate of sedimentation was low enough to allow the sediment to become thoroughly bioturbated. Towards the top, evidence of density currents prevails and indicate a gradual shallowing of the basin and a prodelta –lower delta-front regime. At Golbini, the basin appears to have been shallower, the influx of sand-sized particles greater and the upper half of the formation consists of sediments that repeatedly came under the influence of storm waves. At other localities (e.g. at Parvar) bioturbated silts are abruptly followed by delta-front– delta-top sediments of the Dansirit Formation.

Dansirit Formation (Lower Bajocian) The Dansirit Suite was erected by Bragin et al. (1976). It corresponds to the middle unit of the Diktash Formation of Nabavi & Seyed-Emami (1975) and Nabavi (1980), and partly to the Upper Coal Member of Assereto (1966). Repin (1987, p. 243) elevated it to the rank of a formation without any discussion. Type section. At Tazareh coal mine (Razmja area; N368260 3100 , E548290 4800 : Fig. 13). Additional sections. Golbini, Kuhe Bashm-eDehsufian (Fig. 11), Parvar (N368020 0100 , E538280 2400 ), Sharif-Abad (N358420 5900 , E538260 0100 ) and Fillzamin (N358500 4600 , E518 530 2800 : Fig. 15G, H). Thickness. Strongly varying; 319 m at Tazareh (type section), 188 m at Golbini, 261 m at Kuhe Bashm-e-Dehsufian, 171 m at Parvar, 10.5 m at Sharif-Abad and 108 m at Fillzamin. Lithology. The Dansirit Formation consists of large-scale cross-bedded sandstone packages, between which silt carbonaceous silt and coal layers are intercalated (Fig. 17H). The sandstones are fine- to coarse-grained and either coarsen or fine upwards. At some levels conglomeratic lenses

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Fig. 16. Type section of the Fillzamin Formation at Fillzamin near Ira Valley. For the key of symbols see Figure 5.

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Fig. 17. New lithostratigraphic scheme of the Shemshak Group of the Alborz Mountains. Absolute ages after Gradstein et al. (2004). Abbreviations: Ca, Carnian; No, Norian; Rh, Rhaetian; He, Hettangian; Si, Sinemurian; Plb, Pliensbachian; To, Toarcian; Aa, Aalenian; Baj, Bajocian; Bat, Bathonian.

consisting of well-rounded quartz pebbles occur. The sandstones are highly immature arkoses to very mature quartz arenites. The finer-grained units exhibit ripple lamination and interlayered bedding. Sandstone beds intercalated between silt commonly exhibit hummocky cross-stratification. The mature quartz arenites often show low-angle lamination. Plant debris, wood fragments and, in a few cases, large wood logs, as well as root horizons, are common. At Golbini the formation consists of a thick package of well-bedded fine-grained sandstones, the lower two-thirds of which are crossbedded, whereas the upper third is largely bioturbated. Further west, at Tappehe Chefteh Qazi, the

formation consists of an about 30 m-thick package of sandstone, the quartes quarter of which is finegrained and exhibits m-thick planar cross-bedding units alternating with well-bedded units. The remaining part is medium- to coarse-grained and planar-bedded throughout, whereby the foresets all dip in the same direction. At Sharif-Abad, the Dansirit Formation is very thin. A 0.5 m-thick level of reddish ferruginous sandstone occurs about 3 m above the base. It is followed by a 0.3–0.5 m-thick layer of quartz and sandstone pebbles, with large trigoniid and other bivalves belemnites and corals. The medium-grained cross-bedded sandstone above contains lenticles rich in bioclasts,

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reworked ferruginous concretions, bivalves and serpulids. West of the road, between Sangsar and Shahmirzad, milk quartz conglomerates, composed of three units 1, 7 and 9 m thick and fining upwards into sandstones, occur below the Dalichai Formation in a heavily tectonized section. Most probably they are part of the Dansirit Formation. Boundaries. Although the top of the underlying Fillzamin Formation is generally coarser than the bulk of that formation, the lower boundary of the Dansirit Formation is easily identified by a sudden increase in grain size, generally in combination with large-scale trough cross-stratification. The top of the formation is an unconformity and characterized by a distinct change from sandstone to silt and silty marl of the Dalichai Formation. At some localities, such as Tooy (N378110 4500 , E578040 3400 ), it is even an angular unconformity. Fossil content. In the connection with coal seams, a rich flora has been determined (Bragin et al. 1976; Schweizer & Kirchner 2003). Taxa include Otozamites, Ptilophyllum, Pterophyllum, Nilssoniopteris and Williamsonia. Invertebrate fossils are rare and consist mainly of marine bivalves (e.g. Isognomon, Myophorella, Pholadomya, Ceratomya and Bakevellia), rarely of corals. Marine trace fossils such as Ophiomorpha, Planolites, Chondrites, Dactyloidites, Thalassinoides, Lophoctenium, Taenidium and Rhizocorallium irregulare are more widespread. The top bed of the formation commonly represents a transgressive lag deposit with bivalves (e.g. oysters, Ctenostreon), brachiopods, belemnites and, rare, unidentifiable ammonites. Age. Finds of ammonites in the underlying Fillzamin Formation and the overlying Dalichai Formation confine the age of the Dansirit Formation to the Early Bajocian. Palaeoenvironmental interpretation. In most areas, the Dansirit Formation represents deltaplain–marginal marine-environments. Well-sorted fine-grained sandstones with low-angle lamination and the trace fossil Ophiomorpha are interpreted as beach deposits, whereas coal seams, carbonaceous silt, interlayered clay–silt units and small coarsening-upwards sequences are typical of delta-plain embayments, swamps and lakes. Fining-upwards sandstone packages with erosional base are interpreted as the fills of distributary channels. Hummocky cross-stratified sandstones suggest storm-influenced foreshore environments. As one would expect in such marginal environments, the facies pattern differs from this general picture at some localities. For example, at Golbini the formation is represented

by 190 m of cross-bedded fine-grained sandstone, which is interpreted as representing upper delta-front deposits, whereas at Tappehe-Chefte-Qazi (Jajarm area) well-sorted, medium- to coarse-grained, planar cross-bedded sandstones are best interpreted as a near-shore bar complex. The thin development of the formation at Sharif-Abad in connection with red beds probably indicates a predominance of terrestrial environments, which left only traces in the sedimentary record. The well-rounded milk quartz conglomerates south of Shahmirzad are deposits of braided rivers and document erosion of metamorphic rocks. On the whole, the Dansirit Formation is the sedimentary expression of pronounced shallowing of the Shemshak Basin. This shallowing was already initiated towards the end of the underlying Fillzamin Formation, as is indicated by the ubiquitous coarsening-upwards trend, but was not completed at the base of the Dansirit Formation. The base of the formation therefore corresponds to a distinct relative sea-level fall (Fu¨rsich et al. 2009).

Javaherdeh Formation (Lower Jurassic – Lower Bajocian) The Javaherdeh Formation was erected by Bragin et al. (1976) for a conglomeratic unit developed along the northern margin of the Alborz Mountains and best observed in the Ramsar area. Type section. 1.5 km NE of Javaherdeh village (N368510 , E508280 ). Additional sections. Sedimentological observations were made along the new road leading uphill to the radar station at Javaherdeh (N368510 3800 , E508280 18.700 – N368510 5500 , E508270 37.400 ) and along the road between Rahimabad and Qazvin (N368510 0500 , E508130 3000 ). Thickness. 1050 m at type locality 1.5 km NE of Javaherdeh village (Bragin et al. 1976), 800 m at Ekrasar and Geshlag (Repin 1987), decreasing further towards the south. As defined here, the formation is less than 100 m at Galanderud. Lithology. Polymict conglomerates consisting mainly of well-rounded pebbles of quartzites, milk quartz, volcanics and metamorphics, rarely of granite (Fig. 7H). Pebble size reaches up to 40 cm, but is mostly in the 3–10 cm range. The matrix is a coarse-grained arkose. At some levels chaotically arranged tree trunks occur. The conglomerates are highly lenticular, alternating with gravelly or coarse-grained, very poorly sorted sandstone intercalations, which usually display large-scale trough cross-bedding. The conglomerates form packages 3–6 m, occasionally more than 10 m, in thickness, generally have a sharp

LITHOSTRATIGRAPHY OF THE SHEMSHAK GROUP

base, but do not fine upwards. Rarely, m-thick intercalations of argillaceous silt and fine-grained sandstone rich in plant debris occur. According to Repin (1987, pp. 215– 216), the formation contains Lower and Middle Jurassic plant remains in the middle part at Ekrasar. At Galanderud (Fig. 7F) it consists of about 70 m of fine conglomerates and sandstones. The overlying 150–250 m of argillites, siltstones and sandstones with coal seams, and the following 300 m of sandstones, siltstones and argillites of mostly lacustrine– origin (as demonstrated by bivalves and ammonites), were regarded by Repin as part of the formation, but are better placed in the Alasht Formation. At Gheshlagh, the succession consists mostly of interbedded fine quartz conglomerates, sandstones, siltstones and mudstones with palaeosol horizons. Boundaries. In places, the base of the formation is an angular unconformity, seen, for example, at the slope west of the road between Javaherdeh and Ramsar (Fig. 7G; see the Appendix). The top has not been observed by us. Fossil content. Apart from plant remains, wood pieces and tree trunks, no fossils are present. Age. Unfortunately, no reliable information is available concerning the age of the Javaherdeh Formation. The formation unconformably overlies the Kalariz Formation. The lower boundary therefore most probably corresponds roughly to the Triassic –Jurassic boundary. According to Bragin et al. (1976), the formation is overlain by the Dalichai Formation. Consequently, the age of the Javaherdeh Formation appears to be Hettangian –Early Bajocian. This is supported by finds of Early–Middle Jurassic plants at Ekrasar. Palaeoenvironmental interpretation. The Javaherdeh Formation is interpreted as deposits of mid to distal humid alluvial fans, and partly represents debris flow, partly proximal braided river deposits. Towards the south, it interfingers with the fluvial– lacustrine sediments of the Alasht Formation.

Conclusions Palaeoenvironmental evolution of the Shemshak Group As has been noted by several authors (Bragin et al. 1976; Nabavi 1980; Repin 1987; Aghanabati 1998; Fu¨rsich et al. 2005), the Shemshak Group displays a distinct zonation parallel to the strike of the mountain range. For this reason, the distribution of the various lithostratigraphic units is not uniform across the Alborz Mountains (Fig. 17). In the Triassic part of the group,

155

marine, partly deeper offshore, conditions existed in the north and fine-grained siliciclastic sediments were deposited (Ekrasar Formation). The unconformity with the underlying Elikah Formation is a drowning unconformity. Starting in the Norian, this basin was gradually infilled, which is evidenced by the prodelta and delta front sediments of the Laleband Formation. By Rhaetian times, fluvial–lacustrine conditions, shown by the Kalariz Formation, became established over much of the area. In the southern Alborz, in contrast, the shallowmarine platform carbonates of the Elikah Formation became subaerially exposed during the Carnian. Karst features are widespread and locally thick bauxitic –lateritic deposits accumulated, evidence of long-term subaerial exposure and a warm, comparatively humid climate. Volcanic activity produced small lava flows and tuff deposits. When sedimentation was resumed in the Norian, it occurred in fluvial, coastal plain and in marginal-marine brackish-water environments (Shahmirzad Formation). Around the Triassic –Jurassic boundary the situation became reversed: in the northern Alborz, the Kalariz Formation is overlain, with angular unconformity, by the Javaherdeh Formation, which persists throughout the Early Jurassic up to the Bajocian. The thick, coarse conglomerates represent alluvial-fan –braided-river deposits, and are evidence of uplift and erosion of the Cimmerian Mountain Range to the north, situated in the southern part of the present-day Caspian Sea. Towards the south, these conglomerates give way to finer-grained siliciclastic sediments (sandstones, silts) of the Alasht Formation, which represent fluvial –lacustrine environments, with the widespread development of coal swamps. These rocks interfinger with marginal-marine deposits in the south-central Alborz Mountains (e.g. the Damavand area). The marine character of the succession increases throughout the Early Jurassic. Shallowmarine conditions became established in the Shahmirzad area as early as the Late Sinemurian, whereas in other areas (e.g. Tazareh and Golbini) this was not the case before the Late Pliensbachian or Early Toarcian. Increased subsidence of the basin in the area of the southern Alborz started during the Toarcian, but was initially compensated for by equally high rates of sedimentation, so that environments remained in the range of the storm-dominated shelf (Shirindasht Formation). During the Aalenian, however, subsidence by far outpaced rates of sedimentation, so that the basin deepened distinctly with the accumulation of dark-grey, fine-grained siliciclastic sediments of the Fillzamin Formation. This trend was reversed towards the end of the Aalenian and during the

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Early Bajocian (Fu¨rsich et al. 2005). Forced regression led to the re-establishment of deltaic and flood-plain environments with coal swamps (Dansirit Formation).

An uplift –subsidence pulse occurred in the Bajocian. It produced two major unconformities that have been subsumed under the term Mid-Cimmerian, unconformity, and can be documented also in other

Fig. 18. Palaeogeographic evolution of the Iran Plate during the Middle Triassic–Jurassic and associated geodynamic events. 1, Eo-Cimmerian Event: initial plate coupling around the Middle– Late Triassic boundary, onset of Cimmerian orogeny; 2, Triassic–Jurassic boundary event: main uplift phase of the Cimmerian mountain chain (transition from syn- to post-orogenic deposits). 3, Mid-Cimmerian Event (Bajocian): break-up unconformity of the Neotethyan back-arc rift-basin in northern Iran (see Fu¨rsich et al. 2009). See the text for further explanation.

LITHOSTRATIGRAPHY OF THE SHEMSHAK GROUP

parts of the Iran Plate (Fu¨rsich et al. 2009). With the last of these unconformities around the beginning of the Late Bajocian the deposition of the Shemshak Group ends. Shelly pebble lags are commonly developed on top of the Dansirit Formation, and mark the beginning of the next major transgression and change in style of sedimentation. The long dominance of siliciclastic sedimentation ended and was replaced by the carbonate platform –basin system of the Dalichai and Lar formations.

Geodynamic significance of the Shemshak Group In contrast to former interpretations, the Shemshak Group is now thought to document not only postcollisional, but also synorogenic, sedimentation of the Cimmerian orogeny and the onset of Neotethyan back-arc rifting in northern Iran (Fig. 18). The Upper Triassic part of the succession represents synorogenic deposits in a peripheral foreland basin after initial coupling of the Iran and Turan plates (i.e. the onset of the Cimmerian orogeny), and the development of the Eo-Cimmerian unconformity. The main uplift of the Cimmerides, associated with the transition from the syn- to the early Liassic post-orogenic (molasse) stage,

157

occurred around the Triassic –Jurassic boundary (conglomerates of the Javaherdeh and coarse arkosic sandstones of the lower Alasht formations). From the Toarcian onwards, increasing subsidence rates in the southern Alborz indicate extension in northern Iran, interpreted as an eastwards-progressive opening of a Neotethys back-arc rift-basin. The Bajocian Mid-Cimmerian event represents the break-up unconformity of this back-arc rift basin (Fu¨rsich et al. 2009), developing into the South Caspian Basin during the remainder of the Middle and Late Jurassic. The detailed tectono-stratigraphy of the Shemshak Group and its implications for the timing of geodynamic events of the Cimmerian orogeny will be the subject of another paper. The research was carried out within the framework of MEBE and an institutional partnership between the Institute fu¨r Pala¨ontologie of Wu¨rzburg University and the Faculty of Engineering of the University of Tehran, financially supported by the Alexander von Humboldt Foundation. We would like to thank the Geological Survey of Iran for logistic support during fieldwork, and J. Taheri (Wu¨rzburg/Mashad), A. Shekarifard (Tehran) and F. Cecca (Paris), who helped us in the field for various periods of time. M. Gaetani (Milano) and M. Simmons (Abingdon) kindly reviewed the manuscript and made numerous useful suggestions.

Appendix Co-ordinates (WGS84) of measured sections and investigated localities in the Shemshak Group of the Alborz Mountains Locality

Golbini/Kuh-e-Ozom Base of section at Section point 1040 m at Tappehe Chefteh Qazi Tooy Kuh-e-Tepal Gheshlagh Tazareh Base of section at Section point at 1250 m Section point at 1620 m Section point at 2410 m Section point at 3565 m Section point at 3570 m Tazareh, west of mine Shahmirzad Base of section at Continuation at Kuhe bashm-e-Dehsufian at Sharif-Abad (east of Diktash) Base at

Co-ordinates

Information

N

E

378030 2200 378030 5900 368540 3300 378110 4500 368230 0300 368540 1300

568320 0200 568310 5700 568160 0100 578040 3400 548440 1500 558210 3000

368230 4400 368240 3700 368250 0500 368250 1700 368260 1500 368260 3100 368250 1200

548290 3100 548290 1000 548290 2500 548280 1000 548290 1100 548290 4800 548220 5200

358470 1500 358510 0300

538200 0700 538260 5900

Complete section measured (1980 m) Shahmirzad Formation Top Shirindasht Formation Shirindasht Formation Shemshak – Dalichai contact Shemshak – Dalichai contact Observation on facies succession Complete section measured (3830 m) Ekrasar Formation Top Shahmirzad Formation Alasht Formation Shirindasht Formation Topmost Fillzamin Formation Basal Dansirit Formation Shemshak – Dalichai contact Complete section measured (1900 m) Shahmirzad Formation Shirindasht Formation

358420 5900

538260 0100

Top 365 m of Shemshak measured Shirindasht Formation (Continued)

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Appendix Continued Locality

Parvar East of road to Parvar East of road to Parvar

Co-ordinates

Information

N

E

358590 2900

538290 4700

0

36802 01

00

0

53828 24

00

Rian

358460 0300

538390 0300

Ekrasar Road 5 km north of Bajan Emamzadeh Hashem Pass

368460 4800 358560 3000 358470 1300

508320 0800 528150 3600 528020 0200

Upper Ruteh valley (type locality)

358590 2000

518320 4800

Lalun valley (near type locality); section top at New road at Javaherdeh

368000 2500

518380 0600

368510 3800

508280 1800

368510 2800

508290 4700

368510 4600

508290 5900

368520 2200

508300 3900

368510 4200

508300 4100

368510 3300

508310 1100

368510 0500

508130 3000

Road between Javaherdeh and Ramsar Road between Javaherdeh and Ramsar West of road between Javaherdeh and Ramsar Road between Javaherdeh and Ramsar Road between Javaherdeh and Ramsar Road between Rahimabad and Qazvin Road between Javaherdeh and Ramsar Exposures above Ekrasar village

368510 420

508300 4500

368470 3700

508310 5400

Galanderud village, roadside section Base of section at Top of section at East of Baladeh

368250 5800 368260 4900 368120 3000

518530 5100 518520 5900 518490 2700

Paland

368000 5400 368010 0300

528530 2000 528530 4400

Mobarak Abad Valley, west of Emamzadeh Hashem Pass Base of section at Section point at 1080 m Section point at c. 1300 m Fillzamin, west of Damavand

358470 5600 358480 3900 358490 0400

518580 3800 518580 2200 518570 4700

Base of section at Top of section at Fillzamin

358500 2700 358500 4600 358500 5300

518520 5500 518530 2800 518530 2200

Basal 250 m of Shemshak measured (Ekrasar Formation) Top 322 m of Shemshak measured (upper Fillzamin – Dansirit) Topmost 54 m of Shemshak measured (Dansirit Formation) Contact Elikah– Ekrasar Contact Shemshak – Dalichai Lower and middle parts of Shemshak: observations on facies Lower and middle parts of Shemshak: observations on facies Upper part of Shemshak: observations on facies Sedimentological observations on Javaherdeh Formation Sedimentological observations on Kalariz Formation Sedimentological observations on Kalariz Formation Angular unconformity between Kalariz and Javaherdeh formations Sedimentological observations on Kalariz Formation Elikah– Ekrasar contact Sedimentological observations on Javaherdeh Formation 130 m of Kalariz Formation measured Sedimentological observations on Ekrasar Formation Section through Ekrasar and Laleband formations (c. 700 m) Contact Elikah– Ekrasar Contact Laleband –Kalariz Contact Elikah and Ekrasar, sedimentological observations on basal Ekrasar Contact Elikah and Ekrasar Section through Ekrasar and basal Laleband formations (134 m) Section through Shemshak Group (Ekrasar – Shirindasht: c. 1300 m) Contact Elikah– Ekrasar Shirindasht Formation Contact Shirindasht-Fillzamin Section through upper Shirindasht Formation to basal Dalichai Formation (590 m) Upper Shirindasht Formation Contact Shemshak – Dalichai Section through Dansirit Formation (115 m)

LITHOSTRATIGRAPHY OF THE SHEMSHAK GROUP

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Characterization of organic matter in the fine-grained siliciclastic sediments of the Shemshak Group (Upper Triassic– Middle Jurassic) in the Alborz Range, Northern Iran ALI SHEKARIFARD1– 3, FRANC¸OIS BAUDIN2,3, JOHANN SCHNYDER2,3 & KAZEM SEYED-EMAMI1 1

University College of Engineering, School of Mining Engineering, University of Tehran, Tehran, Iran (e-mail: [email protected]) 2

UPMC Universite´ Paris 06, UMR Tectonique, case 117, 4, place Jussieu, F-75005 Paris, France 3

CNRS, UMR Tectonique, F-75005 Paris, France

Abstract: Bulk organic geochemical and microscopic studies (Rock-Eval pyrolysis, light transmitted–uv microscope) were carried out on the shales of the Upper Triassic–Middle Jurassic Shemshak Group in the northern, central and southern Alborz Range of northern Iran. Total organic carbon (TOC) values range from 0 to 29.4 wt% (1.2 wt% on average) indicating a generally poor–moderate organic carbon content. Upper Triassic shales in the lower part of the Shemshak Group have been mainly deposited in marine/lake settings under dysoxic– anoxic conditions, with TOC ¼ 0.7 wt% on average. Toarcian– Aalenian shales in the upper part of the Shemshak Group were deposited under comparatively deeper marine oxic–dysoxic conditions with the lowest TOC contents recorded (0.3 wt% on average). Carbonaceous shales at different stratigraphic levels of the Shemshak Group show the highest TOC contents (14.2 wt% on average). Tmax values range from 439 to 599 8C (average 500 8C), indicating that the organic matter has experienced high temperatures during deep burial and active post-sedimentary tectonics. The hydrogen index (HI)–Tmax diagram shows the presence of Type IV kerogen of altered organic matter with a very low mean HI value. The palynofacies is characterized by the dominance of amorphous organic matter probably predominately derived from degradation of marine–nonmarine phytoplankton. The Upper Shemshak Group has low potential to produce petroleum, whereas the Lower Shemshak Group is an important effective petroleum source rock in the Alborz Range. The latter may have generated a considerable amount of petroleum at some localities (e.g. Tazareh and Paland) in the geological past.

The closing of the Palaeotethys seaway during the Early Late Triassic (Early Cimmerian orogeny), and thereby the collision of the Iran Plate with the southern margin of Eurasia, resulted in deposition of a thick pile of siliciclastic sediments in front of tectonically active and uplifted areas (Alavi 1996; Seyed-Emami 2003). The thick Upper Triassic – Middle Jurassic Shemshak Group is widely distributed throughout central, east and north Iran (Seyed-Emami 2003). It attains thickness up to 4000 m and disconformably overlies the carbonate rocks of the Lower–Middle Triassic Elikah Formation. In turn, it is disconformably overlain by marls and limestones of the Middle Jurassic Dalichai Formation. The Shemshak Group of the Alborz Range was deposited in marine–non-marine settings, including lacustrine, fluvio-deltaic–deep marine and even basinal environments that partly experienced anoxic conditions (Fu¨rsich et al. 2005). It mainly consists

of fine- to coarse-grained sandstones, siltstones, shales and highly carbonaceous shales, with numerous coal seams. The Shemshak Group seems to be a source rock in the eastern Alborz and is considered as the probable source rock for the east Caspian oil and gas fields, e.g. Khangiran (Rad 1982). In addition, a few oil and gas seepages would suggest that organic-rich sediments of the Shemshak Group may be a promising petroleum source rock in the Alborz Mountains (particularly in Gorgan and Mazandaran provinces). Moreover, thick and clean sandstone units (e.g. beach deposits: Fu¨rsich et al. 2006) within the group could provide good hydrocarbon reservoir rocks. Abundant overlying and interbedded shaly strata may serve as useful seals and source rocks in some cases. Because previous studies of the organic content of the Shemshak Group were devoted to the coal seams, the aim of our investigations is the characterization of sedimentary organic matter of the

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 161–174. DOI: 10.1144/SP312.7 0305-8719/09/$15.00 # The Geological Society of London 2009.

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shaly parts of the Shemshak Group in the northern, central and southern Alborz Range in order to evaluate its potential as a petroleum source rock.

The shaly units Using stratigraphy, age and depositional environment, the fine-grained siliciclastic sediments of the Shemshak Group have been categorized into three units: Upper Triassic shales (UTS), Toarcian –Aalenian shales (TAS) and carbonaceous shales (CS). The samples were collected from 12 sites in the northern, central and southern Alborz Range. The localities Javaherdeh, Ekrasar, Galanderud and Paland are situated in the northern Alborz; Tazareh, Shahmirzad, Damavand (Westof-Emamzadeh-Hashem), Sharif-Abad and Djam belong to southern Alborz. Baladeh and Parvar are located in central Alborz (Fig. 1). In the present study, only lower and upper parts of the Shemshak Group have been sampled. The sections in the Lower Shemshak Group include Tazareh, lower Parvar, lower Shahmirzad, Galanderud, Paland and West-of-Damavand. The sections in the Upper Shemshak Group belong to upper Parvar, Sharif-Abad, upper Shahmirzad and West-of-Emamzadeh-Hashem sections (Fu¨rsich et al. 2009a, b; Seyed-Emami et al. 2008). In addition, 23 samples from the Djam, Javaherdeh,

Baladeh and Ekrasar areas were investigated. Figures 2 and 3 show lithostratigraphic profiles of selected sections where the UTS and TAS units have been sampled and investigated. The UTS comprises two shaly main units, the so-called basal black shale (BBS) and the Ekrasar Formation (Fu¨rsich et al. 2009a). The BBS unit is only exposed close to the base of the Shemshak Group at Tazareh and Parvar (southern and central Alborz). This unit, with thicknesses up to 300 m, contains predominantly homogeneous blackish shales without evidences of bioturbation (Fig. 2). Towards the top, the dark shales grade into fine-grained sandstone and shale intercalations with large trough cross-bedding and hummocky cross-bedding. Some occurrences of freshwater bivalves (Unionites) in the lower Parvar section have also been observed (solely within the sandstones parts). These features have been interpreted as indicating a lacustrine setting corresponding to the BBS deposition (Fu¨rsich et al. 2009a). The Upper Triassic shales that form the basal part of the Shemshak Group in the northern Alborz belong to the Ekrasar Formation at Galanderud, Ekrasar and Paland. At Galanderud, the thick Ekrasar Formation consists predominantly of monotonous greenish to dark-grey silty-clay sediments. Dark limestone intercalations at the base

Fig. 1. Location map showing distribution of the Shemshak Group with the studied localities in the northern, central and southern Alborz Range, northern Iran. 1, Shahmirzad area (lower and upper Shahmirzad sections); 2, Parvar area (lower and upper Parvar sections); 3, Sharif-Abad section; 4, Djam area; 5, Tazareh section; 6, Damavand area (West-of-Damavand and West-of-Emamzadeh-Hashem sections); 7, Baladeh area; 8, Paland section; 9, Galanderud section; 10, Ekrasar area; 11, Javaherdeh area.

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Fig. 2. Lithostratigraphical columns of the studied sections representative of the Lower Shemshak Group (UTS unit). Open arrows indicate the position of the samples taken for Rock-Eval analysis. Black arrows show the samples taken for palynological observations. Lithostratigraphical profiles after Fu¨rsich et al. (2009a).

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Fig. 3. Lithostratigraphical columns of the studied sections representative of the Upper Shemshak Group (TAS unit) in the southern Alborz Range. M23 corresponds to carbonaceous shale at West-of-Emamzadeh-Hashem. Open arrows indicate the position of the samples taken for Rock-Eval analysis. Black arrows show the samples taken for palynological observations. Lithostratigraphical profiles after Fu¨rsich et al. (2009b) (upper Parvar section), Seyed-Emami et al. (2008) (upper Shahmirzad section) and Fu¨rsich et al. (2009a) (West-of-Emamzadeh-Hashem section).

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contain a marine fauna (bivalves and ammonoids) of late Carnian –early Norian age (Fu¨rsich et al. 2009a). It is overlain by a thick sequence (300– 500 m) of initially argillaceous, later increasingly fine-sandy brownish-grey siltstones (Fig. 2). At Ekrasar, the facies is similar to Galanderud and contains the same fauna. Occasionally there are layers of blackish argillaceous siltstone enriched in pyrite. At Paland, the Ekrasar Formation starts with black shales. Up-section, there are intercalations of laminated sandstones and dark shales. The Paland section is lithologically different from the former sections and consists predominantly of black shales with pyrite-rich levels. The thick TAS unit is highly bioturbated and consists largely of greenish–dark-grey argillaceous siltstones with fossiliferous carbonate concretions (Fillzamin and Shirindasht formations of Fu¨rsich et al. 2009a). It has been deposited in an outer shelf setting (Fu¨rsich et al. 2005). The lack of lamination, strong bioturbation and the colour of the sediments indicate that the TAS unit was largely deposited in oxic conditions with fewer suboxic– anoxic intervals. Parts of the shaly sediments in the Sharif-Abad section may indicate oxygen-poor conditions at the sea floor. This facies type also occurs in the upper part of the Shemshak Group in the southern Alborz in the upper Parvar, upper Shahmirzad and West-of-Emamzadeh-Hashem sections. The TAS is overlain by sandstones packages with coaly and carbonaceous layers (Dansirit Formation: Fu¨rsich et al. 2009a) (Fig. 3). Carbonaceous shales, with a thickness of a few cm to 2 m, occur at different stratigraphic levels of the Shemshak Group at Javaherdeh, Shahmirzad, West-of-Emamzadeh-Hashem and Baladeh. These have been deposited in swamps and small lakes on the flood, as well as coastal plains (Fu¨rsich et al. 2005). In some cases, carbonaceous shales are associated with marine trace fossils (e.g. Diplocraterion parallelum), which indicate a paralic environment (Fu¨rsich et al. 2005). In the West-of-Emamzadeh-Hashem section, near the top of the group there is a carbonaceous shale layer (sample M23), sandwiched between sandstones of the Dansirit Formation, overlaying the TAS unit. This was probably deposited in a coastal swamp on an upper deltaic coastal plain (Fu¨rsich et al. 2009a, b) (Fig. 3). Other carbonaceous shales are overlain by the TAS in southern Alborz and by thick conglomeratic beds in northern Alborz. At Baladeh the carbonaceous shale occurs at the basal part of the Shemshak Group.

Samples and methods A total of 236 shaly samples were taken for analysis of their organic content, source and thermal

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maturity of the organic matter of the Shemshak Group in the northern, central and southern Alborz Range. The samples belong to greenish, grey and black shales, and to several highly carbonaceous shales of the Shemshak Group. An aliquot of each sample was open-air dried and crushed in a rotary-mill to obtain a 5 mm-powder. The total carbon, as well as source and thermal maturation of the organic matter, were estimated using a Rock-Eval OSA instrument (Espitalie´ et al. 1985a, b, 1986). Standard notations are used: S1 and S2 in mg hydrocarbons (HC) per g of rock; Tmax is expressed in 8C; total organic carbon (TOC) content in weight% (wt%) and hydrogen index (HI ¼ S2/TOC  100) in mg HC per g of TOC. To obtain pure kerogen, 32 samples with the highest TOC and HI values were selected for kerogen isolation in order to perform palynological studies. After drying and removal of the altered parts, the samples were concentrated by demineralization using 37% HCl and 70% HF acids without heavy liquid separation. After concentration, the kerogens were filtered using a 10 mm-mesh. For all isolated kerogens three filtered slides and three unfiltered slides were prepared. Qualitative palynofacies observations were performed on 30 samples using both transmitted and uv (ultraviolet) light. In addition, quantitative observations were performed on some samples, using a Zeiss Axioplan 2 imaging microscope. Relative surface proportions of the various categories of particles were measured using a square grid. The counting was considered to be relevant for a counting surface that corresponds to at least 500 counted particles.

Results Rock-Eval data Interpretable Tmax values from the Shemshak Group show a large variability, being in the range of 439– 599 8C (average 500 8C). This indicates that most of the organic matter experienced high temperature during burial and is now over-mature with respect to oil-generation (Fig. 4). It should be mentioned that the precise assessment of the thermal maturity levels of high mature samples using Rock-Eval Tmax data is less reliable than vitrinite reflectance data. Total organic carbon contents of the shales (except a real coal) fluctuate between 0 and 29.4 wt%, with a mean value around 1.2 wt%. Maximum TOC values were found in carbonaceous shales (up to 29.4 wt%) and in black shales of the basal part of the Paland section (up to 3.5 wt%). Eleven highly carbonaceous shales and one real coal were sampled along the studied sections. The organic content of

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carbonaceous shales shows medium HI value (248 mg HC/g TOC sample M23). It comes from the upper part of the Shemshak Group at Westof-Emamzadeh-Hashem section. This sample revealed the lowest Tmax value (439 8C) of all samples and is thermally early mature. The occurrences of high Tmax values, together with very low HI and generally low– moderate TOC values, suggest the presence of hydrogendepleted organic matter that is, at least partly, a consequence of high thermal maturity level, generation and expulsion of petroleum (Fig. 5). Because the organic matter is very mature, precise determination of its origin using a modified Van-Krevelen diagram is not diagnostic, and microscopic study of organic matter is necessary.

Palynofacies

Fig. 4. HI– Tmax diagram for the shales of the Shemshak Group in the Alborz Range, northern Iran.

the carbonaceous shales is around 14.2 wt% (range 3.3–29.4 wt%) and of the coal 53 wt%. Average TOC of the BBS unit is around 0.62 wt%. The TOC contents of the basal Ekrasar Formation range from 0.5 wt% at Ekrasar and Galanderud sections to 1.2 wt% at Paland section. Some black shales in the basal part of Paland section have a high TOC value of about 3.5 wt%. Generally, the mean TOC value of the Upper Triassic shales of the Shemshak Group is about 0.7 wt%. TOC values of the Toarcian –Aalenian shales ranges from 0 to 0.65 wt%, averaging at 0.3 wt%. The results of the Rock-Eval pyrolysis of the studied shaly units are summarized in Table 1. Hydrogen index values are generally very low, between 0 and 248 mg HC/g TOC, being 28 mg HC/g TOC on average. Only one sample from

Dispersed organic matter (OM) in the shaly sediments of the Shemshak Group in the studied samples is derived from both continental and marine sources that can be classified into three groups including: (1) amorphous OM (AOM); (2) terrestrial OM; and (3) marine and brackish palynomorphs. The classification of the different types of organic matter is presented in Table 2. Amorphous organic matter (AOM) corresponds to relatively large, flaky, granular, cloudy orange – dark brown or even grey-black particles (Fig. 6A). It is generally thought to be derived from the degradation of marine or lacustrine algal and bacterial constituents, rich in hydrocarbons. Alternatively, it can also be derived from waxy coatings of terrestrial plants (Tyson 1995). Terrestrial particles are subdivided into a palynomacerals group, which is composed of vascular plants fragments, and a palynomorphs group, composed here of spore pollens, a few fungal filaments and a few possible occurrences of freshwater – brackish algae (Botryococcus-type algae), which are however oxidized and altered and therefore difficult to recognize. Palynomacerals are subdivided into opaque and semi-opaque particles, corresponding mainly to woody lignin debris and translucent phytoclasts linked to cuticles and other related phytoclasts remains (Steffen & Gorin 1993) (Fig. 6B–D). Woody components are weak and break easily during transportation into smaller particles. Sometimes they are introduced into a basin by wind after woodlands and swamps have been on fire (Taylor et al. 1998). They can be abundant in fluvial–lacustrine –deltaic systems (e.g. coastal plain –prodelta, freshwater swamps and lagoonal deposits) to marine and pelagic deposits (Tyson 1995; Ercegovac & Kostic 2006). They show no fluorescence. Cuticles are found in fluvial, deltaic, estuarine, shelf and submarine-fan

Table 1. Organic geochemical (Rock-Eval) characteristics of the studied shales in the northern, central and southern Alborz Range The shaly units

Sample No.

Tmax(8C) range (average)

TOC (wt%) range (average)

HI range (average)

Carbonaceous shales

11 (four areas)

439– 579 (520)

3.34 – 29.4 (14.2)

0 – 248 (36.5)

Toarcian –Aalenian shales

39 (four areas)

459– 506 (473)

0 – 0.65 (0.3)

0 – 28 (12)

0– 0.15 (0.01) 0– 0.14 (0.04)

108 (two areas)

510– 599 (557)

0 – 1.74 (0.62)

0 – 36 (5)

0– 0.3 (0.02)

0 – 64 (17.1)

0– 0.12 (0.06) 0– 0.42 (0.1)

Upper Triassic shales

Basal black shales Ekrasar Formation

59 (three areas)

443– 579 (485)

0.3 – 3.5 (0.7)

S1 range (average)

S2 range (average)

0.01– 1.11 (0.42) 0– 19.06 (3.45)

0– 0.44 (0.03)

Depositional environment Coastal plain– swamps/ lakes (anoxic) Deeper marine setting (oxic – dysoxic) Lacustrine/non-marine settings (dysoxic-anoxic) Deep-marine setting (oxic – anoxic)

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Fig. 5. HI v. TOC values for the studied samples of the Shemshak Group in the Alborz Range, northern Iran.

deposits, and are indicative of the vicinity of the source (Tyson 1995). They show yellow– orange fluorescence. Marine and brackish palynomorphs correspond to some oxidized dinoflagellate cysts and acritarchs (Micrhystridium sp.) (Fig. 6E, F). Based on qualitative observations of palynofacies, almost all slides of every shaly unit of the Shemshak Group show the same palynofacies characters. Generally AOM dominates in the total, non-filtered slides, together with minor –moderate amounts of vascular-plant debris and some palynomorphs (Fig. 7). This AOM is mainly composed of dark-brown, grey –black heterogeneous AOM that is often associated with black fine-grained particles

(pyrite framboids and/or inertinite particles). This dark-brown and black AOM shows no fluorescence, indicating strong degradation of OM and loss of hydrocarbons. In addition, there are dark-brown and black sporomorphs in some samples of the UTS unit that show high thermal alteration (Fig. 6G– I). These results are in accordance with Rock-Eval pyrolysis data, suggesting that these samples present a very altered and thermally mature –over-mature suite of organic matter. Even though the main part of AOM probably originates from algal and bacterial production in marine and/or lacustrine settings, alteration and oxidation of terrestrial-derived particles cannot be ruled out. Especially considering the frequent occurrences of

Table 2. Classification of sedimentary organic matter of the shales of the Shemshak Group in the northern, central and southern Alborz Range Group Terrestrial OM

Marine(OM) palynomorphs Amorphous OM

Constituent

Original maceral/ kerogen type

Origin

Opaque phytoclasts Semi-opaque phytoclasts Translucent phytoclasts Fungal phytoclasts Spores and pollen Resins Botryococcus sp.? Acritarchs, Dinocysts

Inertinite/IV Vitrinite/III Cutinite/II Inertinite/IV Sporinite/II Resinite/I Alginite/I Liptinite/II

Oxidized wood debris, charcoal. Wood tissues Cuticle (leaf epidermal tissue) Fungal Higher land plants Higher land-plant secretions Freshwater algae Marine phytoplankton

Grey–black AOM Reddish AOM

Liptinite/I and II InertiniteVitrinite/IV–III

Degraded phytoplankton, algae Degraded higher-plant debris

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Fig. 6. Microphotographs of palynofacies slides of less mature samples (A–F) from the TAS unit, Shemshak Group, northern Iran. (A) Brown amorphous organic matter (AOM) (planktonic remains). (B) Planar translucent phytoclast showing cuticular structure. (C) Brown pollen with opaque phytoclasts. (D) Brown spore. (E) Dinoflagellate cyst. (F) Acritarch. Microphotographs of palynofacies slides of over-mature samples (G–I) from the UTS unit. (G) Altered AOM with some oxidized particles. (H) and (I) Black sporomorphs.

wood fragments in the sections of the Shemshak Group, it seems that a mixture of AOM and cuticles –woody debris is present. This could point to degradation and amorphization of lignocellulosic tissues from terrestrial plants as an organic matter source, as shown by other authors (e.g. Lallier-Verge`s et al. 1998), especially in the carbonaceous shales of the Shemshak Group. Some homogeneous, red, orange and yellow particles with a glossy and vitreous appearance, that show brown–orange fluorescence colours, were observed in the Toarcian–Aalenian shales at the Sharif-Abad section and have been interpreted as resin particles (resinite).

Some dinoflagellate cysts and acritarchs (Fig. 6E, F) have been found in the TAS of the upper part of the Shemshak Group (e.g. at SharifAbad). Palynological observations of the basal shales of the Shemshak Group (Ekrasar Formation) at the Galanderud section showed a relatively rich assemblage of dinoflagellate (Ghasemi-Nejad et al. 2004). These palynomorphs indicate the occurrence, at least in some levels, of restricted marine conditions during deposition of the TAS and the Ekrasar Formation. Altered forms have been interpreted with doubts as possible Botryococcus-type algae in some samples of the basal black shales (BBS) from the Tazareh section. This may point

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Fig. 7. Ternary diagram of sedimentary organic matter composition of the shaly sediments from the Shemshak Group in the Alborz Range, northern Iran (AOM, amorphous organic matter; TOM, terrestrial organic matter; MP, marine palynomorphs).

to lacustrine –brackish environments for the BBS unit, at least at some levels or places, in accordance with sedimentological observations (Fu¨rsich et al. 2009a).

Discussion The TOC content of the Shemshak Group was previously measured on 58 samples from the Jajarm and Shahpasand areas (Rad 1982), on 179 samples from the Shemshak, Pol-e-Dokhtar and Vaneh sections (Baudin & Teherani 1991), and on 80 samples from the Galanderud section (Rahimpour-Bonab et al. 2002). The organic carbon contents ranged from 0.2 to 15.6 wt% according to Rad (1982), from 0.2 to 51 wt% according to Baudin & Teherani (1991) and from 0.01 to 72.4 wt% according to Rahimpour-Bonab et al. (2002). In the present study, the TOC content of the shaly samples varies between 0 and 53 wt%, with a mean of 1.5 wt%. Thus, based on TOC contents, the Shemshak Group in the Alborz Range contains poor, fairly good, good and even excellent source rocks. For example, the UTS unit at Paland, as well as Tazareh and Parvar sections, together with the carbonaceous shales have the highest organic richness and can be considered as fair –excellent gas-prone source-rock horizons. On the other hand, the TAS unit shows the lowest OM content, indicating a poor source rock. Because the studied samples are very mature, present TOC values are residual amounts. At least half of the original organic carbon has been lost owing to hydrocarbon and carbon dioxide generation and expulsion (Littke & Leythaeuser 1993). Owing to the high maturity of the samples studied and variation in the palynofacies and OM types, exact determination of the masses of volatile products (hydrocarbon and carbon dioxide) expelled from the source rocks are impossible.

Rock-Eval pyrolysis data show mature–overmature organic matter with low hydrogen index (HI) values mainly between 0 and 81 mg HC/g TOC. Only one early mature carbonaceous shale has a medium HI value (248 mg HC/g TOC). On the HI–Tmax diagram, this sample plots as a Type III organic matter. The S2 v. TOC diagram shows mixed oil/gas-prone organic matter. All other samples contain a hydrogen-depleted and oxygen-rich material that is classified as Type IV kerogen of altered organic matter (Fig. 4). Because of the high maturity level of the samples, the original kerogen type cannot be determined from Rock-Eval data. The reasons for the low HI values of the organic matter in sediments are: (1) the strong mineral– matrix effect; (2) the presence of high percentages of oxidized and opaque phytoclasts or inertinite macerals; (3) oxidation of organic matter during early diagenesis; and (4) thermal alteration of the organic matter and petroleum generation (Tissot & Welte 1984; Tyson 1995; Nzoussi-Mbassani et al. 2003). The low HI values in organic-poor silty-clay samples can be a result of a mineral–matrix effect and adsorption of hydrocarbons generated on clay mineral surfaces during pyrolysis (Espitalie´ et al. 1985a, b, 1986). Thus, some isolated kerogens were prepared and analysed using Rock-Eval. They showed more or less the same results. Hydrogen index values of most kerogens are between 5 and 110 mg HC/g TOC, whereas the bulk of the samples range from 0 to 81 mg HC/g TOC. The mineral– matrix effect is therefore not important. Palynological observations indicate the dominance of AOM in all the studied slides (in counted slides: between 65 and 87%), and a maximum of oxidized higher-plant debris of 18% from the UTS unit of Galanderud section. Thus, the low mean HI values are not a result of the presence of large amounts of opaque phytoclasts. According to sedimentological evidence and the presence of marine palynomorphs (acritarchs and some dinoflagellate cysts) in the shaly sediments, the amorphous organic matter was originally a mixture of Type II –I kerogen OM. This OM most probably was derived from degradation of marine or non-marine phytoplankton and/or bacteria. A minor contribution of AOM from land-plant tissues may be also possible at some levels, as suggested by palynofacies observations. In the black shales of the lower part of the Shemshak Group (BBS) at the Tazareh and Parvar sections, some questionable, but considerably altered, Botryococcus-type algae may be indicative of freshwater –brackish settings. Based on TOC values, the lack of bioturbation in the shaly sediments of the BBS unit and the presence of pyrite in some samples, most probably the oxidation of organic matter during early diagenesis,

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was not significant and conditions for preservation of organic matter were fair –good. In general, AOM with fluorescence is derived from planktonic or bacterial organic matter deposited in an anoxic setting and shows yellow colour in the immature state (Lewan 1986). Nearly all of the AOM in the studied samples shows no or only very weak fluorescence, which suggests strong alteration of the original material. Because all of the samples are thermally mature –over-mature,

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the fluorescence extinction and very low mean HI values could be due to thermal alteration (Tissot & Welte 1984; Tyson 1995; Nzoussi-Mbassani et al. 2003). Only some resinite particles show brown fluorescence, the colour of which could be related to the high resistance of resinite material. Marine and lacustrine AOM of phytoplanktonic and bacterial origin (less resistant material) is very sensitive to the maturity level, in contrast to more resistant resin, spore, pollen and cuticle material.

Fig. 8. Range of Tmax values for eight sections analysed in the present work by Rock-Eval pyrolysis. Additional data from Vaneh, Pol-e-Dokhtar and Shemshak sections from Baudin (1989). Average Tmax values for each section are figured by grey and black circles. Note that Ekrasar, Galanderud, Sharif-Abad and Shahmirzad sections show lower Tmax values, always below 522 8C, and therefore lower thermal maturity of OM when compared to other sections. Generally, the sections located on the axis of the Alborz Range (figured by black circles) present higher Tmax values than Ekrasar, Galanderud, Sharif-Abad and Shahmirzad sections, located at the northern and southern margins of the Alborz Range (grey circles). This may indicate partly a tectonic control during the formation of the Alborz Mountains, evidenced by the observed high variability of OM thermal maturity and OM alteration. This tectonic effect was superimposed to primary burial of OM during basin evolution to produce the observed high–very-high OM thermal maturity.

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Based on the high maturity level of the samples (average Tmax 500 8C), thermal alteration was the major factor in organic matter alteration. Position of the samples on the HI –Tmax diagram (Fig. 4) indicates altered organic matter with Type IV kerogen. Thus, the samples are close to the end of the kerogens evolution– alteration pathway. In general, high thermal maturity can be caused by deep (syn-/post-sedimentary) burial, by postburial tectonic activity or by high heat-flow rates in the vicinity of magmatic intrusive –extrusive bodies (Petmecky et al. 1999). As mentioned before, the Shemshak Group, with thicknesses of up to 4000 meters, experienced locally very high subsidence and sedimentation rates, characteristic for (young) rift basins (Fu¨rsich et al. 2005). In addition, the Alborz Range experienced strong tectonic and magmatic activity during the Cretaceous and Cenozoic (Sto¨cklin 1968; Alavi 1996; Allen et al. 2003). Our average Tmax values at each section draw a contrasted picture of the sections with respect to the thermal alteration of OM. Mean Tmax values and the range of Tmax values of each section are indicated in Figure 8. Original data from the Vaneh, Pol-e-Dokhtar and Shemshak type locality sections, based on the work of Baudin (1989), are also plotted in Figure 8. In comparison to samples from sections located in central positions along the axis of the Alborz Range it clearly appears that the samples from the Ekrasar, Galanderud, Shahmirzad and Sharif-Abad sections on the northern and southern margins of the Alborz Range show significantly lower Tmax values. They therefore experienced lower temperatures. Specifically, maxima of Tmax values from the Ekrasar, Galanderud, Shahmirzad and Sharif-Abad sections are all below 522 8C, whereas, apart from Vaneh section, all other sections show mean Tmax values

close to or above 520 8C (Fig. 8). These data suggest that despite spatially highly variable burial rates in the complex Late Triassic –the Middle Jurassic tectonic setting, there is a link between OM thermal maturity/alteration and post-burial tectonic stress in the formation of the Alborz Range thrust belt (Fig. 9). To summarize, the variations of Tmax values (from 439 to 599 8C: Figs 4 and 8) within the studied sections seem to be related, besides the strong tectonic stress, especially to the emplacement of thrust sheets, to the deep burial and, to a lesser degree, to the post-sedimentary magmatic activity. Grey–black amorphous organic matter dominates in the BBS (more than 86%). Other main organic constituents of the BBS are small quantities of oxidized higher-plant debris. Spherical pyrite aggregates occur in some the palynological slides. According to the dysoxic– anoxic lacustrine depositional environment proposed for the BBS unit, the AOM is probably of algal and bacterial origin. It seems that the original organic matter from these beds was rich in hydrogen and has been altered during deep burial due to the loss of hydrogen. Amorphous organic matter with probably originally high HI values belongs to the shaly facies and generated petroleum during deep burial. The only early mature sample (Tmax ¼ 439 8C) from a carbonaceous shale (West-of-EmamzadehHashem section) shows moderate/high hydrogen contents (HI ¼ 248 mg HC/g TOC, S1 þ S2 ¼ 20.1 mg HC/g rock), whereas the others are overmature (average Tmax ¼ 510 8C) with a very low HI (17 mg HC/g TOC on average). Palynological studies indicate a low percentage of terrestrial organic matter and a dominance of amorphous organic matter. This reveals altered Type II kerogen. Based on its HI value the sample shows

Fig. 9. A structural cross-section of the Alborz Range modified from Sto¨cklin (1968). Approximate positions of the Ekrasar (the samples are from the lower part of the Shemshak Group, with lower Tmax values) and Shemshak (the samples are from the upper part of the Shemshak Group, with higher Tmax values) areas have been indicated. Generally, towards the central part of the cross-section, near the axis of the Alborz Range, complexity and intensity of tectonic processes increase, whereas the southern and northern margins show more simple tectonic styles with lower deformation.

DISPERSED OM IN THE SHEMSHAK GROUP, IRAN

the geochemical characters of Type III kerogen. This supports hydrocarbon generation from the carbonaceous shales. The AOM shows a very weak fluorescence colour. Microscopic observations indicate the same palynological matter in all carbonaceous shales. Some carbonaceous shales on the HI –Tmax diagram show a trend of decreasing HI values along with increasing Tmax values. Thus, it seems that some over-mature carbonaceous shale samples have generated hydrocarbons in the past and are effective petroleum source rocks. In spite of their high thermal maturity level, few carbonaceous shales still have a good source rock potential (e.g. sample J4 from the Javaherdeh area, S1 þ S2 ¼ 12.9 mg HC/g rock, Tmax ¼ 510 8C). This indicates that these levels were probably very rich in oil-prone organic matter at immature state. Thus, most probably carbonaceous shales have produced a lot of petroleum in the Alborz Range. The low mean TOC value (0.3 wt%), low pyrite contents, high bioturbation and marine palynomorphs suggest that the organic matter of the TAS has been largely deposited in a marine settings under oxic conditions with rare suboxic–anoxic episodes. The TAS of the Shemshak Group has a low potential to produce petroleum in Alborz Ranges. Moderate–high residual TOC contents, the presence of pyrite, lack of bioturbation, and the colour of sediments of the BBS unit (at the Tazareh and Parvar sections) and the Ekrasar Formation (at the Paland section) show organic-rich sediments in the lower part of the group, being mainly deposited in an anoxic setting. The Upper Triassic organic-rich shales in the lower part of the Shemshak Group at the Tazareh, Parvar and Paland sections, and the carbonaceous shales with moderate –high TOC, have the best potential to have generated oil and gas in the geological past. The fine-grained siliciclastic sediments of the lower Shemshak Group are an effective source rock in the northern, central and southern Alborz Range.

Summary and conclusions There are six major conclusions from this study. † Fine-grained siliciclastic sediments of the Upper Triassic –Middle Jurassic Shemshak Group in the northern, central and southern Alborz Range of northern Iran can be classified into three groups including: (1) Upper Triassic shales; (2) Toarcian –Aalenian shales; and (3) carbonaceous shales. † Upper Triassic shales in the lower part of the Shemshak Group occur both in the Ekrasar Formation at Galanderud, Ekrasar and Paland, and at Tazareh and the lower Parvar sections.

†

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They have been mainly deposited in marine – lacustrine settings under dysoxic– anoxic conditions, with a 0.7 wt% residual TOC value on average. Greenish–grey Toarcian –Aalenian shales in the upper part of the Shemshak Group occur at the Shahmirzad, Sharif-Abad, West-ofEmamzadeh-Hashem and upper Parvar sections. They have the lowest residual TOC contents (average 0.3 wt%). The latter group was deposited under deeper marine oxic–dysoxic conditions. Carbonaceous shales come predominantly from the Javaherdeh and Shahmirzad areas at different levels of the Shemshak Group. The highest TOC values are found in carbonaceous shales (mean 14.2 wt%) and in the basal black shales of the Tazareh and Paland sections. Carbonaceous shales were mainly deposited in coastal swamps under anoxic conditions. Dispersed organic matter of the fine-grained siliciclastic sediments of the Shemshak Group is categorized into three groups: amorphous OM; terrestrial OM; and marine and brackish palynomorphs particles. Palynological observations indicate the dominance of amorphous OM in all samples, being most probably derived from degradation of marine –lacustrine algae and bacteria, together with degradation of cellulosic vascular-plant tissues within some levels. The shales predominantly contain late –overmature OM with very low mean HI values (28 mg HC/g TOC). According to geochemical characters, the organic matter of the Shemshak Group is rich in hydrogen-depleted materials. It can be classified as altered OM that corresponds in its geochemical composition largely to Type IV kerogen. Organic matter of the Shemshak Group has experienced high temperature during burial and strong tectonic stress. Extensive alteration of organic matter is predominately due to thermal alteration during and after the catagenesis stage. Based on Rock-Eval data and palynological studies, the Lower Shemshak Group, especially thick Upper Triassic black shales facies, along with carbonaceous shales are important effective petroleum source rocks in the northern, central and southern Alborz, which may have generated significant volumes of petroleum at some localities (e.g. Tazareh, Parvar and Paland). These sediments are present-day gas-prone source rocks. The Upper Shemshak Group, i.e., the Toarcian –Aalenian shaly unit, is a poor source rock.

We thank Prof. Dr R. Littke (RWTH Aachen) and Dr J.-P. Houzay (TOTAL) for their review of the manuscript. This research has been carried out within the institutional partnership between the School of Mining

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Engineering, University of Tehran and the Institute of Palaeontology, University of Wu¨rzburg, sponsored by the Alexander von Humboldt Foundation and within the framework of the Middle East Basin Evolution programme (MEBE) based at the Universite´ Paris 6, Paris. We thank Prof. Dr F. T. Fu¨rsich for his critical reviews of an earlier version of the manuscript, and Prof. Dr F. T. Fu¨rsich and Prof. Dr M. Wilmsen for the geological profiles. We also would like to thank logistic support by Geological Survey of Iran. Florence Savignac (Universite´ Paris 6) is acknowledged for her technical support.

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tectonic event (Bajocian) in the Alborz Mountains, Northern Iran: evidence of the break-up unconformity of the South Caspian Basin. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 189– 203. G HASEMI -N EJAD , E., A GHA -N ABATI , A. & D ABIRI , O. 2004. Upper Triassic dinoflagellate cysts from the base of the Shemshak Group in north of Alborz Mountains, Iran. Review of Palaeobotany and Palynology, 132, 207–217. L ALLIER -V ERGE` S , E., P ERRUSSEL , B. P., D ISNAR , J.-R. & B ALTZER , F. 1998. Relationships between environmental conditions and the diagenetic evolution of organic matter derived from higher plants in a modern mangrove swamp system (Guadeloupe, France West Indies). Organic Geochemistry, 29, (5– 7), 1663–1686. L EWAN , M. D. 1986. Stable carbon isotopes of amorphous kerogens from phanerozoic sedimentary rocks. Geochimica et Cosmochimica Acta, 50, 1581– 1591. L ITTKE , R. & L EYTHAEUSER , D. 1993. Migration of oil and gas in coals. In: L AW , B. E. & R ICE , D. D. (eds) Hydrocarbons From Coal. AAPG, Studies in Geology, 38, 219 –236. N ZOUSSI -M BASSANI , P., D ISNAR , J.-R. & L AGGOUN D EFARGE , F. 2003. Organic matter characteristics of Cenomanian– Turonian source rocks: implications for petroleum and gas exploration onshore Senegal. Marine and Petroleum Geology, 20, 411– 427. P ETMECKY , S., M EIER , L., R EISER , H. & L ITTKE , R. 1999. High thermal maturity in the Lower Saxony Basin: intrusion or deep burial? Tectonophysics, 304, 317–344. R AD , F. K. 1982. Hydrocarbon potential of the eastern Alborz Region, NE Iran. Journal of Petroleum Geology, 4, 419–435. R AHIMPOUR -B ONAB , H., Z AMANI , Z. & K AMALI , M. R. 2002. Source rock evaluation of Shemshak Formation in the central Alborz basin: A preliminary investigation. Iranian International Journal of Science. 3(2), 235–262. S EYED -E MAMI , K. 2003. Triassic in Iran. Facies, 48, 91–106. S EYED -E MAMI , K., F URSICH , F. T., W ILMSEN , M., M AJIDIFARD , M. R. & S HEKARIFARD , A. 2008. Lower and Middle Jurassic ammonoids of the Shemshak Group in Alborz, Iran and their palaeobiogeographical and biostratigraphical importance. Acta Palaeontologica Polonica, 53(2), 237– 260. S TEFFEN , D. & G ORIN , G. 1993. Palynofacies of the upper Tithonian deep-sea carbonates in the Vocontian trough (SE France). Bulletin des Centres Recherches Exploration–Production Elf Aquitaine, 17(1), 235–247. S TO¨ CKLIN , J. 1968. Structural history and tectonics of Iran: A Review. AAPG Bulletin, 52(7), 1229–1258. T AYLOR , G. H., T EICHMU¨ LLER , M., D AVIS , A., D IESSEL , C. F. K., L ITTKE , R. & R OBERT , P. 1998. Organic Petrology. Borntraeger, Berlin. T ISSOT , B. P. & W ELTE , D. H. 1984. Petroleum Formation and Occurrence. Springer, Berlin. T YSON , R. V. 1995. Sedimentary Organic Matter: Organic Facies and Palynofacies. Chapman & Hall, London.

The Shemshak Group (Lower – Middle Jurassic) of the Binalud Mountains, NE Iran: stratigraphy, depositional environments and geodynamic implications ¨ RSICH1 & JAFAR TAHERI3 MARKUS WILMSEN1,2,*, FRANZ THEODOR FU 1

GeoZentrum Nordbayern der Universita¨t Erlangen-Nu¨rnberg, Fachgruppe Pala¨oUmwelt, Loewenichstrasse 28, D-91054 Erlangen, Germany

2

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany 3

Geological Survey of Iran, NE Branch, P.O. Box 91735-1166, Mashad, Iran *Corresponding author (e-mail: [email protected])

Abstract: The Lower–lower Middle Jurassic non-marine sedimentary succession of the Binalud Mountains of NE Iran is correlated with the Jurassic part of the Shemshak Group of the Alborz Mountains and subdivided into three formations: the Arefi, the Bazehowz and the Aghounj formations. The succession rests, with angular unconformity, on a metamorphic basement deformed during the Late Triassic Eo-Cimmerian orogeny. The lowermost unit, the Arefi Formation, is subdivided into a lower Derekhtoot Member and an upper Kurtian Member. The Derekhtoot Member (up to 750 m thick) consists of very coarse-grained, chaotic boulder beds, breccias and conglomerates representing rock-fall deposits and proximal–middle alluvial fans, deposited along steep fault scarps. The succeeding Kurtian Member (.300 m) comprises finer-grained conglomerates with well-rounded clasts, reflecting deposition in a proximal braided river system. The overlying Bazehowz Formation is more than 1000 m thick and consists of vertically stacked, decametrescale channel-fill cycles of the middle reaches of a braided fluvial system. The uppermost unit, the Aghounj Formation, consists of at least 400 m of granule- to pebble-size, thick-bedded and large-scale trough cross-bedded quartz conglomerates and sandy interbeds of a proximal braided fluvial system. The overall succession fines upwards due to erosion, down to metamorphic basement, of a high-relief source area in the NE, and rests on Cimmerian basement, suggesting that the strata are intramontane deposits of the Cimmerian mountain chain in NE Iran. This interpretation has important implications concerning the position of the NW– SE-trending Eo-Cimmerian suture in NE Iran, which should be placed further SW than formerly assumed.

After the Early Permian detachment of the Cimmerian terranes from the northern margin of Gondwana, the Neotethys Ocean rapidly opened at the expense of the Palaeotethys (Stampfli & Borel 2002). The Iran Plate and Afghan Block as adjoining parts of the narrow, north-drifting ‘Cimmerian continent collage’ (Sengo¨r et al. 1988) closed the Palaeotethys, collided with Eurasia (Turan Plate) and created the Cimmerian mountain chain (e.g. Sengo¨r 1984, 1990). The exact timing of this orogenic event has remained controversial (e.g. Boulin 1988; Alavi et al. 1997; Sadi et al. 1997; Stampfli & Borel 2002; Seyed-Emami 2003), but it probably occurred during the Late Triassic (Fu¨rsich et al. 2009a). The position of the Palaeotethys suture is inferred to lie north of the present-day Alborz Mountains, running ESE between the Koppeh Dagh and the Binalud Mountains into northern Afghanistan (Boulin 1988; Alavi 1991, 1996; Alavi et al. 1997; Saidi et al. 1997; Zanchi et al. 2006) (Fig. 1).

In the Binalud Mountains, the southeastern extension of the Alborz chain, Palaeozoic–Triassic meta-sediments and Triassic –Lower Jurassic granitoids are widely distributed (e.g. Alavi 1992) (Fig. 1). The Palaeozoic–Triassic meta-sediments are interpreted as Palaeotethys remnants and are overlain, along a NW–SE-striking fault zone (Shandiz–Sangbast Fault), by a thick and poorly dated siliciclastic unit variably attributed to the Lower or Middle Jurassic (e.g. Aghanabati & Shahrabi 1987). The purpose of the present paper is a first description and interpretation of these hitherto poorly understood post-Eo-Cimmerian deposits in the Binalud Mountains. This includes a formal lithostratigraphic subdivision of the succession, an analysis of the depositional environments of the formations, and the correlation to litho- and chronostratigraphic equivalent units of the Alborz Mountains, northern Iran (see Fu¨rsich et al. 2009a). Finally, the inferred

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 175–188. DOI: 10.1144/SP312.8 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Geological map of NE Iran around Mashad, with the location of principal study areas (star symbols). The measured sections are located mainly in the southern part of the Binalud Mountains (Arefi, Kurtian, Bazehowz, Derekhtoot Farm). See text for detailed co-ordinates.

geodynamic significance of the succession is briefly discussed. The study is based on four measured sections and observations made at several other outcrops in the Binalud Mountains.

Geological setting Northeast Iran is characterized by NW –SE-striking geological units in both, the Koppeh Dagh and Binalud Mountains (Fig. 1). The Binalud Mountains are the SE continuation of the Alborz Mountains of northern Iran (e.g. Lammerer et al. 1984; EftekharNezhad & Behroozi 1991; Alavi 1992), where the succession consists of Palaeozoic–Triassic metasediments and Triassic –?Lower Jurassic granitoids (Berberian & Berberian 1981). These are overlain by Jurassic sediments consisting of poorly dated non-marine successions, the equivalents of the Middle Jurassic (Upper Bajocian –Bathonian) Kashafrud Formation (Fig. 2) (Aghanabati & Shahrabi 1987; Taheri et al. 2009), and Upper Jurassic and Lower Cretaceous carbonates. The (Upper) Triassic– ?Lower Jurassic granites (e.g. Mashad Granite and equivalents in the area of Torbat-e-Jam

further SE) are collision-related (Karimpour et al. 2006) and were exhumed by Middle Jurassic times, when they acted as a sediment source (e.g. in the basal Kashafrud Formation north of Torbat-e-Jam). The Permo-Triassic meta-sediments exhibit flysch features such as turbiditic greywackes, olistoliths, deep-sea sediments (radiolarian cherts) and ultrabasic rocks interpreted as obducted remnants of the Palaeotethys accretionary prism (Alavi 1991). The Palaeotethys suture, placed by Alavi (1991, 1992) in the Mashad area, strikes NW–SE. The Koppeh Dagh, NE of the Palaeotethys suture, is part of the Eurasian Plate. However, outcrops of pre-Middle Jurassic rocks are rare, being largely restricted to the tectonic window of the Aghdarband area (Fig. 1) (Baud & Stampfli 1989; Baud et al. 1991; Ruttner 1991; Alavi et al. 1997). The 1500 m-thick Triassic succession of Aghdarband consists, in ascending order, of the Scythian – Lower Anisian Sefid Kuh Formation (grey, micritic limestones with andesitic pyroclastic rocks at the base) and the Upper Anisian –Lower Carnian Sina Formation (volcanic and volcaniclastic rocks), unconformably overlain by the dark shales and fine-grained sandstones of the coal-bearing

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deformation in the Koppeh Dagh area clearly postdates the Norian–?Rhaetian Miankuhi Formation and predates the coarse-crystalline post-orogenic leukogranites (such as the Torbat-e-Jam Granite), which were exhumed by early Middle Jurassic times.

Stratigraphy and depositional environments

Fig. 2. Generalized Lower–Middle Jurassic sedimentary succession and formations of the Shemshak Group in the Binalud Mountains, NE Iran.

Miankuhi Formation (Norian – ?Rhaetian). The succession is intruded by coarse-crystalline leukogranites (Torbat-e-Jam Granite) and overlain, with angular unconformity, by the Upper Bajocian– Bathonian Kashafrud Formation, and a thick sequence of Upper Jurassic and Cretaceous strata (Eftekhar-Nezhad & Behroozi 1993). The Triassic of Aghdarband is inferred to have been deposited on basement rocks of the Turan Plate in an Eo-Cimmerian back-arc basin (Baud & Stampfli 1989; Baud et al. 1991). The main Eo-Cimmerian

The Eo-Cimmerian metamorphics to the south, west and NW of Mashad are overlain by a variable succession of poorly dated siliciclastic rocks regarded as ‘Middle Jurassic’ on most geological maps (e.g. Aghanabati & Shahrabi 1987; Taheri & Ghaemi 1999) (Fig. 2). The contact is usually mapped as a thrust fault (e.g. Alavi 1992, p. 364, fig. 4), along a NW –SE-striking fault zone (Shandiz–Sangbast Fault). However, at several localities (e.g. at Derekhtoot Farm, N368050 3300 , E598360 3000 , or near Kurtian, N368100 2600 , E598300 5900 , south and SSW of Mashad) (Figs 3 and 4A– C) the contact is clearly an angular unconformity without any faulting. At Derekhtoot Farm, the beds below the Eo-Cimmerian unconformity dip 628 towards 2428 (WSW), whereas the post-Eo-Cimmerian succession dips with 458 towards 2258. At Kurtian, pre-EoCimmerian beds dip approximately 0558/2308, post-Eo-Cimmerian strata at approximately 0308/ 2258. The sedimentary contact between the metamorphic succession (interpreted by Alavi 1991, 1992 as Palaeotethys remnants) and the non-marine Jurassic succession has important implications for the geodynamic interpretation of the area (see below). The age of this non-marine succession, subdivided below into three formations (Fig. 2), is poorly constrained. The underlying Eo-Cimmerian Mashad Phyllite locally yielded questionable Rhaetian –early Liassic plant fossils (Javadi & Latifi 2006). Plant macrofossils from the nonmarine succession may indicate an Early –Middle Jurassic age (Khatonie Malayossefi 2000). The presence of granitoid pebbles in the conglomerates of the succession is also more compatible with a Jurassic than a Late Triassic age (in the Torbat-e-Jam area, post-orogenic leukogranites intrude the Norian– ?Rhaetian Miankuhi Formation). In the Fraizi area (northwestern Binalud, NW of Mashad), these rocks are overlain by marine rocks yielding Late Bajocian ammonites (Taheri et al. 2006, 2009) (Fig. 2). These stratigraphic relationships clearly indicate that those strata were deposited after the (Late Triassic) Eo-Cimmerian deformation and before the widespread marine transgression of the ‘mid-Bajocian’ Mid-Cimmerian event (see Fu¨rsich et al. 2009b). Thus, a Liassic–early Bajocian age for the non-marine sequence is inferred, equivalent to the the Jurassic part of the Shemshak

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Fig. 3. Reference section of the Arefi Formation measured near Kurtian (N368100 2600 , E598300 5900 –N368100 2500 , E598300 3400 ) showing the Derekhtoot and Kurtian members.

Group in the northern Alborz (Fu¨rsich et al. 2009a). Based on lithological and sedimentary features, as well as on the stratigraphic succession, three formations are differentiated and formalized within the Shemshak Group of the Binalud Mountains. The minimum thickness of the succession is 2000 m (Fig. 2).

Arefi Formation Type area and type section The type area is located near the small village of Arefi (N368080 4600 , E598320 3500 ), south of Mashad, where the Arefi Formation rests on steeply inclined Permian

SHEMSHAK GROUP OF THE BINALUD MOUNTAINS

Fig. 4. Continued.

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phyllites. Unfortunately, the contact is strongly faulted and obscured by younger deposits in this area. Thus, a reference section of the Arefi Formation was measured near Kurtian along the wadi (N368100 2600 , E598300 5900 –368100 2500 , E598300 3400 ; Fig. 3).

Thickness Up to 750 m (near Derekhtoot Farm; base at N368050 3300 , E598360 3000 , top at N368050 1500 , E598360 1100 ).

Subdivision The Arefi Formation is subdivided into two members: a lower Derekhtoot Member and an upper Kurtian Member (Fig. 3). Derekhtoot Member Lithology. The lower Derekhtoot Member consists of up to 750 m-thick, very coarse and chaotic cobble and boulder beds, megabreccias and conglomerates, consisting mainly of Permo-Triassic metamorphic and granitoid rocks, quartzites and milky quartz (Figs 3 and 4A–C, E). Boulders may reach up to 5 m in diameter. The largest boulders are commonly granitoids and quartzites, the latter often exhibiting a brown weathering rim. Further components include andalusite schists, garnet schists, micaschists, milky quartz, black cherts, dark phyllites and meta-sandstones. In many places, the granitoids do not occur at the very base of the Derekhtoot Member. Rounding and sorting is very variable, ranging from poor to moderately sorted and rounded. The rocks are predominantly clast-supported (Fig. 4E), but occasional matrix-supported beds occur. The matrix is a poorly sorted gravelly greywacke with platy phyllite granules and pebbles, or a gravelly mudstone. Colours vary from greenish-grey to brown, occasionally with a touch of red. Bedding, if present, is only poorly defined, cross-bedding was not observed in the coarse-grained beds. On a decametre-scale, sharp- to erosively based fining-upwards cycles occur, comprising coarse boulder beds at the base grading into finer conglomerates and thin intercalations of pebbly, coarse arkosic sandstones with occasional decimetre-scale trough cross-bedding.

Kurtian Member Lithology. The upper Kurtian Member consists of more mature, clast-supported polymict conglomerates (.300 m in thickness, Fig. 3). It contains a diverse suite of well-rounded cobbles and pebbles made of metamorphic and igneous rocks (different schists and gneisses, migmatites, phyllites, metasandstones, quartzites, milky quartz, granitoids), similar in composition to the Derekhtoot Member. Maximum diameter, however, is much smaller and may reach 30 cm, but the mean size is 5 – 7 cm. Occasionally, wood logs occur. The rounding and sorting of clasts is moderate–good and they are set within a greenish-grey coarse- to mediumgrained greywacke –arkosic sandstone matrix. The conglomerate beds are commonly 1–5 m in thickness, sometimes lenticular, with sharp bases, and often exhibit large-scale trough cross-bedding and clast imbrication along foresets (Fig. 4D), both indicating a southwesterly palaeo-transport direction. The conglomerates are intercalated with decimetre- to metre-scale fine- to coarse-grained, greenish greywackes –sub-arkoses. They are structureless or may show large-scale trough crossbedding or ripple cross-lamination, and yield fine plant debris, more rarely complete leaves. Conglomerates and sandstones form characteristic metre-scale fining- and thinning-upwards cycles. Depositional environments. The sedimentary succession of the Arefi Formation indicates the erosion of a metamorphic basement to the east to NE of the Binalud Mountains, representing the core of the former Cimmerian Mountains. The large maximum clast size, poor rounding and sorting, as well as the chaotic fabric of the sediments of the Derekhtoot Member, indicate a very high relief, steep gradients and a short distance of transport. The rocks were deposited by rock-fall avalanches (chaotic boulder breccia), by debris flows (matrix-supported conglomerates) and in ephemeral stream channels (massive –crudely bedded, erosional conglomerates) of proximal– middle alluvial fans (e.g. McGowen & Groat 1971; Nilsen 1982). The common brown weathering rims and occasional brownish-red colours supports the inference of a semi-arid climate. The

Fig. 4. (Continued) Field aspects of the Arefi Formation. (A) Boulder beds of the Arefi Formation (Derekhtoot Member) near Derekhtoot Farm south of Mashad. Maximum boulder diameter is 4 m. (B) Contact of Permo-Triassic meta-sediments (right) and boulder beds of the Arefi Formation near Kurtian. Boulders (arrowed) reach 2.5 m in diameter. (C) Angular unconformity between metamorphic Permo-Triassic rocks (lower-right part of image with a 30 cm-long hammer) and the lower Arefi Formation, the latter consisting of poorly sorted cobble and boulder breccia. (D) Upper part of the Arefi Formation consisting of crudely bedded and large-scale trough cross-bedded, clast-supported, pebble-size conglomerates. Hammer (length 30 cm, arrowed) is given for scale. (E) Close-up view of finer-grained breccia of the lower part of the Arefi Formation consisting of poorly sorted pebbles and cobbles derived from Permo-Triassic meta-sediments and granitoids set in an immature matrix of arkosic sandstone and angular phyllite granules.

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broadly similar facies and thickness of the unit along the NW–SE-striking Shandiz–Sangbast Fault suggests fault-controlled deposition in the form of laterally extensive, overlapping fan wedges parallel to the fault (bajada deposits). The decametre-scale fining-upwards cycles are inferred to have formed by rapid pulses of faulting (uplift of source area), followed by retreat of the scarp front and lowering of relief in the highlands (e.g. Heward 1978a; Ethridge 1985). The better sorting and rounding, smaller grainsize, sedimentary structures, clast-supported fabric and small-scale fining-upwards sequences suggest a lower alluvial fan –proximal braided river origin for the conglomerates of the Kurtian Member (e.g. Miall 1985, 1996). They are the downstream equivalents of the coarse-grained Derekhtoot Member, indicating longer transport and decreasing relief energy. The conglomerate beds probably represent multistorey gravel units, originating as channel bars (e.g. Miall 1985). The thin lenses of sand were deposited in abandoned channels or are sand wedges at the edge of bars. The predominance of greenish colours and plant remains indicate more humid conditions for the conglomerates of the Kurtian Member of the Arefi Formation. In terms of lithofacies and bedding, the conglomerates of the Kurtian Member are very similar to those of the Javarherdeh Formation of the northern Alborz (Fu¨rsich et al. 2009a) and to intramontane molasse conglomerates of the Prado Formation from the Upper Carboniferous of northern Spain (pers. observ.; see also Heward 1978b).

Bazehowz Formation Type area and type section The type area of the Bazehowz Formation is around the small village of Bazehowz, approximately 25 km south of Mashad (Fig. 1). The base of the type section (Fig. 5) is at N368030 2300 , E598330 2500 , and the top is at N368030 3600 , E598390 1200 . The beds dip approximately 408 towards the west.

Thickness No complete section could be measured owing to the relatively poor exposure of the strata and internal faulting and folding. The type section (Fig. 5) has a thickness of 260 m and represents only a part of the succession. However, large areas of the outcrop belt of the Shemshak Group south, west and NW of Mashad are covered by strata of the Bazehowz Formation, and a minimum thickness of 1000 m is assumed.

Fig. 5. Type section of the Bazehowz Formation (N368030 2300 , E598330 2500 – N368030 3600 , E598390 1200 ), representing meandering stream and alluvial plain deposits (see Fig. 3 for the key to symbols).

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quartz-rich with moderate textural maturity, and often contain granules and pebbles of milky quartz, with quartzite and black cherts being subordinate constituents. Conglomerate beds are clast-supported and also consist predominantly of moderately to well-rounded and -sorted milky quartz granules and pebbles, and are characterized by large-scale trough cross-bedding (Fig. 6). The fine-grained interbeds (argillaceous –finesandy silts) are dark grey– greenish grey. Thin (centimetre-thick) beds of ripple cross-laminated or parallel-laminated, fine-grained sandstone are occasionally intercalated. The succession is rich in plant remains, and numerous well-preserved plants (horsetails, ferns, cycads, conifers) were collected from the Bazehowz area. Layers of carbonaceous silts and coal seams occur, but workable thicknesses of more than 1 m are rarely reached (Fig. 7A).

Depositional environments

Fig. 6. Section of part of the fluvial Bazehowz Formation south of Derekhtoot Farm (base at N368040 5500 , E598350 2700 , top at N368050 0500 E598350 3200 ) showing pebbly channel sandstones alternating with carbonaceous siltstone and sandstone of floodplain origin (see Fig. 3 for the key to symbols).

Lithology The Bazehowz Formation consists of argillaceous silt– fine-sandy silt alternating with large-scale trough cross-bedded, medium- to coarse-grained sandstone, pebbly sandstones and fine-grained conglomerates forming beds up to 10 –30 m in thickness (Figs 5 and 6). These beds exhibit invariably sharp, often erosional, bases and numerous internal erosion surfaces. Foresets of trough cross-beds may reach a thickness of 1–3 m, being replaced up-section by small-scale trough cross-bedding and ripple-lamination. The sandstones are commonly

The Bazehowz Formation consists essentially of stacked, metre- to decametre-scale fining- and thinning-upwards cycles, starting at the base with an erosional surface overlain by trough crossbedded pebbly sandstones–fine-grained conglomerates grading into sand- and siltstones. The sediments were deposited in the middle reaches of a braided fluvial system dominated by pebbly sands. In such a setting, downstream and laterally migrating sand and gravel bars form planar- or trough cross-bedded units, while lateral migration and sudden abandonment of channels due to avulsion generates fining-upwards channel-fill sequences similar to those described earlier (Miall 1985). The finer sediments are thought to be deposits of wet interchannel areas, yielding common plant remains and occasional coal seams. An interpretation of the fining-upwards cycles as point-bar deposits of a large-scale meandering, coarse-grained fluvial system is less likely, although coarse-grained gravelly streams tend to have somewhat lower sinuosities than sandy streams of similar size (McGowen & Garner 1970; Collinson 1986; Miall 1996). The dark colours of the floodplain/ overbank sediments, abundant plant remains, and the presence of carbonaceous beds and the coal

Fig. 7. (Continued) Field aspects of the Bazehowz and Aghounj formations. (A) Bazehowz Formation south of Derekhtoot Farm with pebbly sandstone beds of fluvial channel origin (fcs) embracing floodplain fines (sandstone– siltstone intercalations, carbonaceous siltstones and coal seams). (B) Reference section of the Aghounj Formation NW of Kurtian showing approximately 100 m of thick-bedded conglomerates consisting of well-rounded milky quartz granules and pebbles alternating with siltstones and fine- to medium-grained sandstones. (C) Close-up view of a conglomerate bed of the Aghounj Formation NW of Kurtian showing the clast-supported fabric with abundant well-rounded milky quartz granules and pebbles, as well as rare platy, dark-coloured phyllite pebbles. The hammer length is 30 cm. (D) Sharp-based, fining-upwards conglomerate– sandstone channel-fill sequence of the Aghounj Formation with large-scale trough cross-bedding in its lower part (near hammer; section NW of Kurtian, the hammer is 30 cm long).

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Fig. 7. Continued.

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Aghounj Formation Type area and type section The type area of the Aghounj Formation is around the small village of Aghounj, south of Mashad, where the formation crops out extensively in the centre of an anticline, the core of which now forms the highest elevation (1540 m) in the study area (N368040 480 , E598330 4200 for the lower part of the formation, and N368040 170 , E598320 4400 for the upper part). However, the top of the formation is not exposed there and outcrops are poor, making it difficult to measure a section. Thus, a reference section was measured in a gorge near Kurtian (Figs 7B and 8: base at N368100 2900 , E598300 1700 , dip 248ESE) where the strata are well exposed. However, the top of the formation here could not be measured, but it appears to be followed by finer-grained strata.

Thickness The Aghounj Formation is estimated to be at least 400 m thick.

Lithology The Aghounj Formation consists of metre-scale beds and decametre-scale amalgamated bedsets, some of them lenticular, of poorly–well-rounded quartz conglomerates with only very minor phyllite, chert and other metamorphic components (Fig. 7B, C). Sorting is generally poor. Grain size normally does not exceed pebble size (mean diameter c. 3 cm). Occasionally, cobbles up to 15 cm in diameter occur, and sorting is rather poor. Most beds exhibit large-scale trough cross-bedding, sharp, erosional bases and a weak fining-upwards trend (Fig. 7D). Occasionally, load structures and injection structures are present at the base. Wood fragments and logs are common, the latter being concentrated at the base of the conglomerate units. Between the conglomerate beds and bedsets, fine-grained, partly ripple-cross-laminated sandstones containing some plant debris and rare bioturbation occur. These finergrained interbeds are usually thinner than the conglomerates but are often poorly exposed. Fig. 8. Reference section of the Aghounj Formation NW of Kurtian (base at N368100 2900 , E598300 1700 ) showing channel-fill conglomerates alternating with siltstones and sandstones (see Fig. 3 for the key to symbols).

seams suggest humid climate conditions for the deposition of the Bazehowz Formation by perennial streams (cf. Collinson 1986).

Depositional environments Based on the coarse grain size, sedimentary structures, the lenticular shape and amalgamated architecture of conglomerate beds, as well as small-scale fining-upwards sequences, the Aghounj Formation is interpreted as the product of a proximal braided river system (Miall 1985, 1996). The compositional maturity of the gravel deposits is high;

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textural maturity is less so. The Aghounj Formation indicates erosion of a deeply denudated (Cimmerian) basement with slightly higher fluvial gradients than in the underlying Bazehowz Formation. A lithologically similar, 150 m-thick quartz conglomerate was recorded from the upper part of the Shemshak Group (?Aalenian) of the eastern Alborz (Gheshlagh) by Stampfli (1978).

Discussion and conclusions The post-Eo-Cimmerian and pre-Upper Bajocian non-marine sedimentary succession of the Binalud

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Mountains of NE Iran can be subdivided into three formations: i.e. the Arefi, the Bazehowz and the Aghounj formations (Fig. 9). Based on sparse biostratigraphic data (mainly plants, ammonites in the overlying marine strata of the Kashafrud Formation), an Early–early Middle Jurassic (Early Bajocian) age can be inferred. Unfortunately, no further chronostratigraphic data are available for the three formations, which replace each other up-section, even though their boundaries are not known (except for the base of the Arefi Formation). The lowermost unit, the Arefi Formation, is subdivided into a lower Derekhtoot Member and an

Fig. 9. Lithostratigraphic framework of the Shemshak Group of the Binalud Mountains and correlation with the succession in the Alborz Mountains of northern Iran (cf. Fu¨rsich et al. 2009a).

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upper Kurtian Member. The Derekhtoot Member (up to 750 m thick) consists of very coarse-grained, chaotic boulder beds, breccias and conglomerates representing rock-fall deposits and proximal – middle alluvial fans, deposited as bajada along a steep fault scarp (Shandiz –Sangbast Fault) under a semi-arid climate. The succeeding Kurtian Member (.300 m) consists of finer-grained and better-rounded conglomerates, and reflects deposition of a proximal perennial braided river system. The clasts of both members of the Arefi Formation are dominated by a diverse suite of metamorphic and magmatic components. The overlying Bazehowz Formation is more than 1000 m thick and consists of stacked, decametre-scale channel-fill cycles of the middle reaches of a pebbly–sandy braided river system. Abundant organic matter, plant fossils and greenish-grey colours in floodplain fines indicate fairly humid conditions for the Bazehowz Formation. The uppermost unit, the Aghounj Formation, consists at least of 400 m of granule- to pebble-size, thickbedded and large-scale trough-cross-bedded quartz conglomerates, and sandy interbeds of a proximal–medial braided river system. This unit is inferred to be overlain by marine Upper Bajocian –Bathonian strata of the Kashafrud Formation (see Taheri et al. 2009). The Jurassic non-marine succession of the Binalud Mountains reflects the erosion of a highrelief source area down to the metamorphic basement. Lateral facies variations indicate that the clastic material was transported from NE to SW (vaguely supported by a few palaeo-current data). These data, combined with the clast spectrum, allows for an interpretation of the strata as the Cimmerian molasse deposits in northeastern Iran. The succession represents a large-scale fining-upwards megacycle from the Derekhtoot Member to the Kurtian Member of the Arefi Formation, and into the Bazehowz Formation. The Aghounj Formation reflects a renewed increase in fluvial gradients, which perhaps was associated with an early Middle Jurassic tectonism known in Iran as the Mid-Cimmerian tectonic event (Fu¨rsich et al. 2009b). Lithologically similar and stratigraphically broadly equivalent quartz conglomerates of inferred Aalenian age are also known from the eastern Alborz, the westward continuation of the Binalud Mountains (Stampfli 1978). The absence of marine Toarcian– Aalenian strata from the easternmost Alborz (Stampfli 1978) and the Binalud, in contrast to the western and central Alborz where they are thick and well documented (Shirindasht and Fillzamin formations; Fu¨rsich et al. 2005) (Fig. 9) support the view of a zip-like, eastwardproceeding diachronous opening of a Neotethys back-arc basin in northern Iran during the Jurassic (Fu¨rsich et al. 2005, 20009a, b). This basin developed into the strongly subsiding South Caspian

Basin (Brunet et al. 2003) and continued eastwards into the Late Bajocian–Bathonian Kashafrud Basin of the Koppeh Dagh Mountains in NE Iran (Taheri et al. 2006, 2009). The Lower– lower Middle Jurassic succession of the Binalud Mountains is correlated with the Jurassic part of the Shemshak Group of the Alborz Mountains (Fu¨rsich et al. 2009a) (Fig. 9). In the Alborz Mountains, coarse-grained Lower Jurassic siliciclastic rocks of the Javaherdeh and Alasht formations likewise indicate the rapid uplift and denudation of the Cimmerian mountain chain with a north –south-directed transport of the erosional products. The Alborz units represent the Cimmerian molasse of the peripheral foreland basin (Fu¨rsich et al. 2009a). However, they rest on Upper Triassic synorogenic (‘flysch’) sediments of the foredeep and not on metamorphic Cimmerian basement, as in the case of the Binalud succession, and the conglomerates of the Javaherdeh Formation are not as proximal as the boulder beds and conglomerates of the Derekhtoot Member of the Arefi Formation. The Binalud succession, therefore, represents deposits in closer proximity to the Cimmerian mountain chain. The Binalud Basin is interpreted as an intramontane basin developing during the late phase of the Cimmerian orogeny. Based on their extreme textural and compositional immaturity, and numerous internal (local–regional) erosional unconformities, the Derekhtoot Member of the Arefi Formation may also represent a synorogenic wedge-top deposit (cf. DeCelles & Giles 1996; Chiang et al. 2004). However, the Arefi – Bazehowz formations are molasse-type sediments and form an overall fining-upwards megasequence that records the opening of a basin by extensional tectonics. The laterally extensive, bajada-like distribution of the conglomerates parallel to the NW– SE-striking Shandiz–Sangbast Fault and the sediment transport towards the SW suggest that the fault was synsedimentarily active as a normal fault, forming the northeastern boundary of the basin. Thus, the Binalud succession may have been deposited in a NW–SE elongated half-graben basin developing in response to orogenic collapse (Fig. 10). An interpretation of the Binalud succession as an intramontane deposit (i.e. formed within the Cimmerian mountain belt) has important geodynamic implications concerning the position of the Eo-Cimmerian suture in NE Iran, which is usually placed with a NW–SE strike in the Mashad area (e.g. Alavi 1991, 1992, 1996; Alavi et al. 1997): the suture must be placed further SW than formerly assumed, as intramontane basins develop within the thrust belt on top of the overriding plate (e.g. Einsele 2000). This is also suggested by the position of the Eo-Cimmerian granites of Mashad and Torbat-e-Jam (Fig. 1), which should be an upper plate feature. Their intrusion may

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Fig. 10. Sketch showing the Early Jurassic Binalud Basin (NE Iran) as an intramontane basin within the Cimmerian Mountains. See the text for further explanations.

have been triggered by slab break-off in the course of the Eo-Cimmerian collision, which is inferred to have occurred at the Triassic– Jurassic boundary (Fu¨rsich et al. 2009a; see Hann et al. 2003 for a Variscan example). The post-Norian and preMiddle Jurassic age of the Torbat –Jam Granite corroborates this scenario. We acknowledge constructive reviews by B. Turner (Durham) and T. Voigt (Jena). The German Research Foundation (DFG) sponsored F. T. Fu¨rsich, M. Wilmsen and J. Taheri for investigations in NE Iran (FU 131/ 32-2). We furthermore acknowledge logistic support by the Geological Survey of Iran (GSI) and financial support by the Middle East Basin Evolution (MEBE) Programme. H. deWall (Wu¨rzburg) and J. Kley (Jena) are thanked for discussion of the geodynamic situation of the Binalud succession.

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University of Texas, Report of Investigations, 72, 1– 57. M IALL , A. D. 1985. Architectural-element analysis: a new method of facies analysis applied to fluvial deposits. Earth Science Reviews, 22, 261– 308. M IALL , A. D. 1996. The Geology of Fluvial Deposits: Sedimentary Facies, Basin Analysis and Petroleum Geology. Springer, Berlin. N ILSEN , T. H. 1982. Alluvial fan deposits. In: S CHOLLE , P. A. & S PEARING , D. R. (eds) Sandstone Depositional Environments. AAPG, Memoir, 31, 49–66. R UTTNER , A. W. 1991. Geology of the Aghdarband area (Kopet Dagh, NE-Iran). Abhandlungen der Geologischen Bundesanstalt Wien, 38, 7–79. S AIDI , A., B RUNET , M.-F. & R ICOU , L.-E. 1997. Continental accretion of the Iran Block to Eurasia as seen from Late Paleozoic to Early Cretaceous subsidence curves. Geodinamica Acta, 10, 189 –208. S ENGO¨ R , A. M. C. 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia. Geological Society of America, Special Paper, 195. S ENGO¨ R , A. M. C. 1990. A new model for the late Palaeozoic– Mesozoic tectonic evolution of Iran and its implications for Oman. In: R OBERTSON , A. H. F., S EARLE , M. P. & R IES , A. C. (eds) The Geology and Tectonics of the Oman Region. Geological Society, London, Special Publications, 49, 797–831. S ENGO¨ R , A. M. C., A LTINER , D., C IN , A., U STAO¨ MER , T. & H SU¨ , K. J. 1988. Origin and assembly of the Tethysides orogenic collage at the expense of Gondwana Land. In: A UDLEY -C HARLES , M. G. & H ALLAM , A. (eds) Gondwana and Tethys. Geological Society, London, Special Publications, 37, 119–181. S EYED -E MAMI , K. 2003. Triassic in Iran. Facies, 48, 91–106. S TAMPFLI , G. M. 1978. Etude ge´ologique ge´ne´rale de l’Elburz oriental au S de Gonbad-e-Qabus, Iran N-E. The`se no. 1868, Faculte´ des Sciences de l’Universite´ de Gene`ve. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters, 196, 17–33. T AHERI , J. & G HAEMI , F. 1999. Geological Map of Mashhad, in the Scale 1:100 000. Geological Survey of Iran, Tehran. T AHERI , J., F U¨ RSICH , F. T. & W ILMSEN , M. 2006. Stratigraphy and depositional environments of the Upper Bajocian– Bathonian Kashafrud Formation (NE Iran). Volumina Jurassica, 4, 105–106. T AHERI , J., F U¨ RSICH , F. T. & W ILMSEN , M. 2009. Stratigraphy, depositional environments and geodynamic significance of the Upper Bajocian–Bathonian Kashafrud Formation, NE Iran. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 205–218. Z ANCHI , A., B ERRA , F., M ATTEI , M., G HASSEMI , M. & S ABOURI , J. 2006. Inversion tectonics in central Alborz, Iran. Journal of Structural Geology, 28, 2023– 2037.

The Mid-Cimmerian tectonic event (Bajocian) in the Alborz Mountains, Northern Iran: evidence of the break-up unconformity of the South Caspian Basin ¨ RSICH1*, MARKUS WILMSEN1,2, KAZEM SEYED-EMAMI3 & FRANZ THEODOR FU MAHMOUD REZA MAJIDIFARD4 1

Geozentrum Nordbayern der Universita¨t Erlangen-Nu¨rnberg, Fachgruppe Pala¨oUmwelt, Loewenichstrasse 28, D-91054 Erlangen, Germany (e-mail: [email protected])

2

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany 3

School of Mining Engineering, University College of Engineering, University of Tehran, P.O. Box 11365-4563, Iran 4

Geological Survey of Iran, P.O. Box 131851-1494, Tehran, Iran

*Corresponding author (e-mail: [email protected]) Abstract: The Mid-Cimmerian tectonic event of Bajocian age can be documented all across the Iran Plate (Alborz Mountains of northern Iran, NE Iran, east-central Iran) and the southern Koppeh Dagh (northeastern Iran). In the Alborz area, the tectonic event consisted of two main pulses. A distinct unconformity (near the Lower– Upper Bajocian boundary) at or near the base of the Dansirit Formation is the sedimentary expression of rapid basin shallowing due to uplift and erosion. Another unconformity is developed in the early Upper Bajocian, close to or at the top of the Dansirit Formation. Locally, it is expressed as an angular unconformity due to block rotation and is overlain by a thin transgressive conglomerate followed by silty marls of the deep-marine Upper Bajocian–Callovian Dalichai Formation. This upper unconformity signals a rapid subsidence pulse. On the Tabas Block of east-central Iran, a single unconformity can be documented that is time-equivalent to those bounding the Dansirit Formation (i.e. ‘mid-Bajocian’). Local folding gives direct evidence of compressional tectonics, and conglomerates indicate subaerial denudation of older Mesozoic or Palaeozoic strata. After a stratigraphic gap, transgressive sediments of ?Late Bajocian–Bathonian age follow, suggesting a fusion of the lower and upper Mid-Cimmerian unconformities in east-central Iran. Along the southern margin of the Koppeh Dagh Mountains (NE Iran), a Late Bajocian subsidence pulse initiated the opening of the strongly subsiding Kashafrud Basin, an eastwards extension of the South Caspian Basin. In all of these areas, one phase of uplift and erosion took place followed by a pronounced pulse of subsidence running counter to trends of the eustatic sea-level curve. Thus, what is generally understood as the Mid-Cimmerian tectonic event is now thought to consist of a tectonic phase, confined to the Bajocian. This phase is explained as the expression of the onset of sea-floor spreading within the South Caspian Basin situated to the north of the present-day Alborz Mountains. This strongly subsiding basin developed close to the Palaeotethys suture during the Toarcian– Aalenian and went through a change from the rifting- to the spreading-stage during the Bajocian. The Mid-Cimmerian event therefore reflects the break-up unconformity of the South Caspian Basin.

Unconformities in the sedimentary record may have different causes. They can be the result of eustatic sea-level fluctuations, may reflect large-scale climatic changes or, most commonly, are the expression of tectonic movements at different scales and produced by a variety of processes. In the Late Triassic– Jurassic three significant unconformities of tectonic origin have been defined. Based on structural observations in Crimea, Stille (1924) recognized a Late

Triassic Early Cimmerian orogeny and a preTithonian Late Cimmerian orogeny. Ziegler (1975) defined a Mid-Cimmerian (Middle Jurassic) deformation phase, which can be widely recognized in northern Europe. These tectonic phases are widely used in the literature for time-equivalent events (e.g. Kyrkjebø et al. 2004). In the Mesozoic of Central and Northern Iran (Iran Plate; Fig. 1), the earliest and probably most

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 189–203. DOI: 10.1144/SP312.9 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Structural framework of Iran (A, above) and distribution of the Shemshak Group in the Alborz Mountains with locations of sections (B, below). (a) Damavand area (Fillzamin section); (b) Sharif Abad section; (c) Shahmirzad (Kuhe-Bashm-e Dehsufian) section; (d) Parvar area; (e) Tazareh section; (f) Jajarm area (including Golbini and Tappehe-Chefte-Qasi sections); and (g) Rian area.

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distinct of these unconformities is found around the Middle–Upper Triassic boundary. Although not an angular unconformity in most places, it can be easily recognized by the abrupt change in facies from platform carbonates (Shotori and Elikah formations) to siliciclastic rocks of the Shemshak Group (see Fu¨rsich et al. 2009) (Fig. 2). Moreover,

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at many localities across northern and east-central Iran, the unconformity is characterized by a karst morphology and the development of lateritic soils and bauxite deposits indicative of subaerial exposure. This unconformity is clearly the expression of orogenic movements, the Early Cimmerian orogeny, and has been interpreted as the

Fig. 2. Main lithostratigraphic units of the Upper Triassic–Jurassic succession of northern and central Iran. A, Early Cimmerian unconformity; B, Mid-Cimmerian unconformity.

Fig. 3. Sections through the Dansirit Formation in the central and eastern Alborz. Key to the symbols applies also to Figures 5 and 7.

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signal of the initial collision of the Iran Plate with the Turan Plate of Eurasia (Fu¨rsich et al. 2009). In the Alborz Mountains, the thick pile of siliciclastic rocks of the Upper Triassic –lower Middle Jurassic Shemshak Group is terminated by yet another unconformity and an equally sharp change back to a carbonate regime, the carbonate platform –basin complex of the Middle–Upper Jurassic Dalichai and Lar formations. This latter unconformity formed around the Early– Late Bajocian boundary. Its signal – a sharp change in facies, commonly associated with the formation of conglomerate layers – can be documented in all areas of the former Iran Plate and beyond, from Kerman in the south to the Koppeh Dagh in the north, and has been interpreted as tectonic in origin. Seyed-Emami & Alavi-Naini (1990) named this tectonic event Lutian, but subsequently the already existing term ‘Mid-Cimmerian tectonic event’ became entrenched in the literature on the area (Aghanabati 1998; Seyed-Emami et al. 2004). Rarely, the Mid-Cimmerian unconformity is developed as an angular unconformity (e.g. at Mazinu, 60 km WSW of Tabas: Wilmsen et al. 2003; SeyedEmami et al. 2004, plate 1, fig. 1). Whereas the tectonic origin of the Mid-Cimmerian unconformity is beyond doubt, no satisfactory explanation has been offered so far as to the precise causative processes. Moreover, as will be shown later, there are, in fact, two closely related unconformities that in the past have not been differentiated. It is the purpose of this paper to describe and analyse the sedimentary framework of these unconformities and to offer a geodynamic interpretation by relating them to the onset of sea-floor spreading in the South Caspian Basin, to the north of the present-day Alborz Mountains.

Geological framework Iran consists of a mosaic of plates and microplates (Fig. 1). Among them is the Iran Plate, part of the Cimmerian continent collage (Sengo¨r et al. 1988) that became detached from the northern margin of Gondwana in the Early Permian and, by moving northwards, collided with Eurasia. In the case of the Iran Plate, collision with Eurasia took place around the Middle–Late Triassic, thereby closing the Palaeotethys in the north (e.g. Stampfli & Borel 2002). The Iran Plate consists of several fragments, mainly the Central –East Iranian Microcontinent, the Alborz Mountains and NW Iran, which apparently moved northwards together. While drifting northwards during the Triassic, the Iran Plate was covered by extensive carbonate platforms (Shotori and Elikah formations; Fig. 2). After collision and formation of the Cimmerian

mountain chain (e.g. Sengo¨r 1990) the carbonate regime was replaced by a thick pile of siliciclastic sediments, the Upper Triassic –lower Middle Jurassic Shemshak Group (Fu¨rsich et al. 2009) (Fig. 2). The lower part of this group corresponds to Cimmerian flysch and molasse deposits, the upper part to the fill of an extensional basin. After the Bajocian Mid-Cimmerian tectonic movements, the siliciclastic regime of the Shemshak Group was replaced by a carbonate system (e.g. Dalichai, Lar, Esfandiar Limestone and Qal-eh-Dokhtar formations) that covered large parts of the Iran Plate until the end of the Jurassic System.

Facies and environmental framework of the Dansirit Formation The two unconformities occur at or close to the base and top, respectively, of the Dansirit Formation, the youngest formation within the Shemshak Group (Fu¨rsich et al. 2009). Seven sections have been measured in detail (Figs 1 and 3); in addition, several others have been investigated, but not logged (Fig. 1). The Dansirit Formation (Bajocian) overlies the Fillzamin Formation (Aalenian), a monotonous succession that in the west of the study area consists of dark-grey mudstones with scattered calcareous and ferruginous concretions. In the east the formation is slightly coarser grained and is composed of poorly developed coarsening-upwards parasequences. The succession has been interpreted by Fu¨rsich et al. (2005) as outer shelf/basinal in origin, deposited well below storm-wave base. In most sections (Fig. 3) the Fillzamin Formation coarsens towards the top, and intercalated sandstones show hummocky cross-stratification (HCS) and parallel lamination. The clear evidence of a storm-influenced regime indicates gradual shallowing of the basin. At most localities the boundary to the Dansirit Formation coincides with the position of the lower unconformity, labelled ‘1’ in Figure 3. The Dansirit Formation has everywhere a shallow-water, nearshore to coastal/delta-plain character (Fu¨rsich et al. 2009). In the eastern Alborz Mountains (e.g. at Jajarm, Golbini, Tappehe Chefteh Qazi and Parvar) it consists of marine sandstone. Towards the west (e.g. at Tazareh, Sharif-Abad and Fillzamin) the formation is heterolithic, being composed of argillaceous silt, silt/siltstone and sandstone of partly marine, partly non-marine origin. At Jajarm in the eastern Alborz Mountains (Fig. 3F) 188 m of thick-bedded, fine-grained sandstones, most of them with trough cross-stratification, indicate a high-energy environment above fairweather wave-base. Increased bioturbation and, at

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some levels, a slightly finer grain size in the upper third of the section point to more tranquil conditions. Wood fragments and plant remains are signs that the coastline was not far away. We interpret the succession to record an upper delta-front regime, which deepened slightly up-section. Two kilometres further east, at Golbini, the timeequivalent succession is finer grained (silt to finegrained sand), only 30 m thick, highly bioturbated and arranged in small coarsening-upwards cycles. Marine fossils are common, and include bivalves (e.g. Ctenostreon, Isognomon, Chlamys, Trichites and Pholadomya), gastropods, corals, serpulids and belemnites. Most probably, the Dansirit Formation at this locality represents a somewhat protected, but fully marine, embayment. Some 20 km further west, at Tappehe Chefteh Qazi, the Dansirit Formation consists of approximately 30 m of thick-bedded, fine-grained sandstone (Fig. 4A). Near the base structureless beds alternate with planar cross-bedded units, which prevail up-section. All foresets dip in the same direction (WNW; Fig. 4B). The top 1.5 m are poorly sorted, medium- to coarse-grained sandstone; the grain size decreases again at the bioturbated very top. The sandstone package can be interpreted as a submarine coastal-bar complex. The unidirectional cross-stratification may be due to transport by longshore currents. The most expanded section of the Dansirit Formation occurs at the Tazareh coal mine (313 m; Figs 3E and 4C). It starts with 29 m of very coarsegrained, poorly sorted, arkosic, lenticular sandstone packages, each 5– 7 m in thickness, with large-scale trough cross-bedding and erosional bases. The lowermost of these packages contains chaotically arranged tree trunks (Fig. 4D) and cuts down into marine sediments. This unit, interpreted as fluvial channel fill, is overlain by sharp-based, finingupwards sandstones and argillaceous silt/sandstone alternations with rootlet horizons and coal seams, which correspond to small fluvial channels, floodplain sediments and coastal marsh land. Rare marine incursions are documented by bivalves such as Pholadomya and Ceratomya, and by the trace fossil Rhizocorallium. Approximately 100 m above the base, the character of the sediment changes: coarsening-upwards sequences predominate, and trace fossils (Ophiomorpha, Taenidium and Chondrites), repeated hummocky crossstratification and well-sorted sandstone packages with low-angle lamination record storm-influenced upper offshore –foreshore environments. In the upper third of the Dansirit Formation, a brief return to continental conditions is shown by dm-thick coal layers and a rootlet horizon. The overlying cross-bedded sandstones, occasionally with lenses of well-rounded quartz granules and

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milky quartz pebbles, are interpreted as documenting a return to high-energy, nearshore conditions (sand and gravel bars). Laterally, the thickness of the conglomerate lenses varies considerably, reaching up to 80 cm, as does the maximum pebble size (1–3 cm). Up-section, this unit grades into bioturbated, bedded, fine-grained sandstones. The topmost bed is a calcareous, fine- to mediumgrained sandstone with occasional belemnites and bivalves. At Parvar, situated further west and north (Fig. 1), the Dansirit Formation is composed of fine- to rarely medium-grained sandstones with rare intercalations of silty fine-grained sandstone (Fig. 3D). Only part of the succession exhibits primary sedimentary structures (large-scale trough cross-bedding, HCS), bioturbation being widespread. Discrete trace fossils include Thalassinoides, Planolites, Rhizocorallium and Lophoctenium. Plant and wood debris is common. The trace fossils and scattered bivalves point to a shallow-marine setting, most probably a lower– upper delta front. A 30 cm-thick layer of coal in the middle part of the formation is most probably allochthonous in nature not being associated with a root horizon. The Dansirit Formation at Kuhe-Bashm-e Dehsufian (Fig. 3C) is relatively expanded (261 m thick), but half of the section is covered by scree. The succession appears to be predominantly sandy, partly arkosic, and the grain size varies between fine-sandy silt and gravel. The lower 60 m are composed of thickening- and coarseningupwards sequences. Sedimentary structures are common (ripple lamination, parallel lamination, HCS and large-scale trough cross-bedding), as are signs of bioturbation (Rhizocorallium irregulare) and rare shell debris. These features point to a shallow-marine environment. The exposed upper part starts with well-sorted, low-angle laminated fine-grained sandstones with Ophiomorpha, which we interpret as upper foreshore deposits. Trough crossbedded medium- to very-coarse-grained ferruginous sandstones, in places gravelly and with lenses of quartz conglomerates and wood fragments, are thought to represent fluvial deposits. The succession at Sharif-Abad (Fig. 3B), approximately 9 km further SSE of Kuhe-Bashm-e Dehsufian, differs markedly in being very thin (10.5 m). As is the case at most localities, the underlying Fillzamin Formation coarsens towards the top (Fig. 4E), but the lower unconformity is drawn at the change from interbedded bioturbated silty finegrained sandstones and ripple-laminated sandstones to fine- to medium-grained sandstones with flaser bedding, a coal lenticle and a reddish ferruginous layer, which we interpret as pedogenic in origin. The following 5 m start with a 0.5 m-thick, sharpbased, trough cross-bedded quartz conglomerate,

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Fig. 4.

Continued.

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fining upwards into fine-grained sandstone (Fig. 4F). The conglomerate consists of well-rounded milky quartz, lydite and, more rarely, sandstone pebbles up to 4 cm in diameter. Lenticular concentrations of bivalve shells (large Myophorella and Isognomon), abraded coral heads and belemnites indicate a marine origin, and the unit can be interpreted as a transgressive lag with a considerable hiatus at its base. The overlying fine-grained sandstone is bioclastic, exhibits low-angle cross-stratification and, at the top, there is yet another shell-cum-pebble lag with serpulids, bivalves (e.g. large oysters), ammonites, belemnites and brachiopods. About 1km east of this section a fragment of the ammonite Kumatostephanus? has been found in sandy limestones of the topmost Fillzamin Formation by one of the authors (K. Seyed-Emami), along with bivalves and belemnites. The ammonite indicates the Propinquans to Humphriesianum zones of the upper Lower Bajocian. At Rian, approximately 35 km NE of Semnan (N358460 0300 , E538390 0300 ), the upper part of the Shemshak Group is even more strongly condensed (Fig. 5). The Upper Toarcian– Bajocian succession comprises less than 60 m, of which the lower 45 m consist of mixed carbonate –siliciclastic sediments (silty marl, sandy bio-packstones and rudstones), with 17 m of coral meadow and patch reefs in the upper part (Pandey & Fu¨rsich 2006). The top 8.5 m of the group are composed of large-scale trough cross-bedded sandstone with lenses of quartz conglomerates at the base (Fig. 4G), which overlie, with distinct unconformity, the debris facies (biorudstones) that terminated the coralreef growth. This debris facies records shallowing, which led to the destruction of the patch reefs. The conglomeratic sandstones contain oyster debris and are, thus, clearly of marine origin. They document a transgressive pulse, during which coarse-grained fluvial material, the sedimentary expression of a significant phase of erosion most probably associated with uplift, was reworked. Another transgressive pulse is recorded by a 5 cm-thick layer of bioclastic

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Fig. 5. The Mid-Cimmerian unconformity at Rian (see also Fig. 4G). For the key to the symbols see Figure 3.

calcareous sandstone on top of the trough crossbedded sandstone unit. The boundary to the overlying Dalichai Formation is less abrupt than at most other localities: the basal beds of the Dalichai Formation consist of 3 m of hummocky crossstratified fine-grained sandstones interbedded with layers of fine-sandy marly silt, before the characteristic uniform silty marl facies sets in. The westernmost section investigated is situated west of Damavand at Fillzamin (Figs 3A and 6D). The lowermost few metres of cross-bedded mediumgrained sandstone (basal Dansirit Formation; Fig. 6F) are marine in origin, as is evidenced by the trace fossil Ophiomorpha. The unconformity

Fig. 4. (Continued) The Mid-Cimmerian unconformities in the Alborz Mountains, Iran. (A) Sharp contact (arrow) of the soft siltstones of the Fillzamin Formation and medium- to coarse-grained sandstones of the Dansirit Formation (lower unconformity, Tappehe-Chefte-Qasi, Jajarm area). (B) Unidirectional, tabular cross-bedding in sandstones of the Dansirit Formation (Tappehe-Chefte-Qasi, Jajarm area, exposed thickness c. 30 m). (C) Type section of the Dansirit Formation north of Tazareh. The cliff in the background is formed by Upper Jurassic limestones (Lar Formation). (D) Large wood log from the lower part of the Dansirit Formation at Tazareh. Hammer (30 cm long) for scale. (E) Marly Fillzamin Formation (c. 100 m thick) at Sharif Abad, overlain by sandstones of the Dansirit Formation (lower unconformity arrowed). (F) Upper unconformity (arrowed): pebbly transgressive sandstones of the upper Dansirit Formation (left) truncating red-coloured (pedogenic?) fine- to medium-grained sandstones of the lower Dansirit Formation (right). (G) The two Mid-Cimmerian unconformities are fused into one surface at Rian where fossiliferous marly–silty bioclastic limestones are truncated by transgressive quartz conglomerates and trough cross-bedded sandstones (cf. Fig. 5; hammer handle is 40 mm thick). (H) Well-rounded and sorted fine-grained quartz conglomerate of the Dansirit Formation west of Shahmirzad. Note the well-rounded quartz granules (pencil sharpener is 25 mm long).

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Fig. 6.

Continued.

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is placed at the base of the overlying coarse-grained sandstone, which exhibits signs of lateral accretion (epsilon stratification) and rip-up clasts, and which fines upwards. It represents a meandering river deposit, as does another sandstone body higher up. Coal layers, irregular iron crusts and rootlet horizons also point to a flood-plain– low-lying coastal-plain environment, whereas thinner sandstones with Ophiomorpha record short-lived marine ingressions.

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environments to distinctly shallower settings, such as upper delta front (e.g. Jajarm) or delta top (e.g. Tazareh). Owing to a lack of index fossils, the precise stratigraphic position of the unconformity cannot be established. Based on the age of the uppermost ammonites of the Fillzamin Formation (Late Aalenian–late Early Bajocian) and that of ammonites found at top of the Dansirit Formation (early–middle Late Bajocian), however, a latest Early–earliest Late Bajocian age is beyond doubt.

The unconformities Upper unconformity Lower unconformity The lower unconformity is invariably characterized by a distinct increase in grain size, even though a gradual coarsening trend is developed at the top of the underlying Fillzamin Formation at several localities, most pronounced at Tazareh (Fig. 3E). At other localities, for example at Kuhe-Bashm-e Dehsufian, this coarsening trend is not developed and bioturbated silts of the Fillzamin Formation are overlain by medium-grained sandstones of the Dansirit Formation. At Parvar, the unconformity is less conspicuous and has been placed at the change from prodelta silts with intercalated HCS beds to thick-bedded, predominantly cross-bedded, sandstones of the upper delta front. It is not clear whether the tectonized package of quartz conglomerates exposed south of Shahmirzad (Fig. 4H) comprises the bulk of the Dansirit Formation or only its top. The relationship of the conglomerates to the lower unconformity, therefore, cannot be evaluated. In general, the change in grain size and sediment texture (a change from biogenic to primary sedimentary structures such as trough cross-bedding) coincides with a change from offshore shelf

Throughout the outcrop belt, an upper unconformity is characterized by an abrupt change from the nearshore, shallow-marine –terrestrial sandstones of the Dansirit Formation to the deep marine silty–marly sediments of the Dalichai Formation (e.g. Fig. 6B, D). The unconformity thus clearly corresponds to a rapid deepening pulse of high-magnitude (drowning unconformity). Sedimentary evidence of this event is a 10–20 cm-thick ferruginous lag deposit of relatively uniform character across the outcrop belt. The sediment is generally a highly bioturbated, poorly sorted, bioclastic, fine- to medium-grained sandstone, in some cases a biorudstone, with scattered fossils, among them bivalves (Ctenostreon, Pleuromya, Liostrea), gastropods, brachiopods, serpulids, belemnites and ammonites (Garantiana, Strenoceras). The latter indicate an early– middle Late Bajocian age. The relative concentration of shells suggests a somewhat condensed character of the bed. At some localities (e.g. at Parvar), the unconformity is less spectacular. The uppermost sandstone of the Dansirit Formation is muddy and contains the deep burrowing bivalve Pleuromya in growth position (Fig. 7). In contrast to many other sections, the lowermost marl– limestone beds of the Dalichai

Fig. 6. (Continued) The Mid-Cimmerian unconformities in the Alborz Mountains (A, B, D and F) and east-central Iran (C, E, G and H). (A) Geological sketch of the upper unconformity at Tooy: tilted strata of the Dansirit Formation are overlain, with angular unconformity, by ammonite-bearing transgressive calcarenites and deep-water marls of the Dalichai Formation. (B) Local angular unconformity (arrow) due to block rotation as expression of the upper Mid-Cimmerian unconformity at Tooy. Note the abrupt superposition of deep-water marls on shallow-water sediments (the sandstones of the upper Dansirit Formation are c. 20 m thick). (C) The Mid-Cimmerian unconformity (arrow) at Kalshaneh, east-central Iran: Lower Bajocian Hojedk Formation (left) sharply overlain by dark-coloured limestones of the ?Upper Bajocian–Lower Bathonian Parvadeh Formation (right). The lower and upper unconformities are fused (thickness of exposed strata c. 50 m). (D) Type section of the Fillzamin Formation (300 m thick) near Fillzamin (Ira Valley): soft, dark-green siltstone are capped by sandstones of the Dansirit Formation (arrow at contact, cf. Fig. 6F). The Dalichai Formation is largely covered by scree of the cliff-forming Lar Formation (in the background). (E) Folded Lower Bajocian Hojedk Formation below transgressive limestone of the ?Upper Bajocian– Lower Bathonian Parvadeh Formation near Mazinu, east-central Iran (height of cliff c. 25 m). (F) Lower unconformity at Fillzamin: sharp contact of deltaic sandstones of the Dansirit Formation on deeper marine, organic-rich silts of the Fillzamin Formation (exposed thickness c. 50 m). (G) Basal conglomerate of the ?Upper Bajocian–Lower Bathonian Parvadeh Formation on Permian quartzites (south of Sikhor, southern Shotori Mountains, east-central Iran; 35 cm-long hammer for scale). (H) Parvadeh Formation onlapping Permian quartzites south of Sikhor, southern Shotori Mountains, east-central Iran. The Parvadeh Formation is c. 50 m thick.

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Fig. 7. Upper part of the Dansirit Formation and lower part of the Dalichai Formation north of Parvar. For the key to the symbols see Figure 3.

Formation contain here a considerable amount of silt and sand. The most spectacular unconformity is developed at Tooy, southeastern Alborz Mountains (N378110 7900 , E578040 6100 ; Fig. 6A, B). The uppermost sandstones of the Dansirit Formation with trough crossstratification and questionable Ophiomorpha, interpreted to represent a delta-front environment, are followed by a 20 m-thick succession of dark-grey clay with plant debris (interpreted as interdistributary bay fill), which in turn is overlain by a 1.5 m-thick, fining-upwards, stratified, matrixsupported conglomerate with an erosional base. The pebbles are moderately rounded, up to 10 cm in diameter, and consist of sandstone, ferruginous concretions and yellowish fine-grained limestones reworked from underlying beds. The matrix is a bioclastic calcareous sandstone and contains belemnites, gastropods, crinoid ossicles and oyster shells, clearly documenting the marine character of this high-energy bed. This unit is followed by another 3 m of thick-bedded, sandy bioclastic grainstones and rudstones, with bryozoans, crinoid ossicles and belemnites. The uppermost layer is a nodular biopackstone containing the Late Bajocian ammonite Garantiana and followed by the silty marls of the deeper marine Dalichai Formation. A few 100 m further south, the clay unit and the

basal conglomerate gradually disappear and the deltaic sandstones are directly overlain, at a sharp erosional boundary, by a 1 m-thick crinoidal grainstone. The latter two units form an angle of 5–108. When followed laterally, it becomes clear that this angular unconformity is younger than the clay unit, but older than the conglomerate unit, which clearly is transgressive in nature, even though it is only locally developed. Most probably it accumulated in shallow depressions between highs formed by lithified sediments of the Dansirit Formation. The angular unconformity is evidence of tectonic tilting associated with erosion and onlap of marine sediments. This unconformity should be equivalent to the deepening pulse between the Dansirit Formation and the Dalichai Formation observed at all localities studied by the authors (Fig. 1). Thus, the top unconformity is by no means the only transgressive pulse that can be documented in the Dansirit Formation, but by far the most significant one. For example, the repeated fluctuations between terrestrial and shallow-marine environments seen in the Tazareh section record such subordinate pulses, albeit of a far minor magnitude. Also, the shell-bearing quartz conglomerate at Sharif-Abad, overlying terrestrial red beds, clearly documents a transgressive event. At Golbini (Jajarm area; N378040 1100 , E568280 9900 ) approximately 3 km east of the section shown in Figure 3F, a discontinuous layer of large loaf-shaped cobbles (diameter up to 50 cm) consisting of calcareous sandstone occurs approximately 30 m below the base of the Dalichai Formation. The cobbles have been bored on all sides by bivalves and indicate early diagenetic lithification of the sandstone, followed by exhumation and repeated reworking. Thus, the cobble layer also points to a gap in sedimentation and a phase of erosion most likely in the context of a transgression.

Sequence stratigraphic significance The sandy Dansirit Formation capping the Shemshak Group is bounded by unconformities at the base and top. These two major unconformities are related to pronounced changes in relative sea level and can be interpreted in sequence stratigraphic terms (Fig. 8). As has been shown earlier, additional minor unconformities (mainly flooding surfaces) are evident, to a variable extent, within the formation, but are only briefly discussed in the context of this paper. The basal unconformity clearly represents a major sequence boundary. Sediments of a highstand systems tract (HST) reflecting varying degrees of proximality (offshore mudstones to prodelta–lower delta front siltstone–sandstone intercalations) are

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local/regional tectonic movements or both will be discussed later in this paper.

The Mid-Cimmerian tectonic event in other areas of Iran In their discussion of the Mid-Cimmerian tectonic movements, Seyed-Emami & Alavi-Naini (1990) briefly discussed their manifestations in other parts of Iran, in particular east-central and northeastern Iran. Since their account, extensive additional data concerning the biostratigraphy and sedimentology of these areas have been collected, also with respect to the Mid-Cimmerian tectonic movements, which are summarized in the following.

East-central Iran

Fig. 8. Sequence stratigraphic interpretation of the Mid-Cimmerian tectonic phase (see the text for further explanation). 1, Lower Mid-Cimmerian unconformity; 2, upper Mid-Cimmerian unconformity; SB, sequence boundary; ts, transgressive surface.

sharply overlain by sediments of the lowstand systems tract (LST) (e.g. sandstones of distributary channel, fluvial or high-energy nearshore bar origin). The upper unconformity at the top of the formation is developed as a transgressive surface (TS), followed by a thin transgressive lag, representing the transgressive systems tract (TST) and the expanded succession of the (silty) marls of the Dalichai Formation, which can be interpreted as HST deposits of a carbonate system still under some influence of a siliciclastic regime. Where the Dansirit Formation is expanded as, for example, at Tazareh (Fig. 3E), it is composed of a number of 5–20 m-thick highly asymmetric, coarsening-upwards cycles, which correspond to parasequences (small-scale transgressive– regressive cycles). The transgressive portion is usually represented by a comparatively thin, coarsegrained, sandstone. Such parasequences are also well developed at Golbini. At most other localities, however, these smaller-scale sea-level fluctuations can only be recognized in parts of the sections, if at all. In conclusion, the Dansirit Formation is composed of a number of cycles of different orders of magnitude. Distinct changes in water depth are indicated by unconformities at or near the base of the formation and at its top. Whether the changes in relative sea level are eustatic in nature or reflect

On the Tabas Block, which contains the best record of Jurassic strata within the so-called Central-East Iranian Microcontinent, a distinct unconformity is found at the base of the ?Upper Bajocian – Bathonian Parvadeh Formation. The Parvadeh Formation overlies the Lower Bajocian Hojedk Formation (Fig. 6C), a generally thick siliciclastic unit, which is mainly non-marine in the southern part of the block, where it contains extensive coal deposits in the area between Ravar and Kerman. In the northern part it is either marine or nonmarine. Thicknesses of the Hojedk Formation are known to vary drastically over short distances in the central part of the Tabas Block, from about 60 m at Parvadeh to more than 1000 m a few tens of kilometres further SW. At most localities (e.g. in the Kamar-e-Mehdi area) the overlying Parvadeh Formation starts with a conglomerate or with a pebbly sandstone. The pebbles generally consist of milky quartz, are well rounded and may reach a diameter of several centrimetres. The thickness of the conglomeratic unit varies between 2 m and just a few centimetres. Locally, however (e.g. east of Kuh-e-Echelon), the conglomerate is composed of Palaeozoic and Triassic cobbles (Aghanabati 1977) and is approximately 15 m thick. In the southern Shotori Mountains, east of Kuh-e-Jamal, the Parvadeh Formation unconformably overlies an irregular topography formed by Permian quartzites (Fig. 6G, H). At the base of the formation conglomerate layers comprising quartzite and limestone pebbles in a coarse-sandy matrix occur repeatedly. The succession points to a major hiatus, and to one or several phases of uplift and erosion preceding the transgression of the Parvadeh Formation. In yet other areas (e.g. west of Esfak, north of Kalshaneh), the basal conglomerate consists of bored and encrusted calcareous sandstone

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pebbles; whereas, for example, approximately 65 km SW of Tabas, near the Tabas –Yazd road, there is no sign of an erosional event nor of an influx of coarse siliciclastic material at the base of the Parvadeh Formation. In the southern Tabas Block, between Ravar, Bidou and Kerman, an up to several metre-thick conglomerate layer is commonly developed between the Bajocian Hojedk Formation and the overlying Bidou Formation. The conglomerate consists of sandstone, dolomite and limestone pebbles in a coarse-sandy matrix, in places also of quartz and granitoid pebbles, locally followed by up to 20 m-thick, coarse-grained cross-bedded sandstones (e.g. south of Bidou). The rapid facies changes within the Hojedk Formation, its partly marine and partly terrestrial character, and the strongly varying thickness over short lateral distances suggest that this highly variable facies pattern was produced by syn-depositional block movements. The widespread unconformity and conglomerate layer at the base of the Parvadeh Formation indicate a phase of non-sedimentation, uplift and erosion, followed by renewed deposition. Depending on the relief, the basal sediments of the Parvadeh Formation either represent braided streams (valley fills) or transgressive lags. At a single locality (near Mazinu), the top beds of the Hojedk Formation are gently folded, and overlain by the flat-lying Parvadeh Formation with distinct angular unconformity (Fig. 6E).

Koppeh Dagh (NE Iran) At the southern margin of the Koppeh Dagh, NE of Mashhad, which until the collision with the Iran Plate formed the southern rim of the Turan Plate, the time slice attributed to the Mid-Cimmerian tectonic movements, i.e. around the boundary Lower–Upper Bajocian, is also represented by a distinct angular unconformity (Seyed-Emami et al. 1994; Taheri et al. 2006, 2009). This unconformity separates the locally .2000 m-thick, marine, siliciclastic sediments of the Upper Bajocian–Bathonian Kashafrud Formation from older underlying rocks (e.g. Triassic volcanics, uppermost Triassic– Lower Jurassic granites, Upper Triassic marginal-marine to non-marine siliciclastics). The unconformity cannot be dated with precision, but ammonites such as Garantiana found just above the basal transgressive conglomerate of the Kashafrud Formation are early Late Bajocian in age so that the unconformity must be pre-Late Bajocian. It probably reflects the main collision phase between the Iran and the Turan plates, which took place around the Triassic–Jurassic boundary (Wilmsen et al. 2009b). Even so, the basal conglomerate shows the initiation of rapid

subsidence and the formation of a new large sedimentary basin, the Kashafrud Basin. Initial rapid subsidence took place in the Late Bajocian and is thus coeval with the subsidence pulse observed all across the Alborz at the top of the Shemshak Group, and with that seen at the base of the Parvadeh Formation in east-central Iran. It is logical to assume a common origin for this event.

Geodynamic significance As demonstrated above, the lower Middle Jurassic strata of the Alborz Mountains and east-central Iran (i.e. across the former Iran Plate) share a number of sedimentary and stratigraphic features. Broadly speaking, pronounced subsidence, starting in the Toarcian and culminating in the Aalenian (Fu¨rsich et al. 2005), led to the development of a comparatively deep-marine east –west-trending basin in the Alborz region (Fig. 9). However, subsidence differed markedly normal to strike, as is seen at localities such as Rian and Sharif-Abad, where the thickness of the Toarcian –Aalenian succession is distinctly reduced. This suggests syn-depositional block faulting and thus extensional tectonics as causative factors (e.g. Leeder 1995). Similar, time-equivalent, syn-depositional block faulting can be demonstrated in east-central Iran. There, subsidence differed strongly over short distances, leading to a highly differentiated facies pattern and formation of local troughs (e.g. Zarand Trough on the southern Tabas Block: Huckriede et al. 1962; marked facies and thickness changes in the Hojedk Formation of the northern Tabas Block: Wilmsen et al. 2009a). A reversal of this trend is seen in the late Early Bajocian of the Alborz Mountains where the unconformity at or near the base of the Dansirit Formation (i.e. the lower unconformity) indicates sudden uplift and shallowing (Fig. 9). The Dansirit Formation is generally a shallow-water deposit organized in a number of parasequences and minor unconformities as expressions of changes in relative sea level. While some of these changes most probably record third-order and high-frequency cycles, others may indicate tectonic instability (e.g. the level with encrusted and bored sandstone cobbles near Golbini). The abrupt change in the early Late Bajocian from shallow-water sandstones or calcarenites of the Dansirit Formation to deeper water marl or silty marl of the Dalichai Formation indicates a marked increase in the rate of subsidence, resulting in rapid deepening. The presence of an angular unconformity below this deepening pulse at one locality (Tooy) supports the conclusion that crustal thinning accommodated by block rotation was responsible for the development of this upper

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Fig. 9. Temporal succession of tectonic events in the Lower and Middle Jurassic of the Alborz Mountains and Central Iran. 1, Lower Mid-Cimmerian unconformity; 2, upper Mid-Cimmerian unconformity (see the text for further explanation).

unconformity of the Mid-Cimmerian tectonic event. All these events took place within a short time interval, between the latest Early and the earliest Late Bajocian, i.e. no more than 1 Ma (according to the geological timescale of Gradstein et al. 2004). Thus, what is generally understood as the Mid-Cimmerian tectonic event is now thought to consist of a phase, confined to the Bajocian, during which two major tectonic pulses occurred (Fig. 9). Tectonic instability started during the (Late) Aalenian –Early Bajocian when tectonic subsidence rates were significantly accelerated in the Alborz (Fu¨rsich et al. 2005). Rapid shallowing around the Early–Late Bajocian boundary associated with deformation and local folding must be due to tectonic uplift. The pattern cannot be reconciled with the eustatic sea-level trend at that time (Fu¨rsich et al. 2005), for which a sea-level rise has been postulated (Hardenbol et al. 1998). This compressional ‘mid-Bajocian’ erosion and uplift phase (i.e. the lower unconformity) is followed by a subsidence pulse, indicated by rapid deepening across the upper unconformity. Local tilting associated with this surface is indicative of block rotation created by extensional tectonics with normal faulting. The succession of geodynamic events during the Bajocian of the Alborz share several similarities with the deformation and multiple unconformity development occurring during the Aptian in Mesozoic rift basins of the North Atlantic, which were interpreted as the expression of the break-up

unconformity (Hiscott et al. 1990; Boillot et al. 2002). This unconformity was followed by pronounced post-rift thermal subsidence. Looking at the evolution of the Shemshak Basin during the Early and Middle Jurassic, there is evidence that, after collision between the Iran and Turan plates during the Late Triassic, the basin went through an Early Jurassic molasse stage (Fu¨rsich et al. 2009): coarse clastic material shed from the Cimmerian mountain range in the area of the present-day southern Caspian Sea prograded southwards across the present-day Alborz Mountains. It accumulated in alluvial fans, laterally grading into braided and, further south, into meandering river deposits, the river winding through an extensive coastal plain dotted with lakes and, at times, covered with extensive swamps. By the Toarcian, subsidence increased and marine conditions were established across the Alborz area, whereas previously such conditions were only locally developed and confined to nearshore areas. At that time, the Cimmerian mountain range to the north must have been eroded to a large extent. We speculate that this phase of extension in the Alborz area is the expression of crustal thinning taking place further north, where the South Caspian Basin began to subside in the area of the former Cimmerian mountain range (Brunet et al. 2003). In the Alborz area, this subsidence trend was reversed in the Early Bajocian to become re-established in the Late Bajocian. This phase of short-lived uplift,

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documented by the Dansirit Formation and its unconformities, is best explained as a ‘break-up unconformity’ caused by the onset of sea-floor spreading in the South Caspian Basin to the north (see also Stampfli & Borel 2002; Brunet et al. 2003). If true, this would provide information on the early history of the South Caspian Basin, which so far is not available from the basin fill itself. In particular, it suggests that the rifting phase of that basin started in the late Early Jurassic and that sea-floor spreading was established by the Late Bajocian.

Conclusions Detailed sedimentological and biostratigraphic investigations of the Dansirit Formation, the uppermost formation of the Shemshak Group, in the central and eastern Alborz Mountains reveal the existence of at least two unconformities, which occur at or near the base of the formation and at or near its top. These unconformities document either phases of rapid uplift (lower unconformity) or subsidence (upper unconformity), and run counter to trends of the eustatic sea-level curve. Their tectonic origin supports the existence of the so-called Mid-Cimmerian tectonic movements, which were first described for Iran by Seyed-Emami & Alavi-Naini (1990). Thus, what is generally understood as the Mid-Cimmerian tectonic event is now thought to consist of a phase, confined to the Bajocian, during which at least two tectonic pulses occurred. The onset of tectonic instability culminating in the Mid-Cimmerian event occurred during the Late Aalenian– Early Bajocian when tectonic subsidence rates significantly accelerated across the Iran Plate. Rapid shallowing around the Early– Late Bajocian boundary in the Alborz (deposition of the Dansirit Formation) must have been due to tectonic uplift and erosion of source areas as it contradicts eustatic sea-level trends. The deposition of the Dansirit Formation took less than 1 Ma. This compressional ‘mid-Bajocian’ phase is followed by extensional tectonics, indicated by block rotation, and rapid deepening (deposition of the deep-marine Dalichai Formation) at a time when eustatic sea level is generally thought to have been falling. A roughly time-equivalent unconformity, commonly followed by conglomerates, can be documented at the base of the ?Upper Bajocian– Bathonian Parvadeh Formation of the Tabas Block (east-central Iran). The Parvadeh Formation is underlain by the Lower Bajocian Hojedk Formation, which exhibits pronounced changes in thickness across short lateral distances, as well

as distinct changes in facies. At a single locality (Mazinu, SW of Tabas), folding and thus direct evidence of compressional tectonics is available. After a stratigraphic gap, transgressive sediments of ?Late Bajocian –Bathonian age follow, suggesting a fusion of the lower and upper Mid-Cimmerian unconformity in east-central Iran. Viewed in the framework of the geodynamic evolution of the Iran Plate during the Jurassic (Fu¨rsich et al. 2009), the unconformities and the associated thickness and facies variations are explained as expressions of the onset of sea-floor spreading within the South Caspian Basin situated to the north. This strongly subsiding basin opened close to the Palaeotethys suture, north of the present-day Alborz Mountains, during the Toarcian–Aalenian and went through a change from the rifting- to the spreading-stage during the ‘mid-Bajocian’. The Mid-Cimmerian event is therefore considered to reflect the break-up unconformity of the South Caspian Basin. The research was carried out within the framework of an institutional partnership between the Institut fu¨r Pala¨ontologie of Wu¨rzburg University and the Faculty of Engineering of the University of Tehran, financially supported by the Alexander von Humboldt Foundation, and within the framework of the Middle East Basin Evolution programme (MEBE) based in Paris. We would like to thank the Geological Survey of Iran, Tehran, for logistic support in the field. J. Ineson, Copenhagen and H. Stollhofen (Aachen) provided constructive reviews of the manuscript. Their efforts are highly appreciated.

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S EYED -E MAMI , K., F U¨ RSICH , F. T. & W ILMSEN , M. 2004. Documentation and significance of tectonic events in the northern Tabas Block (east-central Iran) during the Middle and Late Jurassic. Rivista Italiana di Paleontologia e Stratigrafia, 110, 163– 171. S EYED -E MAMI , K., S CHAIRER , G. & B EHROOZI , A. 1994. Einige Ammoniten aus der Kashafrud-Formation (Mittlerer Jura) E Mashhad (NE-Iran). Mitteilungen Bayerischer Staatssammlung Pala¨ontologie und historische Geologie, 34, 145–158. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters, 196, 17–33. S TILLE , H. 1924. Grundfragen der vergleichenden Tektonik. Borntraeger, Berlin. T AHERI , J., F U¨ RSICH , F. T. & W ILMSEN , M. 2006. Stratigraphy and depositional environments of the Upper Bajocian– Bathonian Kashafrud Formation (NE Iran). Volumina Jurassica, 4, 105– 106. T AHERI , J., F U¨ RSICH , F. T. & W ILMSEN , M. 2009. Stratigraphy, depositional environments and geodynamic significance of the Upper Bajocian– Bathonian Kashafrud Formation, NE Iran. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 205–218. W ILMSEN , M., F U¨ RSICH , F. T. & S EYED -E MAMI , K. 2003. Revised lithostratigraphy of the Middle and Upper Jurassic Magu Group of the northern Tabas Block, east-central Iran. Newsletter on Stratigraphy, 39, 143– 156. W ILMSEN , M., F U¨ RSICH , F. T., S EYED -E MAMI , K. & M AJIDIFARD , M. R. 2009a. An overview of the stratigraphy and facies development of the Jurassic System on the Tabas Block, east-central Iran. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 323–343. W ILMSEN , M., F U¨ RSICH , F. T. & T AHERI , J. 2009b. The Shemshak Group (Lower–Middle Jurassic) of the Binalud Mountains, NE Iran: stratigraphy, depositional environments and geodynamic implications. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 175–188. Z IEGLER , P. A. 1975. Outline of the geological history of the North Sea. In: W OODLAND , A. W. (ed.) Proceedings of the 4th Conference an Petroleum Geology, Petroleum and the Continental Shelf of North-west Europe, Volume 1. New York, 131– 149.

Stratigraphy, depositional environments and geodynamic significance of the Upper Bajocian – Bathonian Kashafrud Formation, NE Iran ¨ RSICH2 & MARKUS WILMSEN2,3 JAFAR TAHERI1,2, FRANZ THEODOR FU 1

Geological Survey of Iran, NE Branch, P.O. Box 91735-1166, Mashad, Iran

2

Geozentrum Nordbayern der Universita¨t Erlangen-Nu¨rnberg, Fachgruppe Pala¨oUmwelt, Loewenichstrasse 28, D-91054 Erlangen, Germany

3

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany *(e-mail: [email protected]; [email protected]) Abstract: A strongly subsiding rift basin in NE Iran, the Kashafrud Basin, opened in the Late Bajocian with the accumulation of more than 2000 m of Upper Bajocian–Upper Bathonian siliciclastic sediments. These sediments comprise the Kashafrud Formation, which crops out along a NW– SE stretch of more than 200 km and occupies a width of 50 km, situated between the Koppeh Dagh and the Binalud Mountains. Ten sections of the formation were logged. Sedimentary environments range from non-marine alluvial fans and braided rivers in the lowermost part of the succession to deltas, succeeded by storm-dominated shelf, slope and deep-marine basin. Monotonous mudstones and turbidites prevail in the deep-marine part of the basin. The thickness and facies of the Kashafrud Formation vary strongly between localities, and reflect distance from the rift margins as well as submarine topography, which was shaped by block tectonics. The Kashafrud Basin is interpreted as the eastern extension of the South Caspian Basin, which entered the rifting stage in the late Early Jurassic and the spreading stage in the Late Bajocian.

The Iran Plate, as one of the Cimmerian terranes, had been welded to the Turan Plate of Eurasia by the Late Triassic –Early Jurassic (Stampfli & Borel 2002; Fu¨rsich et al. 2009a). Although this event closed the Palaeotethys Ocean in the area, it was far from becoming stabilized and remained highly tectonically mobile, and several short-lived basins, such as the Sabzevar, Sistan and the Nain– Baft oceans, opened in the Cretaceous (Seyed-Emami et al. 1972; Lindenberg & Jacobshagen 1983; Lindenberg et al. 1983; Tirrul et al. 1983). A strongly subsiding basin started to form in the Late Bajocian in NE Iran (Seyed-Emami & Alavi-Naini 1990; SeyedEmami et al. 1994), termed herein the Kashafrud Basin. It experienced rapid subsidence during the Late Bajocian–Bathonian, with the accumulation of more than 2000 m of siliciclastic sediments, i.e. the Kashafrud Formation. The siliciclastic basin fill was followed, in the study area, by a carbonate platform–basin system during the Late Jurassic (Lasemi 1995) when subsidence rates had considerably slowed down. The Kashafrud Formation crops out across a large area in NE Iran, from NW of Mashhad to Torbat-e-Jam near the Afghan border (Fig. 1). The outcrop region is more than 200 km long and approximately 50 km wide (Aghanabati et al.

1986; Eftekhar-Nezhad & Behroozi 1993). The southwestern margin of the outcrop corresponds to the margin of the Kashafrud Basin, whereas towards the NE the Kashafrud Formation dips below Upper Jurassic–Tertiary strata of the Koppeh Dagh (Fig. 1). As no borehole data are available for the Kashafrud Formation below these younger beds, nothing can be said about the position of the northeastern basin margin. The aim of this paper is to present preliminary stratigraphic, sedimentological and ichnological information on the Kashafrud Formation, and to briefly discuss the origin and geodynamic setting of the basin in which it accumulated.

Geological setting The study area is located between two main sedimentary– structural zones: the Koppeh Dagh in the NE and the Binalud Mountains (the eastern elongation of the Alborz fold belt) in the SW (Fig. 1). The boundary between these two structural zones corresponds to the suture of the Palaeotethys (Alavi 1992), an ocean that had been completely subducted below the Turan Plate as a part of Eurasia in the Late Triassic. This tectonic event is

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 205–218. DOI: 10.1144/SP312.10 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. General geological map of NE Iran with position of all investigated sections and the transect of sections correlated in Figure 3.

Fig. 2. Middle Jurassic lithostratigraphy of the Mashad area.

Fig. 3. Correlation of five sections through the Kashafrud Formation, roughly corresponding to a cross-section through the basin (see Fig. 1). Note that the Radar section (left) is terminated by a fault and therefore does not represent the total thickness of the formation. The key of symbols applies to all figures.

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named the Early Cimmerian orogeny. According to Sto¨cklin (1974), Davoudzadeh et al. (1981), Davoudzadeh & Schmidt (1983), Stampfli et al. (1991), Stampfli & Borel (2002) and Seyed-Emami (2003), collision occurred at the end of the Carnian. Saidi et al. (1997) proposed that collision occurred in the Ladinian, whilst Fu¨rsich et al. (2009a) argue that the collision took place during the Late Triassic, followed by rapid uplift of the Cimmerides at the Triassic –Jurassic boundary. In the Binalud Mountains, Permian low-grade meta-turbidites and ophiolitic rocks, as well as the Triassic (?– lowermost Jurassic) Mashhad Phyllite, are unconformably overlain by Lower –lower Middle Jurassic strata (conglomerates, sandstones, and shales with undeterminable plant fossils and thin coal veins) that have been interpreted as the Shemshak Group (Wilmsen et al. 2009). These deposits (interpreted as Cimmerian molasse) are overlain by marine siliciclastic Kashafrud Formation (Upper Bajocian –upper Bathonian; Fig. 2). The Koppeh Dagh comprises a very thick Middle Jurassic – Tertiary sedimentary succession (Afshar-Harb 1979, 1983, 1994). Outcrops of preMiddle Jurassic rocks, however, are rare, being largely restricted to the tectonic window of the Aghdarband area (e.g. Ruttner 1991; Alavi et al. 1997). Strata of the Shemshak Group (Upper Triassic –lower Middle Jurassic; see Fu¨rsich et al. 2009a) appear to be largely absent from the Koppeh Dagh area, but the Norian –?Rhaetian Miankuhi Formation may represent an equivalent to the lower part of the Shemshak Group. The Kashafrud Formation unconformably rests on folded Triassic or older rocks (e.g. Seyed-Emami et al. 2001) (Fig. 2). The Upper Bajocian–Bathonian Kashafrud Formation is a thick (up to .2000 m) succession of siliciclastic rocks, extending along the southern margin of the NW–SE-trending Koppeh Dagh. Depending on the area, it rests unconformably on Permian ophiolites, Permian metamorphic rocks, uppermost Triassic –Lower Jurassic granites (e.g. north of Torbat-e-Jam), the Middle– lowermost Upper Triassic Sina Formation (a unit of volcanic and volcaniclastic rocks) or the Norian (? –Rhaetian) Miankuhi Formation (a siliciclastic, coal-bearing unit) (Fig. 2). The Kashafrud Formation usually starts with a basal conglomerate, the composition of which reflects the underlying rocks. In order to elucidate the tectono-sedimentary significance of the Kashafrud Formation, 10 sections have been measured, of which five are presented here (Fig. 3). They illustrate the range of facies and environments occurring within the basin, and allow reconstruction of the basin topography along a transect normal to the strike of the basin. The sections were logged in detail and

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Fig. 4. Stratigraphic log of the lower part of the Fraizi section showing the alluvial fan and braided river facies association (for locations and the key to the symbols see Figs 1 and 3).

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sampled for sedimentological, biostratigraphical and palaeontological data (macrofossils and trace fossils). The biostratigraphy is based on ammonites and, in non-marine sequences, on macroplants and palynomorphs.

Stratigraphy Based on an unpublished report of the National Iranian Oil Company, Afshar-Harb (1969) introduced the Kashafrud Formation into the literature for an intercalation of sandstones and shales. Subsequently the term was used by geologists dealing with the stratigraphy and geology of the area (e.g. Aghanabati 1998). However, no formal type section was ever proposed. Madani (1977) interpreted a section near the village of Baghbaghou as representing flysch deposits. Afshar-Harb (1983), Aghanabati et al. (1986) and Eftekhar-Nezhad & Behroozi (1993) mapped the Kashafrud Formation on the Sarakhs, Mashad and Torbat-e-Jam quadrangle sheets, respectively, and Taheri & Ghaemi (1999) used it on the Mashad 1:100 000 sheet. Seyed-Emami et al. (1994, 1996) and Hoseiniun (1995) studied ammonites of the formation. Taheri et al. (2006) and Wilmsen et al. (2006) gave a brief account of the stratigraphy, facies and origin of the Kashafrud Formation. Recently, Poursoltani et al. (2007) characterized the rocks of the formation as representing deep-water fans. The Kashafrud Formation overlies different older rocks with angular unconformity and is followed gradually by the Callovian Chaman Bid Formation – a unit consisting of silty marl and argillaceous limestones – and, in the eastern part of the basin, by limestones (Mozduran area) or sandy limestones (Ghal-e-Sangi area) of the Upper Jurassic Mozduran Formation (Fig. 2). The age of the base of the formation is Late Bajocian, documented by ammonites such as Orthogarantiana, Garantiana and Parkinsonia (Seyed-Emami et al. 1994, 1996; Hoseiniun 1995). The Late Bathonian age of the upper part of the Kashafrud Formation is indicated by the ammonites Prohecticoceras haugi (Provici-Hatzeg), Cadomites claramontanus (Kopik), Procerites (Siemiradzkia) sp. and Homoeoplanulites sp., found in the Jizabad and gas

pipeline sections. This age has been confirmed by the occurrence of Macrocephalites near the base of the overlying Chaman Bid Formation at Radar and Fraizi localities. The thickness of the Kashafrud Formation varies from about 300 m (Radar section) to possibly more than 2900 m (Maiamay section). Towards the west, the Kashafrud Formation interfingers with the more marly–calcareous Dalichai Formation (a unit that is widespread in the Alborz Mountains of northern Iran; see Seyed-Emami et al. 2001; Fu¨rsich et al. 2009b). In the Fraizi area (Fig. 1), a transitional facies is developed consisting of marly silt with intercalations of sandstones, which has been mapped as Dalichai Formation on the geological maps (e.g. Aghanabati et al. 1986), but in our opinion corresponds more closely to the lithology of the Kashafrud Formation than the Dalichai Formation. At this locality, the Kashafrud Formation is overlain by the platform limestones of the Lar Formation (Fig. 2).

Facies associations Going from the basin margin in the SW to the basin centre in the NE (Fig. 3), six major facies associations have been recognized within the Kashafrud Basin: alluvial fan; braided river; delta; shelf; slope to toe-of-slope; and basin plain. Field observations and lateral facies relationships suggest steep SW–NE gradients, and show a main sedimentary input from the southwestern basin margin. Furthermore, a shallowing trend towards the SE can be observed and the greatest thicknesses are recorded from the Maiamay-gas pipeline area, which may represent the depocentre of the basin (Fig. 3).

Alluvial fan and braided river Rocks of this facies association occur in the lower part of the Fraizi section and include boulder beds, conglomerates and largely coarse-grained arkosic sandstones (Fig. 4). The composition of pebbles, cobbles and boulder-sized components varies according to source rocks (Fig. 5H). Lithologies range from argillaceous silt, siltstone, immature,

Fig. 5. (Continued) Field aspects of the Kashafrud Formation: non-marine–shelf facies. (A) Coarsening- and thickening-upwards parasequences of shelf origin near the northwestern margin of the basin; Danesh section. (B) Small-scale parasequence; Danesh section (the length of the pen is 14.5 cm). (C) Bioturbated, fine- to medium-grained, thin-bedded sandstone and silt of outer shelf–upper slope origin, Kol-e-Malekabad area. (D) Thickening- and coarsening-upwards fan delta succession, Fraizi section. (E) Convolute bedding and flame structures in marine fan delta sandstones; Fraizi section. (F) and (G) Pebbly sandstones and conglomerates of braided river origin (note the lenticular shape of the conglomerate bed in G); east of Fraizi. (H) Polymict conglomerate of alluvial fan origin (brown-coloured) unconformably overlying grey-coloured Mashad Phyllite (arrow); east of Fraizi. (I) Plant fossils from fine-grained interbeds within braided river facies; east of Fraizi (the length of the pen is 14.5 cm).

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Fig. 5. Continued.

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angular to subangular, medium- to coarse-grained, mica-bearing sandstone, to pebbly sandstone, microconglomerate and conglomerate. The conglomerates are poorly sorted and have an arkosic sandstone matrix. West of Fraizi, 40 m of light-grey pebbly kaolinitic sandstones and conglomerates (components: milk quartz, metamorphic rocks and greywacke) overlie a 0–50 m thick reddish boulder bed and conglomerates of the basal Shemshak Group (Wilmsen et al. 2009). A nearly 300 m-thick boulder bed and conglomerate overlying the Mashhad Phyllite SE of Fraizi more probably belong to the basal Kashafrud Formation (see later). In the northwestern part of the basin near Fraizi metre- to decametre-thick lenticular packages of highly immature, large-scale trough cross-bedded, gravelly –pebbly, coarse-grained arkoses and conglomerates form a nearly 1000 m-thick succession that overlies the basal conglomerates (Figs 4 and 5F, G). Occasionally finer interbeds with plant fossils (Fig. 5I) are the most important characters of this facies. The boulder beds, conglomerates and pebbly sandstones are interpreted as debris flows and mud flows of proximal alluvial fans (e.g. McGowen & Groat 1971; Miall 1996). Their rapid lateral and vertical changes in facies and thickness suggest that deposition was largely fault-controlled. The thick stack of poorly sorted arkoses and conglomerates near Fraizi represent the distal equivalents of the proximal alluvial fans, which grade into braided river systems basinward. Although the relationship of the thick conglomerate unit and the overlying packages of coarsegrained arkosic sandstone unit to the Kashafrud Formation is not documented by index fossils (plant fossils occurring in the lower part of the arkosic sandstones were not diagnostic enough to yield a precise age), we tentatively place this unit at the base of the Kashafrud Formation because we could not find any unconformity or any other distinct change between the non-marine and the marine (Upper Bajocian –Bathonian) part of the succession in the Fraizi area. It may well be that a large fluvial system entered the basin in the area and that the supply of sediment kept pace with the high subsidence rate.

Delta A fan delta environment (e.g. McPherson et al. 1988) is indicated by several metre- to decameterthick, coarsening-upwards, medium- to coarsegrained, trough cross-bedded sandstones with abundant plant remains and cross-bedded, coarsegrained–pebbly sandstones with intercalations of micro-conglomerates (Figs 5D and 6). Abundant syn-sedimentary deformation structures are the

Fig. 6. Stratigraphic log of the middle part of the Fraizi section showing the delta facies association. Note the predominance of coarsening- and thickening-upwards cycles (for locations and the key to the symbols see Figs 1 and 3).

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result of rapid sediment loading (Fig. 5E). These types of sediments were observed at three localities: SE of the small village Khij (N368300 2500 , E598040 5000 ), where large convolute bedding and flame structures in thick coarse-grained sandstones (Fig. 5D, E) are the first indications of marine sedimentation in the area and show rapid

Fig. 7. Stratigraphic log of the lower part of the Kol-e-Malekabad section showing the shelf facies association overlying a basal trangressive conglomerate. Note the stacked parasequences as the predominating motif (for locations and the key to the symbols see Figs 1 and 3).

Fig. 8. Stratigraphic log of the lower part of the Maiamay section showing the slope and toe-of-slope facies associations (for locations and the key to the symbols see Figs 1 and 3).

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sedimentation in an upper delta-front environment. Similar sedimentary conditions existed at the Fraizi and Kuh-e-Radar sections. At Kuh-e-Radar wide, lenticular structures consist of 0.5–4 m-thick, partly convoluted, fine- to coarse-grained sandstone and conglomerate beds containing wood fragments. The architecture of the units resembles crosssections of fan lobes, which are composed of stacked channel fills. Angular pebbles, cobbles, and large boulders of Mashhad granitoids and pegmatite rocks in the sediments are a characteristic feature and indicate a very short transport distance that can be related to strong tectonic uplift of closely adjacent areas. Large and thick-shelled bivalves of the Middle –Upper Jurassic genus Fimbria are abundant 45– 50 m above the base of the Radar section.

Shelf The shelf facies association comprises cross-bedded sandstones and bioturbated argillaceous siltstones to fine-grained sandstones, which often are arranged in parasequences (Fig. 7). Abundant wood fragments and hummocky cross-stratified sandstones (tempestites) are characteristic sedimentary features of this lithofacies. Limestones dated as Upper Bajocian occur only at Tappeh Nader. They are about 50 m thick, consist of bedded fossiliferous wackestones, and contain abundant ammonites and bivalves. They are replaced up-section by sandy and marly rocks. These limestones are inferred to have accumulated on the crest of a fault block, which formed a positive feature (swell) on the shelf. Corals, bivalves and other benthic faunas of shelf origin occur in low densities in the lower part of the Kashafrud Formation at Ghal-e-Sangi and Torbat-e-Jam, and indicate comparatively shallow-water conditions. Another main feature of shelf deposits are parasequences that are well developed in the Danesh and Kol-e-Malekabad areas (Fig. 5A, B). Bioturbation is common at these localities (Fig. 5C), and the sediment usually varies between argillaceous silt and fine-grained

sandstone. Storm beds were observed in the middle part of the Ghal-e-Sangi section (at 720 m).

Slope to toe-of-slope Lenticular, thick-bedded, sharp-based, medium- to coarse-grained arkosic– lithic sandstones are interpreted as amalgamated and cannibalizing feeder channel fills (Figs 8 and 9E, I, J). Chaotic conglomerates with shallow-water components represent debris flows. They are associated with dark-green silty–fine-sandy shales, sandstones and silty shales. Slump units indicate a considerable relief (Fig. 9F, G). This facies association is considered to represent slope environments. Slope sediments are common, and more obvious in the lower and middle parts of the sections at Kol-e-Malekabad and Maiamay. The lithology is intercalated between marl and fine- to medium-grained sandstones in the Maiamay area and dark-grey, finegrained sandstones at Kol-e-Malekabad. Toe-of-slope facies is represented by fine- to medium-grained, graded sandstones intercalated between silty and fine-sandy shales, often arranged in thinning-/fining-upwards sequences, which are interpreted as submarine turbidite fans. Characteristic features include convolute bedding, as well as tool and flute marks. Graded bedding is best seen in thick sandstones in the lower part of member 2 of the Kashafrud Formation at Maiamay. These sandstones, which have sharp, erosional, gravelly bases, also exhibit convolute bedding and flame structures. They are probably evidence of decreasing slope angles at the toe-of-slope, on which proximal turbidites came to rest (Fig. 9K, L).

Basin plain Although turbidites are common in most parts of the outcrop area (e.g. at Maiamay, gas pipeline, Kol-e-Malekabad, Ghal-e-Sangi and Senjedak districts), field observations show that classic turbidites, with or without CCC turbidites (characterized by clasts, convolution and climbing ripples; cf.

Fig. 9. (Continued) Field aspects of the Kashafrud Formation: deep-marine facies. (A) Highly bioturbated, dark-green –olive siltstone, just above the transgressive basal conglomerate, evidence of fault-controlled rapid subsidence; Senjedak section (arrowed person in the centre gives the scale). (B) Classic thin-bedded basin plain turbidites in the Maiamay area. (C) Monotonous, dark-green, bioturbated siltstone with rare intercalations of thin, fine-grained sandstones, gas pipeline area. (D) Limestone intercalations (arrowed) in the uppermost part of the Kashafrud Formation, gas pipeline section, containing Late Bathonian ammonites. (E) and (H) Massive turbiditic sandstone, Maiamay area. Note the sharp base of those units (arrowed in H). (F) and (G) Convolution (arrow in F) and slumping structures in slope facies, Maiamay and Ghal-e-Sangi areas, respectively (the length of the pen is 14.5 cm; the hammer is 30 cm long). (I) and (J) Deep-water submarine fan sandstones showing cannibalizing feeder channels (arrowed), Maiamay area (thickness shown in I corresponds to c. 7 m; hammer in J is 30 cm long). (K) Convolution (arrowed) and graded bedding in proximal turbidites, Maiamay area (the length of the pen is 14.5 cm). (L) Graded, coarse-grained sandstone, Maiamay area (the diameter of the lens cap is 6.5 cm).

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Fig. 9. Continued.

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Walker 1992), are more obvious and common in the area south of Mazdavand (near the intersection of the Mashhad– Sarakhs–Salehabad roads), and in the middle part of the Maiamay section. The facies includes mostly thin, graded, fine- to medium-grained sandstones (distal turbidites; Figs 9B and 10), dark, silty– fine-sandy shales, in part bioturbated (Nereites ichnofacies) with rare ammonites or the bivalve Bositra, grey–greenishgrey, marly– silty shales and black shales with ferruginous concretions (Fig. 9C). In the gas pipeline area, the facies is more finegrained and monotonous, and represents deeper environments (Fig. 9C). The relationship between this area and Maiamay is not clear because of a major fault. In the Maiamay area, turbidites are more common than in the gas pipeline area, in which turbidites have mainly a distal character. In general, trace fossils are more common at Maiamay than in the gas pipeline area, which may be related to a higher degree of bioturbation at the latter locality and the difficulty in recognizing and differentiating trace fossils in fine-grained sediments. Rip-up clasts are common, and clasts, pebbles and benthic fossils of shallow-water origin occasionally occur in these facies. They demonstrate transport of shallow-water components to deep-marine environments.

Discussion Analysis of the sedimentary fill of the Kashafrud Basin and reconstruction of the evolution of the basin is crucial for understanding the geodynamic changes that affected the region after the Iran Plate had become attached to Eurasia in the context of the Late Triassic Early Cimmerian orogeny. Moreover, the dark shales that dominate in the more central parts of the basin have a great potential as hydrocarbon source rocks (see also Poursoltani et al. 2007). Although during the last few decades stratigraphic and sedimentological aspects of the Kashafrud Formation were studied (e.g. Madani 1977; Seyed-Emami et al. 1994; Hoseiniun 1995; Poursoltani et al. 2007; several unpublished master theses of Iranian universities), this is the first time that the whole basin is considered. The current data support a rift basin model (Fig. 11) rather than a Neotethys Ocean relationship of the Kashadfrud Basin, as proposed by Poursoltani et al. (2007). Our assumption is based on the following criteria. Fig. 10. Stratigraphic log of the upper part of the Maiamay section showing the basin plain facies association (for locations and the key to the symbols see Figs 1 and 3).

† Pronounced lateral changes in facies and thickness exist within short distances and suggest strong syn-sedimentary tectonic activity (Fig. 3). Examples are the shelf carbonates at

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Fig. 11. Diagrammatic cross-section through the southwestern margin of the Kashafrud Basin, showing the lateral changes in facies of the Kashafrud Formation from the fault-controlled basin margin to the deep-marine central basin. Localities refer to rough positions of the sections of Figure 1.

Tappeh Nader, which are replaced laterally by deep-marine, bioturbated siltstones. Basin subsidence appears to have been controlled by normal (i.e. extensional) faults, which strike NW– SE when mapped by facies distribution (e.g. at the radar section). † There is a general rift-related deepening trend also seen in other upper Middle Jurassic rock units in northern Iran (Dalichai and Chaman Bid formations; see Fu¨rsich et al. 2009b). Moreover, a number of diabasic dykes and sills exist, particularly at the southern margin of the basin, especially in the Torbat-e-Jam and Senjedak areas, which do need, however, to be dated and studied in more detail. † The rift model, which implies a direct relationship with the South Caspian Basin further west, corresponds more closely with existing palaeogeographical maps (Fig. 12) (Thierry 2000). The Kashafrud Basin was, as was the case in the South Caspian Basin, separated from the Neotethys Ocean in the south by the Iran Plate and its eastern continuation. † The NE–SW strike of the basin axis, as proposed by Poursoltani et al. (2007, fig. 12), is perpendicular to the NW –SE trend suggested herein based on facies mapping, evidence from numerous sections in the entire basin, the orientation of basin-bounding normal faults and the

general strike of large structures in NE Iran (Figs 1 and 12). The NW-directed palaeoflow reconstructed by Poursoltani et al. (2007) indicates, in our opinion, an axial (i.e. basin-axis parallel) transport of material rather than an input from the basin margin. Axial transport directions are common or even predominate in many rift basins and turbidite depositional systems (e.g. Leeder & Gawthorpe 1987; Leeder 1995; Stow et al. 1996). There is, in fact, also a shallowing trend in the Kashafrud Basin from NW to SE, but this is related to the ‘zip-like’ southeastward opening of the basin, which was deep in the NW, fairly shallow in the Torbat-e-Jam area and graded eastsoutheastward into non-marine strata in northern Afghanistan. In this region, the terrestrial, coal-bearing Bathonian Ispusta Formation, which is up to 800 m thick, overlies with angular unconformity various older stratigraphic units (Wohlfart & Wittekindt 1980). † A high rate of subsidence is indicated by a package of siliciclastic sediments more than 2000 m thick, and by a significant deepening of the basin from above sea level to a basin plain at least a few hundred metres deep, if not more, during a relatively short timespan of less than 5 Ma, according to the geological timescale of Gradstein et al. (2004). This has also been

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Fig. 12. Palaeogeography and plate tectonic reconstruction of the Iran and Turan plates during the late Middle Jurassic (based on Thierry 2000). The geographical and geodynamic position of the Kashafrud Basin and its relationship to the South Caspian Basin is shown in the inset (for simplicity reasons, no palaeogeographic details are shown for the Iran Plate).

documented recently by Brunet et al. (2007) and Shahidi (2008). The field observation that the transgressive basal conglomerate of the Kashafrud Formation occurs in close stratigraphic distance to deep-marine, highly bioturbated, olive-green siltstones and fine-grained sandstones in the Senjedak area (Fig. 9A) may point to rapid subsidence of a fault-bounded basin (Stow et al. 1996). The tectonic subsidence rates of 80 –100 m Ma21 (corrected for compaction, bathymetry, eustatic sea-level and isostasy; unpubl. data of A. Shahidi and M.-F. Brunet, Paris) are characteristic of young rift basins (e.g. Leeder 1995; Ziegler & Cloetingh 2004). Coeval rock units such as the Bashkalateh Formation (Afshar-Harb 1979, 1994) of the easternmost Alborz and the Dalichai Formation of the central and eastern Alborz provide links between the Kashafrud Basin and the South Caspian Basin (see Brunet et al. 2003) that started to open earlier, i.e. during Toarcian–Aalenian times (Fu¨rsich et al. 2005). It becomes increasingly

clear that, from the late Early Jurassic onwards, a rift basin opened in northern Iran in a ‘zip-like’ manner from west to east (Fig. 12). The reason for the basin development may be related to Neotethys back-arc rifting. This rift proceeded more-or-less along the Palaeotethys suture and reached NE Iran in the Late Bajocian. The Bajocian Mid-Cimmerian tectonic event may indicate the break-up unconformity of the South Caspian Basin (Brunet et al. 2003; Fu¨rsich et al. 2009b), which is inferred to have developed into a back-arc ocean during the late Middle– Late Jurassic (Stampfli & Borel 2002). It remains unknown whether the Kashafrud Basin also reached the spreading stage.

Conclusions In the Late Bajocian, a fault-controlled, extensionrelated basin, the so-called Kashafrud Basin, originated in northeastern Iran, close to the suture line of the former Palaeotethys Ocean. The basin axis trended roughly NW –SE. Biostratigraphic and sedimentological data indicate rapid subsidence

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during the Late Bajocian –Bathonian, with the main source of sediment input from the southwestern basin margin (Binalud Mountains). The basin was filled nearly exclusively with siliciclastic sediments ranging from conglomerates to clays, accommodated in the Kashafrud Formation. Palaeoenvironments range from alluvial fans and braided rivers to deltas, the storm-influenced shelf, slope and deep-marine basinal areas; the latter two characterized by turbidite sedimentation. The observed features of the Kashafrud Formation are best integrated in a rift-basin model. The northeastern basin margin is inferred to lie below the Koppeh Dagh. The Kashafrud Basin is the southeastern prolongation of the South Caspian Basin and developed in response to Late Bajocian– Bathonian crustal extension following the Mid-Cimmerian tectonic movements. In addition to the more detailed study of the basin fill, the analysis of related rock units, such as the Mashhad Phyllite as well as the Mesozoic rocks of the Aghdarband area – in particular the Miankui Formation – is a prerequisite to understanding the evolution of the basin in more detail. We wish to thank the GSL reviewers, S. Egan (Keele) and K. Seyed-Emami (Tehran), for their constructive reviews. K. Seyed-Emami is further thanked for identification of the ammonites from the Kashafrud Formation. We thank also the Geological Survey of Iran (GSI) at Tehran and Mashad for logistic support and expert guidance in the field. Financial support was provided by the German Research Society (Deutsche Forschungsgemeinschft, grant FU 131/32-2), which is gratefully acknowledged. We thank M.-F. Brunet and A. Shahidi (Paris) for discussions of the geodynamic background of the Kashafrud Formation.

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Subsidence and uplift mechanisms within the South Caspian Basin: insights from the onshore and offshore Azerbaijan region STUART S. EGAN1*, JON MOSAR2, MARIE-FRANC¸OISE BRUNET3,4 & TALAT KANGARLI5 1

School of Physical and Geographical Sciences, Keele University, Keele, Staffordshire ST5 5BG, UK

2

De´partement de Ge´osciences de l’Universite´, Ge´ologie et Pale´ontologie, Pe´rolles-Ch du Muse´e 6, CH-1700 Fribourg, Switzerland 3

UPMC Universite´ Paris 06, UMR Tectonique, case 129, 4, place Jussieu, F-75005 Paris, France 4

CNRS, UMR Tectonique, F-75005 Paris, France

5

Geology Institute of Azerbaijan, National Academy of Sciences, H. Javid av. 29A, Baku AZ1143, Azerbaijan. *Corresponding author (e-mail: [email protected]) Abstract: A combination of fieldwork, basin analysis and modelling techniques has been used to try and understand the role, as well as the timing, of the subsidence –uplift mechanisms that have affected the Azerbaijan region of the South Caspian Basin (SCB) from Mesozoic to Recent. Key outcrops have been studied in the eastern Greater Caucasus, and the region has been divided into several major tectonic zones that are diagnostic of different former sedimentary realms representing a complete traverse from a passive margin setting to slope and distal basin environments. Subsequent deformation has caused folds and thrusts that generally trend from NW– SE to WNW– ESE. Offshore data has been analysed to provide insights into the regional structural and stratigraphic evolution of the SCB to the east of Azerbaijan. Several structural trends and subsidence patterns have been identified within the study area. In addition, burial history modelling suggests that there are at least three main components of subsidence, including a relatively short-lived basin-wide event at 6 Ma that is characterized by a rapid increase in the rate of subsidence. Numerical modelling that includes structural, thermal, isostatic and surface processes has been applied to the SCB. Models that reconcile the observed amount of fault-controlled deformation with the magnitude of overall thinning of the crust generate a comparable amount of subsidence to that observed in the basin. In addition, model results support the tectonic scenario that SCB crust has a density that is compatible with an oceanic composition and is being under-thrust beneath the central Caspian region.

Some of the deepest sedimentary basins, and largest sources of hydrocarbons, in the world occur within intra-continental settings, but they are poorly understood in terms of the mechanisms that have controlled their subsidence history. The South Caspian Basin (SCB) is an example of one of these deep basins (Fig. 1), with a depth to basement of over 20 km in places (Shikalibeily & Grigoriants 1980; Brunet et al. 2003). Although it is widely accepted that the SCB was initiated by Mesozoic back-arc extension related to the subduction of the Tethys Plate (e.g. Zonenshain & Le Pichon 1986), more than half of the 20 km subsidence presently observed occurred within the tectonic framework

of the evolution of the Alpine –Himalayan orogenic belt. The main focus of this investigation has been to attempt to decipher the role played by various subsidence and uplift mechanisms that have affected the SCB region from Mesozoic to Recent. A combination of fieldwork, basin analysis and numerical modelling techniques has been used to compare and contrast structural styles, stratigraphic architecture and subsidence–uplift history of the SCB region. The investigation has concentrated on the Azerbaijan region of the basin (Fig. 1), mainly due to data availability. Key outcrops have been studied in the eastern Greater Caucasus, including the central part, its termination with the Caspian

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 219–240. DOI: 10.1144/SP312.11 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Map of SCB and surrounding regions showing general location of onshore (yellow box) and offshore (red box) study areas (background image courtesy of NASA World Wind). Points A and B define the ends of the cross-section presented in Figure 2.

Sea, and the northern and southern fold-and-thrust fronts. The overall geodynamic setting of the eastern Greater Caucasus corresponds to a region of continental collision that has inverted former passive margin. The tectonic features are compatible with those of a doubly verging mountain belt with two external fold and thrust belts. One of the main objectives of the fieldwork has been to gather structural data from outcrops of various stratigraphic ages located at the SCB margins, as well as within the neighbouring Greater Caucasus orogenic belt. This structural investigation has been complemented by a sedimentological analysis to provide insights into the dominant depositional environments during the basin’s evolution. The region is divided into several major tectonic zones that are diagnostic of different former sedimentary realms representing a complete traverse from a passive margin-type setting, with sediments from lagoonal to reef

environments, to slope and distal basin environments. The onshore component of this study has provided a precise overview of the structure, stratigraphy and tectonic evolution of the eastern Greater Caucasus along a NNE–SSW transect, which can be used as a proxy for the evolution of the neighbouring SCB. A number of regional-scale cross-sections and horizon depth maps, constructed from public domain and commercial exploration data, have been used to carry out a basin analysis exercise on the offshore Azerbaijan region of the SCB. The crosssections provide information on the regional structural and stratigraphic evolution of the SCB, while horizon depth maps have enabled the identification of regional and temporal structural trends. In particular, the map data show the close spatial relationship between the Apsheron Sill (Fig. 1) and the deepest parts of the basin. In addition, burial history modelling has been used to determine trends in subsidence

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within different parts of the study area. This modelling reveals several distinct subsidence events that have affected the basin, including a relatively shortlived basin-wide event at 6 Ma that is characterized by a rapid increase in the rate of subsidence. Structural and geodynamic modelling has been applied to the SCB that enables the forward modelling of extensional basin evolution due to rifting followed by subsequent deformational events. A kinematic modelling approach has been used, whereby a section of lithosphere is defined in terms of its thickness, density and thermal structure. Once the lithosphere is defined it is possible to numerically model its deformation via structural, thermal and isostatic processes (e.g. Egan & Meredith 2007). This modelling is particularly useful for providing insights into the effects of deep-lithosphere processes that are not well constrained by geological and geophysical data. It has been used to test a variety of tectonic scenarios for SCB evolution, including depthdependent stretching and the possibility that the SCB crust has a density that is compatible with an oceanic-type composition.

Geological background and evolution of the South Caspian Basin The SCB occupies the deep southern part of the Caspian Sea, and extends onshore in West Turkmenia in the east and the Lower Kura in the west (see Fig. 1 for location). It is one of the deepest basins in the world, with an estimated depth to basement of 20–25 km (e.g. Shikalibeily & Grigoriants 1980; Zonenshain & Le Pichon 1986; Brunet et al. 2003). The northern boundary of the SCB lies along the Apsheron Sill, a structural high that links to the Apsheron peninsula and the eastern Greater Caucasus. The Greater Caucasus orogenic belt was formed by compressional deformation resulting from the closure of a Jurassic–Eocene back-arc basin (e.g. Zonenshain & Le Pichon 1986; Ershov et al. 1999). The tectonic history of the region is dominated by folding and thrusting generated by the convergence between the Arabian and Eurasian plates. Recent seismic activity suggests that the SCB is behaving as a rigid block surrounded by weaker seismically active regions (Jackson et al. 2002; Allen et al. 2003). In addition, regional convergence of Arabia and Eurasia has resulted in a series of arcuate folds in the west of the basin that curve from the NW– SE trend of the eastern Greater Caucasus, Apsheron Sill and Alborz regions, to an approximately north – south trend to the west and east of the basin. These folds are mainly anticlinal buckle folds, which mostly detach within the mud-dominated deposits of the Maikop Suite (Allen et al. 2003). A linear, NW–SE-trending zone of folds in the north of the

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basin forms the Apsheron Sill, a topographically elevated region that is thought to be related to the under-thrusting of the SCB beneath the central Caspian basin. There remains controversy regarding the timing of the opening of the SCB and whether there are two different sub-basins or only one. The majority of authors advocate that the SCB is a back-arc basin that formed in the Mesozoic (e.g. Zonenshain & Le Pichon 1986). The opening of the Greater Caucasus back-arc basin started at the end of the Triassic and continued throughout the Jurassic (Ershov et al. 2003). A similar late Triassic –lower Jurassic timing is proposed for rifting in the proto-SCB (Brunet et al. 2003, 2007). Golonka (2000, 2004) proposes that the opening of the SCB took place in several stages from the Jurassic to the Eocene, and even as late as the Miocene in the southern part of the basin. Other authors, using data from the Talesh area (see Fig. 1 for location), propose that the main phase of opening of the SCB occurred during the Eocene (Kazmin 1991) or that there was opening of a southern trough near the Alborz at this time (Mamedov 2004). It has also been suggested that the SCB could be a remnant of the Palaeotethys Ocean (Berberian 1983) and that its relatively great depth is a result of subduction of continental crust during the mid –late Cenozoic (Knapp et al. 2000, 2004). The age of the oldest deposits in the deep central basin are not known. Part of the sedimentary sequence has been determined from onshore drilling in Azerbaijan and Turkmenistan where shallow-water sediments ranging from Late Jurassic to Pleistocene age have been found (Shikalibeily et al. 1988; Mamedov 1992; Ali-Zade et al. 1999; Knapp et al. 2004). The majority of the sedimentary infill of the SCB is Oligocene and younger. The Maikop Suite forms the main hydrocarbon source rock in the SCB. The suite is dominated by mud-based sediments deposited from the Oligocene to the Miocene that have been subject to overpressuring as a result of rapid deposition of overlying sediments. They are commonly found at the surface at selected onshore localities where they form mud volcanoes (e.g. Guliyev & Feizullayev 1997; Devlin et al. 1999). The Productive Series forms the main reservoir unit in the basin and is 5 –7 km thick in the centre of the basin. The series consists of late Miocene –early Pliocene fluvial and deltaic deposits (Reynolds et al. 1998; Allen et al. 2003; Green et al. 2009). The thickest component of the basin infill is represented by sequences of Pliocene– Quaternary in age (Fig. 2), where more than 10 km of sediment were deposited in less than 5 Ma (Brunet et al. 2003); this represents an average depositional rate of approximately 2 km Ma21 for the last 5 Ma (Knapp et al. 2004). Zonenshain & Le Pichon

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Fig. 2. Regional crustal-scale cross-section through the Kura and South Caspian basins (adapted from Baranova et al. 1991; Mamedov 1992). See Figure 1 for location.

(1986) and Nadirov et al. (1997) have suggested that these rapid rates of deposition were due to enhanced tectonic subsidence as a result of compressive forces acting in the last 6 Ma with the sediment supplied by erosion of the uplifted areas at the basin margins. Green et al. (2009), however, explain the increased rate of deposition as being due to a rapid fall in base level at the end of the Miocene related to the Messinian Salinity Crisis. The basin was infilled rapidly by clastic deposits from palaeo-river systems when sea level rose again during the Pliocene. The depth to the Moho varies from about 60 km, beneath the onshore South Caspian region, to approximately 30 km beneath the deepest part of the SCB (Fig. 2). Based on seismic velocity data, it has been suggested that the ‘granitic’ crustal layer is absent in the central part of the basin (Mangino & Priestley 1998). This could be because the crust is of oceanic origin, although it is thicker than normal oceanic crust. Alternatively, it could be continental crust of which the upper section has been removed by erosion or faulting, or it could be thinned and intruded continental crust. It has also been suggested (Zonenshain & Le Pichon 1986) that the crust beneath the SCB is of continental origin but has been subjected to high pressures and high temperatures, such that the granite rocks of the continental basement were metamorphosed to become eclogite. From earthquake studies the SCB crust appears to be a relatively rigid and aseismic block within the framework of the Alpine–Himalayan orogenic belt (Priestley et al. 1994), which is further evidence for a rheological difference between the crust of the SCB and that of the surrounding region (Mangino & Priestley 1998). Currently, the general consensus, mainly derived from geophysical data and gravimetric modelling (e.g. Shikalibeily &

Grigoriants 1980; Berberian 1983; Baranova et al. 1991; Mangino & Priestley 1998), is that the basement of the marine part of the basin comprises a high-velocity, thin complex crust. Controversy remains on its exact nature; thin high-density continental crust or thickened oceanic crust. Earthquakes of depths up to 80 km have been recorded beneath the Apsheron Sill region. Some authors consider this seismicity to be caused by bending and faulting of the SCB lithosphere as it subducts beneath the Apsheron Sill (Priestley et al. 1994). Mangino & Priestley (1998) and Knapp et al. (2004) also suggest that shallow-thrust events along the Apsheron Sill and folding of the sedimentary basin fill indicate that the Apsheron Sill is overriding the subducting crust of the SCB basement. It has also been proposed that the SCB is under-thrusting along its southwestern margin beneath the Talesh and Alborz mountains based on evidence from shallow-angle thrusting (Berberian 1983; Allen et al. 2003). However, Guest et al. (2007) propose a crustal buckling mechanism to explain subsidence in the southern part of the SCB and its relationship to the magnitude of uplift in the neighbouring Alborz mountains in northern Iran.

Tectonic and stratigraphic evolution of the eastern Greater Caucasus Key outcrops have been studied in the eastern Greater Caucasus in Azerbaijan, extending from the frontal thrust range bordering the Kura Basin to the south, to the trailing part of the mountain belt in the north and bordering the eastern Terek Basin, to the eastern limits of the Greater Caucasus along the shores of the Caspian Sea. The main objectives of this field-based investigation were to document the different levels of the stratigraphic

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Fig. 3. Generalized tectonic cross-section through the central part of the eastern Greater Caucasus in Azerbaijan (adapted from Kangarli 1982). The major tectonic zones, and the northern and southern foreland basins, are shown. The tectonic zones also correspond to the major depositional realms, which from north to south are lagoon (restricted), reef/platform, proximal–distal slope (basinal), respectively.

column from Jurassic to present in order to understand the palaeotectonic evolution of the region. The overall geodynamic setting of the eastern Greater Caucasus corresponds to a region of continental collision that has inverted the former passive margin of the Scythian Plate (Philip et al. 1989). To the south, volcanic material is derived from an active subduction-related volcanic arc in the Lesser Caucasus. The tectonic features of the eastern Greater Caucasus are compatible with those of a doubly verging mountain belt with two external fold-and-thrust belts (Fig. 3). The pro-wedge (front) is located to the south and overrides the Kura Basin, whereas the retro-wedge (back) is located to the north and overrides the eastern Terek Basin. The tectonic features and structural relationships are well exposed in the incised river valleys on the northern and southern slopes of the Greater Caucasus, but also in the steep slopes and cliffs of the higher central mountain range where the average altitude is above 2000 m and culminates at 4466 m with mount Bazar Du¨zu¨ at the border between Azerbaijan and Dagestan. Overall, the tectonic structure is one of a typical fold-and-thrust belt, with fault-related folding, thrusts, tectonic imbrications, tectonic klippen, relay structures, faulting and late transverse structures. The mountain-building tectonic events overprint structures that are related to syn-sedimentary palaeotectonics, both extensional and compressional (inversion). Folds and thrusts generally trend NW – SE to WNW –ESE in the eastern Greater Caucasus, cleavage is axial-surface parallel and oriented WNW– ESE with steep –upright dips. Fold-axial

surfaces are steep and dip to the north in the southern and central regions of the mountain range, showing a southwards vergence linked to a transport direction to the south in the frontal part of the mountain belt. Steep structures in the north part are associated with thrusting towards the north in the northern frontal part of the mountain range. It has been possible to identify a succession of tectonic and sedimentary events within the eastern Greater Caucasus region based on data derived from the study of key outcrops. The oldest lithologies encountered are sedimentary sequences of Aalenian in age, located within the central range of the mountain belt. These sequences consist of interbedded shales and sandstones that formed within an extensional basin environment at the southern edge of the Scythian platform. North of the main fold-and-thrust belt, it is possible to observe sanddominated sequences at the base of the Lower Oxfordian resting with angular unconformity on folded Aalenian –Callovian sequences. An additional unconformity can be observed locally between the folded Callovian and the equally folded underlying Bathonian age units. A Jurassic–Cretaceous unconformity is represented by tilted Kimmeridgian beds overlain with erosive unconformity by Berriasian conglomerates that consist of clasts of Kimmeridgian limestone and Aalenian lithologies. The strong angular unconformity between conglomerates and the underlying Kimmeridgian– Tithonian sequences is interpreted to indicate fault-block tilting at the Jurassic–Cretaceous transition. The Valanginian– Barremian sequences are characterized by slump features and olistolith deposits.

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The olistoliths reach sizes of more than 200 m long and 35 m in thickness, and contain fragments of Jurassic limestone. The size and the repeated series of olistoliths suggest ongoing tectonic processes acting upon the carbonate platform forming the edge of the Scythian Plate. Folding and thrusting in the late Albian (pre-Turonian) led to the development of an erosive unconformity followed by sedimentation of neo-autochthonous sequences. The beginning of this neo-autochthonous sequence is marked by a basal conglomeratic sequence of Turonian age in the east and Campanian age to the west. The sequence of events summarized above represent the pre-mountain belt (i.e. pre-Greater Caucasus) evolution of the region that occurred within a back-arc basin setting in response to subduction of the Tethys Plate to the South. The deformational regime was one of extensional tectonics and passive subsidence interrupted by compressional– inversion events. Collision then occurred between Arabia and Eurasia, leading to the closure of this basin system followed by a sequence of mountainbuilding events. An Eocene erosional unconformity marks the beginning of the formation of the

fold-and-thrust belt, with the main uplift of the mountain range starting in Mid-Miocene. This deformation continues to the present with ongoing folding and thrusting, and progressive propagation of the thrust front at the southern leading edge of the orogenic belt. In addition, a general brittle overprinting can be observed throughout the area. This brittle tectonic deformation is represented by transverse faults trending NNE–SSW to NE–SW that cut all existing structures. Similar trends can be found in the fold patterns offshore that have been associated with faults within the basement. The chart provided in Figure 4 provides a visual summary of the tectonic and stratigraphic events that have affected the eastern Greater Caucasus region.

The geological evolution of the South Caspian Basin: offshore Azerbaijan A basin analysis exercise has been carried out on the Azerbaijan region of the SCB (see Fig. 1 for location) using interpretations of offshore seismic data provided by BP. The raw data are based on

Fig. 4. Summary of the tectonic and stratigraphic evolution of the eastern Greater Caucasus produced from field-based investigation. Information on major regional events has been obtained from Zonenshain & Le Pichon (1986), Philip et al. (1989) and Brunet et al. (2003).

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14 regional seismic lines that image down to 20 s TWT (two-way time). Depth maps have been produced for a number of stratigraphic horizons from ‘basement’ to Upper Pliocene, three of which are presented in Figure 5. These maps have been generated from depth-converted interpretations of the original seismic data (see, for example, the subsection on ‘Data constraints and model parameters’ later), which have been sampled and interpolated using a kriging algorithm to generate a regular grid across the study area, enabling contour maps to be generated. Analysis of the depth map for the Top Surakhany horizon (Upper Pliocene, dated at approximately 3.4 Ma; Devlin et al. 1999) reveals several structural trends and subsidence patterns (Fig 5a), including a relatively shallow-depth region with a NW–SE trend in the NE, where the horizon reaches depths of about 400 m. This trend corresponds with the orientation of the Apsheron Sill, which has been formed by compressional deformation that affects Mesozoic –Recent sequences. It is also a trend that is compatible with onshore axial surfaces of the folds that are steep and dip to the north in the southern part of the Greater Caucasus mountain range. The deepest part of the horizon occurs adjacent to the Apsheron Sill where there is a depression reaching depths of about 4750 m. This depression follows the southern margin of the Apsheron Sill and can be explained by flexural deepening due to under-thrusting or, possibly, subduction (Priestley et al. 1994; Granath et al. 2007) of South Caspian basement beneath the middle Caspian region. The west of the map in Figure 5a shows a series of linear, relatively shallow features at depths of about 800 –1200 m. They have a NW–SE trend in the western part of the study area, changing to a more east-west trend in the southern part of the map where they also become more closely spaced. Interpretations of seismic data show that these features represent a set of closely spaced folds that mostly deform Maikop and younger sequences. Similar trends can be found in the fold patterns onshore. There is a NE– SW trend of linear highs that are most evident in the eastern part of the study area. These structures are more widely spaced than those with a NW– SE trend. In additiona, they do not reach the same shallow depths as the NW– SE-trending features. They have a similar trend to a set of transverse faults that cut all of the existing structures onshore. The horizon deepens towards the SE corner of the map where it has depths of 3000–3500 m. This deepening occurs towards the Pre-Alborz trough, which is located in the southeastern corner of the SCB as it approaches the Alborz orogenic belt. Figure 5b shows a similar depth map, but for the Top Mesozoic horizon. This map shows less area of

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coverage and detail compared to the younger horizon described above due to reduced quality and more limited coverage of the seismic data with increasing depth. The horizon in the main part of the study area has an average depth of about 11 500 m. The Apsheron Sill is, again, an area of relative uplift in the north at a depth of 7000 m. Both NE –SW- and NW –SE-trending features are present with typical depths of about 9000 m. However, they form broader features than in the younger horizons, which could be related to decoupling between pre- and post-Mesozoic strata or due to the superposition of folds with different wavelengths and amplitudes. In addition, there are linear regions of relative uplift with a north– south trend in the western part of the study area. The depth to the Top Jurassic horizon is shown in Figure 5c. There is a broad, relatively shallow area in the south of the study area at depths of between 13 000 and 15 000 m. Although the coverage of this map is limited, there is some evidence of uplift of the Apsheron Sill region, illustrated by a decrease in depth in the NE corner of the map to about 12 000 m. The deepest part of the area, at about 20 000 m, is again immediately to the south of the Apsheron Sill. However, this depression does not appear to continue as far to the west as it does in the younger horizons. Burial history modelling is an important analytical technique that provides insights into the history of the deepening of a basin and therefore burial of the infilling sequences (Sclater & Christie 1980). In addition, the burial history of a basin is indicative of the tectonic and isostatic processes operating at any particular time. Basin Mod 1-D software, produced by Platte River Associates, was used to construct burial history plots for selected locations within the SCB study area. It generates time v. depth plots that show the burial history of infill sequences within a basin through time, including the effects of tectonic subsidence, and burial and compaction due to the deposition of younger sequences. In order to carry out the modelling the software requires the ages and depths to each horizon as main input parameters. The lithological composition of each horizon was estimated based on evidence presented in the relevant literature and using Basin Mod 1-D’s default lithology data. The locations selected for burial history analysis are shown in Figure 5a and were chosen for the following reasons: † Location 1 was selected because it is located within the region of uplift forming the Apsheron Sill; † Location 2 represents a relatively deep area of the basin margin in the NW;

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Fig. 5. Contour maps of depth to selected stratigraphic horizons within the SCB showing major structural trends (contours are at 250 m intervals). (a) Top Surakhany horizon (lower upper Pliocene). Labels ‘Loc 1’– ‘Loc 6’ define the locations at which burial history analysis has been carried out (see Fig. 6). (b) Top Mesozoic horizon. (c) Top Jurassic horizon.

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Fig. 5. Continued.

† Location 3 was selected to represent the deepest part of basin to the south of the Apsheron Sill; † Location 4 lies on a NNW –SSE-trending high in the SW corner of the study area; † Location 5 lies on a NE– SW-trending high in the centre of the basin; † Location 6 was selected to represent the SE part of the study area where the basin deepens towards the Pre-Alborz trough. The depths to major stratigraphic horizons were determined from the regional depth maps. It should be noted that the depth data were only available to the base of the Oligocene at most locations owing to the limited data coverage for the older sequences. Location 5, however, situated in the centre of the study area, has data for nine stratigraphic horizons from middle Cretaceous to upper Pliocene. The results from the burial history modelling (Fig. 6) suggest that there are at least three phases of subsidence. During the Cretaceous and most of the Palaeogene the basin exhibits gradual subsidence that decreases almost exponentially. This pattern of subsidence is compatible with a post-rift thermal subsidence phase, along with infill of the basin, following initial rifting of the SCB in the early –mid-Jurassic. This thermal

Fig. 6. Burial history modelling for locations 1 –6 (see Fig. 5a for locations). At least three phases of subsidence can be identified.

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subsidence phase was disrupted at the beginning of the Oligocene when there was an increase in sedimentation rate. This change is likely to be generated by flexurally induced subsidence in response to sediment input from the mountain building occurring onshore. At the same time, compressional deformation affected the Apsheron Sill region. At about 6 Ma the burial data shows a basin-wide event that is characterized by a rapid increase in the rate of sedimentation. The timing of this subsidence phase corresponds to the main uplift event identified onshore from field-based investigation and, therefore, is likely to result from flexural loading. The basin was infilled rapidly by an abundant supply of sedimentary material generated by erosion of the surrounding orogenic belts and a deep incision created by a short and local sea-level drop, equivalent to the Mediterranean Messinian event (Green et al. 2009).

Subsidence mechanisms within the South Caspian Basin: insights from structural and geodynamic modelling Structural and geodynamic modelling has been applied to the SCB that enables the forward modelling of extensional basin evolution due to rifting followed by subsequent inversion events. The modelling is useful for testing the viability of a variety of tectonic scenarios that may explain the subsidence history of basins that are poorly constrained by subsurface data.

Modelling approach A kinematic modelling approach has been used, whereby a section of lithosphere is defined in terms of its thickness, density and thermal structure. Once the lithosphere is defined its deformation is numerically modelled by the following geological and geodynamic processes. Mechanical thinning or thickening of the crust due to movement along crustal faults. Thrust fault movement generates increased surface topography as well as constituting a downwards-acting load upon the lithosphere, which responds by flexing. In contrast, extensional faulting causes negative loading of the lithosphere and generates accommodation space in the form of graben formation. The fault configuration within the model is defined in terms of the number of faults required, and their relative spacing and orientation (i.e. synthetic or antithetic). All of the faults are assumed to terminate or detach at mid –lower crustal levels in order to maintain compatibility with evidence from deep seismic reflection data (e.g. Snyder & Hobbs

1999), seismological investigations (e.g. Jackson & McKenzie 1983) and rheological modelling of the lithosphere (e.g. Kusznir & Park 1987). Movement along each of the faults is defined in terms of a heave value, and the amount of crustal thinning or thickening is calculated by using the Chevron or vertical shear construction (Verrall 1982; White et al. 1986). A listric fault geometry has been used within the modelling carried out in this study in order to simplify the model calculations. A more sophisticated flexural cantilever model based upon planar faults (e.g. Kusznir & Egan 1989) would make little difference to model results at the scale of deformation being investigated here. Mechanical thinning or thickening of both crust and mantle lithosphere by a regionally distributed pure shear mechanism. Deformation by a pure shear (i.e. stretching or squashing) mechanism is assumed to deform the more ductile sections of the lithosphere within the lower crust and mantle lithosphere (Kusznir et al. 1987). Both the lateral position and the width of the pure shear deformation can be defined. Disturbances caused to the lithosphere temperature field. Lithosphere extension and compression causes an overall heating and cooling of the temperature field, respectively. For example, reverse movement along crustal faults causes the emplacement of relatively hot hanging-wall material onto that of a cool footwall, creating a cooling of the geotherm at shallow depth. Compressional deformation by pure shear thickens the lithosphere and also reduces the geothermal gradient. Within the model the lithosphere temperature field is represented as a twodimensional grid. Each node on the grid is assigned a pre-deformational temperature. The temperature of each grid node is then modified according to the amount of deformation experienced. Re-equilibration of the temperature field after deformation. After a major tectonic event the lithosphere temperature field experiences a gradual thermal recovery. For example, in the context of extension the temperature field will cool back to an equilibrated state. The model calculates both lateral and vertical heat flux by the mechanism of conduction. The physical properties of the lithosphere are such that the thermal recovery process takes of the order of 100 ma. Flexural isostatic compensation of tectonic loads. Crustal thinning and thickening, and thermal perturbations caused by tectonic activity, impose loads upon the lithosphere that have to pass through an isostatic filter to give a compensated amount of subsidence or uplift observed at the surface. The flexural isostatic response of the lithosphere to loading is

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calculated within the model in order to generate a realistic surface or basement profile and underlying crustal structure. The methodology used to model flexure assumes the lithosphere behaves as a continuous elastic plate, which is in equilibrium under the action of all applied loads (Turcotte & Schubert 2002). The flexing properties of the lithosphere are defined by its flexural rigidity (Walcott 1970), which is, in turn, set in the model by the parameter effective elastic thickness (Te). Surface processes. The model contains simple algorithms to simulate the infill of accommodation space and the erosion of uplift. Isostatic adjustments are calculated in response to these processes such that sedimentary infill enhances subsidence within a basin, whereas erosion has the combined effect of reducing the size of the uplifts and unloading the lithosphere, which responds by regional isostatic uplift or rebound (Egan & Urquhart 1993). A comprehensive description of all theoretical aspects of the modelling approach are presented in Kusznir & Egan (1989), Egan & Urquhart (1993) and Egan & Meredith (2007), and will not be repeated here. However, Figure 7 illustrates some of the parameters included in the model calculations and examples of results. A typical starting condition for the modelling is shown in Figure 7a, which illustrates a regional cross-section of undeformed lithosphere. The crustal component of this lithosphere is assumed to be 35 km thick with a density of 2800 kg m23, while the mantle lithosphere is assumed to be 90 km thick with a density of 3300 kg m23. The modelled lithosphere is thermally conditioned with an equilibrated geotherm, which has a surface temperature of 0 8C and a temperature at the lithosphere–asthenosphere boundary of 1333 8C. All of these parameters can be varied according to the tectonic scenario to be modelled. The effects of extending this lithosphere are shown in Figure 7b. The model shows how the crust has accommodated extensional deformation by movement along a sequence of faults, while the mantle lithosphere has been deformed by regionally distributed pure shear. The boundary between upper crustal deformation and that in the lower crust and mantle lithosphere is determined by a detachment depth, which is equivalent to the concept of necking depth presented by other authors (e.g. Kooi et al. 1992). The model shows a basement profile consisting of a sequence of closely spaced half-grabens with relative uplift of the footwall due to flexural isostatic processes (Weissel & Karner 1989; Egan 1992). Extension has also caused heating of the lithosphere temperature field, which subsequently has cooled to generate gradual subsidence. The effects of this cooling can be seen in the model by a post-rift stratigraphic

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sequence that blankets the underlying fault blocks and syn-rift stratigraphy. Lithosphere shortening is represented by the model in Figure 7c, which shows a ‘piggy-back’ style of thrusting adjacent to a foreland basin that has been generated mainly by flexure in response to crustal thickening caused by compressional tectonics. In addition, the modelling approach is sufficiently versatile to include multiple extensional and compressional events.

Data constraints and model parameters Data constraint for the modelling is based on several regional cross-sections that have been generated from the interpretation of the original seismic data carried out by BP geoscientists to show structural and stratigraphic components. The cross-sections in Figure 8a and b show structural and stratigraphic components (in TWT) based upon the interpretation of two regional seismic lines acquired from SW to NE across the southern part of the study area. The section in Figure 8a begins about 125 km from the southern Azerbaijan coastline and continues NE for approximately 75 km into the central SCB; the precise location of this section is confidential. There is approximately 9 km of overlap between this cross-section and the SW part of the section in Figure 8b. The cross-sections show evidence of extensional faulting of basement; however, the magnitude of this extension appears to be quite small. In addition, there is evidence of fault-related compressional deformation in the NE part of the section shown in Figure 8b that has uplifted Mesozoic and younger sequences to form the southern part of the Apsheron Sill. The thrust structures in this region appear to detach on or within basement, which is deepened significantly beneath the Apsheron Sill. The cross-section presented in Figure 8c shows structural and stratigraphic components for part of the SCB, Apsheron Sill and Central Caspian Basin (CCB). It has been produced from interpretations of a number of seismic lines from these areas. The SW part of the section is dominated by shortwavelength buckle folds that die out about 60 km along the section at the start of the broad depression representing the central part of the SCB. This deep part of the basin is abruptly terminated by the Apsheron Sill, which forms a broad region of faultcontrolled uplift that has deformed Cretaceous – Upper Pliocene sequences. The thrusts in this area show vergence dominantly to the south. There is distinct thinning of Upper Pliocene (post-Top Surakhany) and Pleistocene sequences across the structure suggesting that this part of the Apsheron Sill has been experiencing uplift until very recently. The CCB is a much shallower structure compared to the SCB. For example, the Top Cretaceous

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Fig 7. An integrated kinematic model of lithosphere deformation. (a) Cross-section of undeformed lithosphere showing initial values of parameters used in the model calculations. This model can be applied to simulate (b) lithosphere extension and associated basin formation, (c) lithosphere shortening and associated thrust-belt and foreland basin formation, or a combination of the two.

boundary reaches depths of 8.5 s TWT in the south compared to about 4 s TWT in the north. The Mesozoic sedimentary sequence is relatively thin, but shows localized thickening of Triassic sequences that is likely to have been caused by normal faulting and small-scale rifting events (Green et al. 2009). The CCB part of the section also shows a broad arch structure that is onlapped by post-Mesozoic

sequences. The axis of this arch is located about 150 km to the north of the Apsheron Sill and can be interpreted as a flexural bulge structure generated in response to crustal thickening and loading created by the formation of the Apsheron Sill. The SW –NE-oriented cross-sections shown in Figure 8a and b have been combined and depth-converted to give coverage over a distance

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Fig. 8. (a) and (b) Cross-sections showing structural and stratigraphic components (in TWT) based on the interpretation of two overlapping regional seismic lines acquired from SW to NE across the study area. (c) Composite section across the SCB, Apsheron Sill and CCB.

of about 200 km from SW to NE of the offshore study area (Fig. 9a). One of the main input parameters required to carry out the modelling is quantification of the magnitude of deformation due to faulting. Each major fault is identified from the data in terms of its position. In addition, deformation is quantified as a positive or negative heave value, representing extension or compression, respectively. Figure 9b and c show major extensional and compressional faults, and their magnitude of heave interpreted from the cross-sections presented in Figure 8.

Model scenario 1 – uniform lithosphere deformation Figure 10 shows a model representation of the SCB and Apsheron Sill region, assuming evolution by regional uniform lithosphere deformation by a coupled faulting –pure shear process. The overall dimensions of the model are based on the regional sections presented in Figure 9, whereby major extensional faults observed in the cross-sections, along with the movement along them, have been reproduced in the model. Deformation by faulting is balanced to regionally distributed pure shear deformation (i.e. stretching) within the lower crust and mantle lithosphere to represent uniform lithosphere extension. Other model

parameters, as well as assumptions made, include rifting is instantaneous; the original (i.e. pre-deformational) crustal thickness is assumed to be 35 km; all faults have a surface dip of 458 and are assumed to detach at a depth of 20 km, below which the lower crust and mantle lithosphere deform by pure shear (Kusznir et al. 1987); accommodation is filled to sea level with sediment (average density of 2500 kg m23); the density of the crust and mantle are assumed to be 2800 and 3300 kg m23, respectively; lithosphere thickness is assumed to be 125 km, with the basal boundary defined by the 1333 8C isotherm (McKenzie 1978). In addition, the flexural isostatic response of the lithosphere is constrained by an effective elastic thickness (Te) of 5 km during rifting, and 10 km during the post-rift and compressional phases. These values are based on a sensitivity test of Te in terms of it generating the most realistic combination of basin geometry and subsidence, and are also compatible with a number of extensional basin settings (e.g. Kusznir et al. 1991; Van Wees & Cloetingh 1996). Figure 10a represents the initial rift phase of the basin that is assumed to have occurred in the middle Jurassic. This timing is compatible with the Mesozoic origins of the proto-SCB as proposed by Golonka (2000, 2004) and Brunet et al. (2003). The model profile in Figure 10b has simulated

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Fig. 9. (a) The SW– NE cross-sections from Figure 8a and b have been combined and depth-converted. (b) Quantification of fault deformation from composite section in (a). (c) Quantification of fault deformation from part of the SCB–Apsheron Sill –CCB composite section (Fig. 8c).

the thermal subsidence phase, combined with the sedimentary infill of accommodation, for a period of 150 Ma. The final phase of basin evolution is represented by the model in Figure 10c, which includes the effects of compressional deformation within the Apsheron Sill region. This compressional deformation is assumed to have occurred at 10 Ma, which approximates to the timing of the main uplift event identified onshore (Fig. 4). It has caused localized uplift of syn- and post-rift stratigraphic sequences in parts of the basin, whereas regions adjacent to the uplift have been deepened due to the crustal loading caused by the deformation. Model results show clearly that it is not possible to reproduce subsidence in the SCB with extensional and compressional deformation constrained by the magnitude of faulting apparent from the data. Maximum observed subsidence within the basin is over 20 km, whereas the model shows a maximum depth of approximately 5 km.

Model scenario 2 – enhancing basin subsidence due to crustal attenuation A potential flaw with the modelling approach used to generate the models in Figure 10 is that the deformation, and therefore subsidence, predicted is constrained by the magnitude of fault-controlled deformation determined from the interpretation of seismic data. It is now acknowledged that crustal extension by fault displacement is much lower than overall crustal thinning in a variety of sedimentary basins (e.g. Moretti & Pinet 1987; Driscoll & Karner 1998; Kusznir et al. 2004). It is very likely, therefore, that the data interpretation presented in Figure 9 does not represent all of the deformation that has occurred within the SCB and, in particular, in the Apsheron Sill region where later compressional deformation will have caused the inversion of extensional faults. In other words, the extensional faulting observed today is

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Fig. 10. Model representation of the SCB –Apsheron Sill region, assuming evolution is due to uniform lithosphere deformation. A sequence of model profiles are presented showing: (a) the initial rift phase; (b) the post-rift, thermal subsidence phase; (c) and the effects of compressional deformation.

only a fraction of that which occurred during the initial rift phase of the basin. In order to counter this potential problem, a modelling approach has been used in which the magnitude of deformation has been calculated using crustal thickness rather than basement faulting. An estimation of Moho depth based on published material (Mangino & Priestley 1998) indicates that the crust beneath the SCB has been thinned to about 10 km along the

line of section being considered. This equates to an extension factor (i.e. Beta value; McKenzie 1978) of 3.5, assuming an original crustal thickness of 35 km. The model profiles presented in Figure 11 reconciles the magnitude of observed faulting with the overall thinning of the crust by assuming depthdependent stretching. The scenario modelled is one where the detachment or necking depth is allowed to progress towards the surface rather than being

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Fig. 11. Model profiles that reconcile the magnitude of observed faulting with the overall thinning of the crust by assuming depth-dependent stretching. (a) Extensional evolution of the basin showing syn- and post-rift sequences. Stratigraphic time lines within the post-rift sequence are at 20 Ma intervals. (b) Compressional deformation to form the Apsheron Sill region. (c) As profile (b), but showing basin-fill sequences and the underlying lithosphere structure.

fixed at mid-crustal levels. The detachment zone may migrate throughout the evolution of a rift or during multiple rifts because the thickness of the brittle and ductile zones would change constantly during deformation in response to both the changing thickness of the crust and due to temperature perturbations (Kusznir & Park 1987). The faulting configuration in the model is based on the regional

sections presented in Figure 9, but extension has been increased in the lower crust and mantle lithosphere to reproduce a realistic attenuation of the crust. In response, the overall magnitude of subsidence has been increased to a maximum of 18 km. In addition, the pattern of subsidence exhibited in this model is more similar to that exhibited by the real data, with a large thickness of post-rift

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Fig. 12. Model profiles assuming that SCB crust is oceanic. (a) Profile showing basin, crust and upper part of the mantle lithosphere. (b) As profile (a) but showing basin-fill sequences and the top of the crust. The oceanic crust generates additional isostatic loading that results in a maximum basin depth of over 20 km, which is very similar to that observed in the basin.

deposition relative to a thinner syn-rift phase. The model also includes compressional deformation (Fig. 11b, c) to generate the Apsheron Sill region, which, like the previous model, has caused localized uplift adjacent to regional, flexurally induced subsidence.

Model scenario 3 – enhanced subsidence due to oceanic crust The model results presented above have assumed that the crust beneath the SCB is continental in affinity, with a density of 2800 kg m23. It has been suggested, however, that the SCB crust is oceanic (e.g. Brunet et al. 2003; Knapp et al. 2004), or a remnant of lower continental crust (Mangino & Priestley 1998), which means that it may have a density of 3000 kg m23 or greater. The model presented in Figure 12 is based on the same parameters and assumption as that in Figure 11; however, the thinned crust beneath the SCB has an increased density of 3000 kg m23 that is compatible with oceanic-type crust. This density increase generates additional isostatic loading and, as a result,

maximum basin depth is increased to over 20 km in the main part of the SCB, which is very similar to that observed in the basin.

Model scenario 4 – subduction of South Caspian crust Several authors have presented convincing evidence that the crust beneath the SCB is being subducted beneath the Central Caspian region (e.g. Priestley et al. 1994; Knapp et al. 2004; Granath et al. 2007). This tectonic scenario has been modelled using the kinematic modelling approach described above. Figure 13a shows a model representation of the extensional evolution of the Caspian Basin. The model has been generated by extending the lithosphere by stretching over a width of 400 km. Although it is acknowledged that the original width of the basin is open to debate, the current width of the basin in a NNE– SSW direction, between the Apsheron Sill and Iranian coastline, is about 325 km. The results from gravity modelling carried out by Granath et al. (2007) suggests that a section of SCB crust

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Fig. 13. Models showing the possible effects of the subduction of the SCB crust. (a) Model representation of the extensional evolution of the basin, which is assumed to be underlain by oceanic-type crust. The maximum basin depth is about 18 km. (b) 80 km of subduction/under-thrusting of the SCB crust causes uplift of the basement, and synand post-rift sequences to form the Apsheron Sill. The adjacent basin has a depth of over 20 km due to isostatic loading generated by subduction and the infill of accommodation space.

at least 80 km in length has been under-thrust beneath the Central Caspian region. The original basin, therefore, must have been at least 400 km wide along the line of section being considered. The magnitude of extension included in the model has been gradually increased to thin the crust to 10 km (cf. original crustal thickness of 35 km) beneath the central part of the basin. It has also been assumed that oceanic-type crust, with a density of 3000 kg m23, exists beneath the basin compared to continental crust at the margins with a density of 2800 kg m23. The basin has been allowed to thermally subside for a simulated time of 150 Ma and any accommodation space has been filled to sea level with sediment with an average density of 2500 kg m23. The maximum depth of the basin is about 18 km. The model presented in Figure 13b has assumed that an 80 km-section of South Caspian crust has been subducted northwards beneath the central Caspian region, as proposed by Granath et al. (2007). The magnitude and timing of subduction of the northern SCB is open to debate. In the absence of definite information, subduction is assumed to have started at 10 Ma to coincide with the main period of mountain building as evidenced from onshore data. The model results show relative uplift of basement, syn-rift and post-rift sequences

over the Apsheron Sill, while the adjacent SCB is deepened to over 23 km. The overall subsidence exhibited by this model shows the closest match to the observed subsidence indicated by the data interpretation presented in Figure 9, suggesting that the subduction of oceanic-type crust is a viable scenario to explain the evolution of the SCB.

Discussion and summary Integration of the onshore and offshore components of the study reveals a more detailed picture of the tectonic and sedimentary events that have controlled the evolution of the Azerbaijan region of the SCB. Analysis of outcrop data within the eastern Greater Caucasus region reveals a succession of events that began with the deposition of sedimentary sequences within the basin environment at the southern edge of the Scythian platform during the middle Jurassic. These sequences are currently located within the central range of the Greater Caucasus mountain belt. The first evidence of compressional deformation is a Mid-Cimmerian unconformity that is represented by folded sequences of Aalenian – Callovian in age overlain unconformably by flatlying Callovian–Oxfordian carbonates and evaporites. There is another unconformity marking the

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end of the Jurassic period where tilted Kimmeridgian beds are overlain with erosive unconformity by Berriasian conglomerates. The lower Cretaceous is characterized by a sequence of olistolith deposits, suggesting ongoing tectonic processes at the edge of the carbonate platform forming the edge of the Scythian Plate. The mid-Cretaceous is marked by another folding event. All of these events occurred within the tectonic framework of back-arc extension related to the subduction of the Tethys Plate to the south beneath the Lesser Caucasus arc system. During this time period the SCB region was part of a much larger basin system, generally referred to as the Greater Caucasus basin or trough (Zonenshain & Le Pichon 1986; Brunet et al. 2003), which was at least 300 km wide and 3000 km in length. The region was dominated by extensional tectonics interrupted by inversion events caused by regional plate movements and interactions. The sequence of Mesozoic tectonic events that have been identified within the eastern Greater Caucasus are difficult to correlate directly with the evolution of the SCB. It is clear, however, from interpretations of the seismic data (Fig. 8) that the basement has been affected by extensional faulting which relates to the extensional regime that dominated the region at the time. It is not possible, however, to identify individual inversion events within the SCB. The sedimentary sequences of Mesozoic age within the basin typically occur at depths of over 10 km at which the resolution of the offshore data is not sufficient to identify relatively minor tectonic events. Although compressional deformation has affected Mesozoic and younger sequences in the basin, it is impossible to separate out episodes of deformation related to mountain building in the Tertiary from older events such as Mid-Cimmerian deformation. These events that are clearly identifiable onshore can be used as a proxy for the evolution of the adjacent basin. During the Neogene period the basin system started to close in response to the collision between Arabia and Eurasia, leading to the main period of mountain building onshore. Folding and thrusting continues to the present day as evidenced by earthquake activity (e.g. Priestley et al. 1994) and progressive propagation of the thrust front at the southern leading edge of the orogeny (Philip et al. 1989). The mountain building has imposed a dominantly NW–SE structural trend on the region that is picked out by the axial surfaces of major folds, as well as cleavage development. The same structural trend continues into the SCB, most noticeably represented by the Apsheron Sill that effectively represents an offshore continuation of the eastern Greater Caucasus. The deeper part of the SCB also exhibits several other structural trends that can be related to onshore observations. In particular, the

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western part of the basin shows folding with a dominantly north–south orientation, which is compatible with fold structures of Pliocene age that can be observed onshore to the south and north of the main mountain belt. The offshore data also reveal a NE– SW-oriented sequence of uplift structures that are dominant in the eastern part of the study area. This structural trend can be identified throughout the whole mountain range as a system of large brittle faults that dissect the tectonic structures. These faults are readily seen on satellite images and are highlighted by the drainage pattern of the major rivers. Structural and geodynamic modelling shows that it is not possible to explain subsidence in the SCB by uniform lithosphere extension when the magnitude of extension is constrained by the observed faulting of basement. This result is not surprising given the magnitude of mountain-building activity that has affected the Caspian region during the Tertiary period, which is likely to have inverted many of the extensional faults generated during the Mesozoic rift phase of the basin. In addition, Walsh et al. (1991) have shown that typically 10–40% of total fault-related extension may not be observed on regional-scale seismic profiles, a factor that is exacerbated in the SCB due to the relatively low data coverage of the basement structure. Model results indicate that it is necessary to combine faultcontrolled deformation with realistic thinning of the crust in order to reproduce a significant magnitude of subsidence. These models are based on a depthdependent stretching mechanism, which does present space problems caused by exaggerated stretching of the lower crust and mantle lithosphere (e.g. Rowley & Sahagian 1986). A number of possible mechanisms may provide a solution to this apparent complication. For example, lower crustal flow has also been advocated as a mechanism for explaining the observed discrepancy between the amount of upper crustal extension due to faulting and extent of crustal thinning (e.g. Bertotti et al. 2000). In addition, Royden & Keen (1980) have argued that magmatic intrusion into the crust may account for some of the inherent space problems, a mechanism that is compatible with the suggestion that the SCB is underlain by relatively dense, oceanic-type crust. Although depth-dependent stretching greatly enhances basin subsidence, the mechanism alone fails to generate the high basin depths in parts of the SCB. It is suggested, therefore, that additional processes have played a role in enhancing the subsidence within the SCB. The preferred model for the region supports evidence derived from geophysical investigation, such that the SCB crust has a density that is compatible with either an oceanic composition or with thinned continental crust that has experienced transformation to a denser phase.

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This conclusion is also compatible with a similar modelling-based investigation carried out by Green et al. (2009). They have used a combination of flexural backstripping (Roberts et al. 1998) and forward kinematic modelling approaches to show that the majority of SCB subsidence can be accounted for by a combination of thermal subsidence and sediment loading of oceanic crust. They do not, however, take into account the effects of any late-stage compressional tectonics or subduction. Model results presented here suggest strongly that the subduction of South Caspian ‘oceanic’ crust beneath the central Caspian region is necessary to generate a realistic subsidence profile and geometry across both basin and the neighbouring Apsheron Sill. Clearly, there is scope for further investigation of these processes using modelling techniques, but any significant improvement on the results presented here requires constraint from additional offshore and onshore data. We are grateful to the Middle East Basin Evolution Programme and its sponsors for funding this study from 2003 to 2005. We would like to thank BP for providing access to offshore data for this investigation. Scientific collaboration with geoscientists at the Geological Institute of Azerbaijan –National Academy of Sciences (GIA) has proven invaluable in enhancing our understanding of the structural and stratigraphic evolution of the Greater Caucasus mountain belt. We also thank N. Abdullayev and F. Lucazeau for reviewing the manuscript and providing valuable suggestions for its improvement.

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Sedimentation and subsidence in the South Caspian Basin, Azerbaijan TIM GREEN1*, NAZIM ABDULLAYEV1, JAKE HOSSACK2, GREG RILEY1 & ALAN M. ROBERTS3 1

BP Caspian Ltd, Chertsey Road, Sunbury TW 16 7LN, Middlesex, UK 2

BP, Chertsey Road, Sunbury TW 16 7LN, Middlesex, UK

3

Badley Geoscience Ltd, North Beck House, North Beck Lane, Hundleby, Spilsby, Lincolnshire, PE23 5NB, UK *Corresponding author (e-mail: [email protected]) Abstract: The South Caspian Basin is believed to contain more than 20 km of Mesozoic and Tertiary sediments deposited on oceanic or thinned continental crust. Mesozoic, Palaeogene and Oligo-Miocene sediments have not been penetrated within the South Caspian Basin itself but are exposed onshore in the basin margins. The Pliocene–Recent sequence has been mapped on a regionally extensive grid of two-dimensional (2D) seismic data and penetrated by recently drilled exploration wells, and is over 7 km thick. Most of this sequence (6 km) is formed of fluvial– lacustrine deltaic sediments of the Pliocene Productive Series that are deposited unconformably above a marine Miocene shale sequence and form the principal hydrocarbon reservoirs in the basin. The Productive Series is overlain by about 1 km of Late Pliocene–Recent marine sediments The thickness of the Pliocene sedimentary sequence implies that relatively rapid, late Tertiary subsidence occurred in the South Caspian Basin; however, there is no geological evidence of a tectonic event capable of generating a major thermal subsidence event at this time. Modelling presented in this paper suggests that it is possible to account for the observed pattern of subsidence and sedimentation in the South Caspian Basin by a process of sediment loading and compaction on a thermally subsiding, late Mesozoic crust without the need for additional Tertiary subsidence mechanisms. Crucially, this model interprets the Pliocene Productive Series to have been deposited in a topographic depression, isolated from the global oceanic system, in which base level was controlled by local factors rather than by global sea level.

Over the past decade BP has carried out geological interpretation on an extensive regional database of two- (2D) and three-dimensional (3D) seismic and well data from the South Caspian Basin (SCB), which shed light on the evolution of the basin. This paper illustrates this interpretation using a regional, balanced cross-section and isopach maps. It uses this information along with published descriptions and subsidence modelling to discuss the stratigraphic history and tectonic evolution of the SCB.

Crustal Structure Evidence for the nature of the crustal structure of the Caspian Sea comes from several sources: (1) Soviet-era deep seismic sounding studies, which indicated the basic velocity profile of the crust (Khalilov et al. 1987; Shikhalibeyli & Grigoriants 1980); (2) teleseismic-receiver function analysis, which has allowed more detailed modelling of crustal structure (Mangino & Priestley 1998); and

(3) the more recent acquisition and interpretation of deep seismic reflection data (Knapp et al. 2004). The earlier studies described the SCB area as underlain by thin, high-velocity crust, which was characterized along its northern edge by deep-focus earthquakes. These studies postulated a Benioff Zone below the Absheron Ridge (Fig. 1). The teleseismic-receiver analysis indicates that western and central parts of the South Caspian Basin are characterized by 15–20 km of lower-velocity sediment (Vp , 4.8 km s21), which overlie 10– 18 km-thick, high-velocity ‘basaltic’ crust (Vp 6.4– 7.4 km s21). This contrasts with the western Turkmenistan area, east of the Caspian Sea, where teleseismic data suggest the crust is composed of 15–20 km of ‘granitic’ upper-crustal material (Vp 4.8–6.4 km s21) and 20 km of ‘basaltic’ lower crust. Deep seismic soundings from the central Caspian Sea area, north of the Absheron Peninsula, show a fundamentally different crustal structure, more typical of continental crust, with a 2–3 km-thick sedimentary

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 241–260. DOI: 10.1144/SP312.12 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Location map of the Caspian region (the locations of cross-sections in Figs 2 and 11 are indicated).

layer, a 10 km-thick granitic layer and a 15– 20 km-thick basaltic layer. The crustal thickness increases from the central Caspian into the Absheron Peninsula and offshore Absheron Ridge, where the Moho depth is at 50 km, and then thins again into the southern Caspian where the Moho averages 34 km depth. Crustal velocities in the South Caspian Basin are thus typical of basaltic crust, and velocities typically associated with granitic crust are not observed. Mangino & Priestley (1998) conclude that the SCB is most likely to be underlain by oceanic crust, although they do not rule out the possibility of thinned continental crust. Jackson et al. (2002) described earthquake focal mechanisms and epicentre depths that show the SCB to be surrounded to the east, south and west by active earthquake belts characterized by shallow earthquakes (,30 km depth), while the northern margin of the basin, the Absheron Ridge, is characterized by earthquakes extending down to 90 km. They interpret these deeper earthquakes as indicative of the subduction of South Caspian oceanic crust beneath the Absheron Ridge. Knapp et al. (2004) describe deep seismic reflection data from the northern SCB that shows a prominent north-dipping reflector at 26–28 km, which

they interpret to be the basement–cover boundary. They interpret south-vergent folds and thrusts in the overlying sedimentary cover to have formed above a subduction zone in which oceanic crust is subducting beneath the Absheron Ridge. In conclusion, the majority of previous researchers have considered that the SCB is most likely to be floored by oceanic crust (Dewey et al. 1973; Berberian 1983; Zonenshain & Le Pichon 1986; Priestley et al. 1994; Brunet et al. 2003), although some have proposed thinned continental crust (Rezanov; & Chamo 1969; Shikhalibeyli & Grigoriants 1980).

Age and mechanism of extension A range of tectonic settings and ages have been proposed for the formation of oceanic basement in the SCB. A number of authors (e.g. Dewey et al. 1973; Kadinsky-Cade et al. 1981; Berberian 1983; Priestley et al. 1994) have considered that the SCB basement could be a relict of Tethyan oceanic crust remnant after the closure of the Black and Caspian seas during Tertiary times. Alternatively, it has been proposed that the South Caspian basement was formed in a back-arc setting north of a

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subducting Neotethyan Ocean in Mesozoic or Early Tertiary time (Berberian 1983; Zonenshain & Le Pichon 1986). Others have suggested that it may have formed in a pull-apart basin during major Cretaceous strike-slip movement in the Caucasus region (Sengo¨r 1990). Although a range of ages have been proposed for the opening of the SCB, the majority view appears to be that the basin opened at some time in the MidJurassic –Cretaceous interval (e.g. Dercourt et al. 1986; Zonenshain & Le Pichon 1986; Granath & Baganz 1996; Otto 1997; Nikishin et al. 1998; Granath et al. 2000; Stampfli et al. 2001; Ziegler et al. 2001; Brunet et al. 2003). Other authors have suggested that opening occurred in the Palaeogene (Berberian & Berberian 1981; Abrams & Narimanov 1997; Boulin 1991).

Tectonic cross-section Structural cross-section from deep reflection seismic data The structural cross-section in Figure 2a shows a part of an interpreted regional section that is derived from a series of 2D deep seismic reflection lines that extend from the southern Caspian to the central Caspian area, and have record lengths up to 20 s TWT (two-way time). The cross-section

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also integrates published information on the deep crustal structure of the area. The cross-section is structurally balanced in 2D, and the restoration of this section is shown in Figure 2b. The central Caspian sequence shown on the northern side of the section consists of a relatively thin Tertiary sequence that onlaps northwards onto a thin Triassic –Cretaceous section. This thin Mesozoic sedimentary section is underlain by Palaeozoic basement that overlies continental basement. Relatively minor localized thickening of the Triassic (and possibly Jurassic) section appears to be associated with normal faulting and small-scale rifting events. The Top Cretaceous reflector is a distinctive, continuous, high-amplitude event on seismic. This section is typical of the Scythian Platform, which existed to the north of the SCB during Mesozoic time and which formed the foreland to the Caucasus and Kopet Dagh fold-and-thrust belt. The South Caspian sequence on the south side of the section is shown to consist of an approximately 15 km-thick Tertiary sequence and 7–8 km of Upper Jurassic– Cretaceous sediment overlying oceanic basement. The basement is interpreted to be oceanic crust that dips below the Absheron Ridge in a northwards-dipping subduction zone. The central part of the section consists of the folded and faulted strata of the Absheron Ridge, which is the offshore continuation of the Greater Caucasus fold-and-thrust belt to the west and the

Fig. 2. (a) Tectonic balanced cross-section and (b) restoration through the central and South Caspian areas (see Fig. 1 for the location of cross-section).

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Kopet Dagh fold belt to the east. At its southern edge this section forms an imbricate thrust stack that verges southwards. It is composed of Late Jurassic–Tertiary sediments interpreted to have been incorporated into an accretionary prism above the subducting oceanic basement. The central part of the fold-and-thrust belt is interpreted to contain middle–late Jurassic syn-rift sediments below the Cretaceous and Tertiary sequence. This interpretation fits with the idea of a middle–late Jurassic rifting and opening of the SCB; it also helps to balance the section volumetrically. An alternative would be to interpret this deeper part of the central section as entirely Late Jurassic and Cretaceous strata. However, this would require considerably greater shortening and repetition in the Mesozoic section than is seen in the overlying Tertiary section. The northern edge of the fold-and-thrust belt is north-vergent and overthrusts the foreland of the Scythian Platform to the north. The seismic data suggests that a Jurassic –Cretaceous sequence overthrusts the Tertiary sequence along this northern boundary. Both the fold and thrust sequence in the central part of this section, and the subduction zone, result from the ongoing collision between the Arabian and Eurasian plates within the Alpine –Himalayan

Collision Zone that has affected the SCB between the Oligocene and the present day. This deformation occurred primarily in two major compressional episodes: (1) an earlier Oligo-Miocene phase; and (2) a later Plio-Pleistocene phase. In summary, the cross-section in Figure 2 can be considered to illustrate the major stratigraphic units and tectonic events in the South Caspian Basin. These tectono-stratigraphic units include: (1) the Jurassic volcanics related to the oceanic crust formation; (2) Late Jurassic and Cretaceous marine shelf and basinal sediments; (3) Palaeogene– Miocene marine rocks deformed during the Miocene compression episode; (4) latest Miocene–Pliocene Productive Series rocks; and (5) Late Pliocene– Recent marine rocks deposited during and after PlioPleistocene compression. The stratigraphic column for the South Caspian Basin is summarized in Figure 3.

Jurassic – Miocene Stratigraphy Although Jurassic sediments have not been penetrated within the South Caspian Basin, they are known from the subsurface of the Kura Sub-Basin and from outcrops in the Caucasus Mountains in the west, in the Kopet Dagh and western Turkmenistan to the east, and the Alborz to the south.

Fig. 3. Generalized stratigraphic column for the South Caspian region. The Jurassic– Miocene sequence is known from the onshore eastern Caucasus, while the Pliocene and younger section has been penetrated during drilling in the offshore South Caspian Basin.

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Jurassic volcanics Jurassic age volcanic rocks are known from the Kura Basin, where the Saatly super deep well (SD-1 TD: 8240 m) penetrated over 4400 m of Jurassic basalts, andesites, diorites and tuffs. The Jurassic volcanics lie below a 710 m sequence of Late Jurassic – Early Cretaceous oolitic and coralline limestones intruded and metamorphosed by basaltic sills. Geochemical analysis, including rare element analysis, suggests derivation from a calc-alkaline magma of island arc type affinities (Buryakovsky et al. 2001). These rocks may be the closest analogues to the pre-sediment crust in the SCB. We speculate that the Saatly rocks may have been elevated above regional level by thrusting to their present tectonic elevation, where they can be reached by the drill bit.

Late Jurassic and Cretaceous sediments onshore The Greater Caucasus Mountains to the west of the South Caspian Basin contain a thrust and folded sequence of Jurassic –Tertiary age strata. The Mesozoic sequence appears to record a shelf to basin margin transition on the northern edge of the basin that has been deformed and thrust predominantly southwards during periods of compressional tectonics in the Tertiary. Egan et al. (2009) interpreted the Mesozoic sediments exposed in the Greater Caucasus as representative of deposition in a passive margin/rift setting, with a transition from the shallow-marine carbonates that fringed the Scythian–Turan Platform to the north. They describe the Caucasus sequence as containing three main Mesozoic facies associations: (1) Late Jurassic –Early Cretaceous shallow-marine fossiliferous limestones interpreted to have been deposited in a higher-energy outer carbonate platform environment; (2) Late Jurassic dolomites and mudstone representative of low-energy, inner-shelf lagoonal deposition; and (3) Late Jurassic–Cretaceous interbedded mudstone and turbiditic sandstones, and calcarenites with abundant olistholiths and olisthostromes interpreted to represent a slope–basin environment. Subsurface mapping of onshore well data suggest thicknesses of 600–2000 m for Upper Jurassic, up to 1500 m for the Lower Cretaceous and 1500 m for the Upper Cretaceous in the Absheron Peninsula area (Geological Institute of Azerbaijan 2003). The Kopet Dagh region to the east of the SCB also contains a relatively continuous, 10 km-thick sequence of Jurassic – Tertiary sediments. This sequence includes Late Jurassic platform carbonate, Neocomian shallow-water carbonates and minor shales and evaporites, and Late Cretaceous and

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Palaeogene carbonates and marls. These sediments record a similar range of shelf –slope environments to those seen in the Caucasus (Brunet et al. 2003). The Scythian and Turan platforms, to the north of the Greater Caucasus Mountains and the Kopet Dagh region, are characterized by a thin sequence of shallow-marine carbonates and clastics of Triassic– Cretaceous age that contain numerous hiati. The region is underlain by a heterogeneous basement composed of Pre-Cambrian and Palaeozoic sediments deformed and metamorphosed in a Late Palaeozoic orogeny. These areas formed the stable inner platforms north of the subsiding shelf areas of the Greater Caucasus and Kopet Dagh and were characterized by shallowmarine –continental sedimentation during the Mesozoic.

Palaeogene sequence onshore Palaeocene and Eocene sediments are exposed in the easternmost Greater Caucasus area as the Ilhidag and Koun formations. These are predominantly argillaceous units of brown, grey and green colour. Subsurface mapping from well data shows that the Palaeocene shales range from 200 to 400 m in thickness, and the Eocene Koun Formation is typically 500–2000 m in thickness (Geological Institute of Azerbaijan 2003).

Mesozoic and Palaeogene sequence offshore The deepest continuous reflections observed on the South Caspian seismic data are a series of four–six low-frequency, high-amplitude events that occur between 10 and 12 s TWT. Although these events are not penetrated by wells, their depth suggests that they correspond to the top of the basement, i.e. the oceanic crust. Seismic mapping suggests that the depth to this event in the SCB varies from about 14 km in the southern part of the mapped basin to 25 km in the syncline immediately south of the Absheron Ridge (Fig. 4). This depth range correlates well with the depth to basement described in the teleseismic studies (Mangino & Priestley 1998) and other deep seismic studies (Knapp et al. 2004). The deepest continuous seismic event that can be mapped in the central Caspian area, i.e. from the Absheron Ridge northwards, is interpreted to be a Top Cretaceous event (Fig. 4). The next overlying seismic event, which can be widely correlated throughout the basin, is interpreted as a Base Maykop (near-Top Eocene) event. Figure 5 shows the isopach map between these two events. The isopach thins to zero north of the Absheron Ridge, where the Base Maykop event onlaps the Top Cretaceous event. To the south in the SCB the isopach map shows a

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Fig. 4. Map of depth to basement (South Caspian)–depth to Top Cretaceous (central Caspian). In the South Caspian Basin the deepest mapped seismic reflector is thought to be the top of the crystalline basement. In the Central Caspian Basin, to the north of the Absheron Ridge, the deepest mapped seismic reflector is interpreted as a Top Cretaceous event.

Fig. 5. Thickness map of the Mesozoic and Palaeogene sequences. This is a depth isochore map between the Near-Base Maykop depth surface and the Top Basement depth surface.

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pronounced thickening in the Mesozoic –Palaeogene sequence from approximately 5 km in the southern SCB to a maximum of around 20 km immediately south of the Absheron Ridge. The thickest part, a 10– 20 km-thick section, forms a wedge of sediment that extends for about 100 km ENE, paralleling the Absheron Ridge, and measures 30– 40 km in width in a NNW direction. The thickened wedge is located in the area of the presentday proposed subduction zone at the north end of the South Caspian Basin.

Oligo-Miocene sequence onshore The Oligo-Miocene Maykop, Chokrak and Diatom formations are exposed in western parts of the Absheron Peninsula and in the Greater Caucasus. They are typically composed of brown–grey, slightly calcareous marine shales with occasional thin sandstone interbeds. Intervals of locally high Total Organic Carbon (TOC) occur within the Maykop and Diatom formations, and these are the main source rocks for the oil and gas fields of the South Caspian region. In onshore exposures they can be observed to have a transitional boundary with shales of the underlying Eocene Koun Formation. Subsurface mapping of well data suggests that the Oligo-Miocene formations thicken from

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about 500 up to 3000 m from north to south across the Absheron Peninsula; however, much of the thickness variation may be tectonic in origin, as this unit is an important de´collement zone in the South Caspian structures. The Maykop and associated formations form a distinctive marine facies that is known to occur over a wide region, as far west as the Eastern Black Sea and as far north as the north of the central Caspian region. During Maykop deposition, the eastern Black Sea and South Caspian are believed to have been linked by the marine seaway of Paratethys (Steininger & Rogl 1984).

Oligo-Miocene sequence offshore Although it has not been drilled in the deep offshore, seismic interpretation suggests that the Oligo-Miocene sequence varies from about 1 km in the central parts of the SCB to 2.5 km in the north in areas adjacent to the Absheron Peninsula and Ridge (Fig. 6). Although the isopach map shows some thickening immediately to the south of the Absheron Ridge, the thickest occurrences are in the syncline to the north of the Absheron Peninsula, and in the fold-and-thrust belt onshore (where there has been significant tectonic thickening and structural repetition).

Fig. 6. Thickness map of the Oligo-Miocene sequence. This is a depth isochore map between the Base Productive Series depth surface and the near-Base Maykop depth surface.

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Late Cretaceous– Miocene compression Although the Late Mesozoic appears to have been characterized by a predominantly passive margin environment, there may be evidence of increasing tectonic activity during the Late Cretaceous and Palaeogene. In the Greater Caucasus, Egan et al. (2009) describe folded and thrusted Late Albian limestones unconformably overlain by Turonian– Cenomanian age conglomerates, as well as an Eocene erosional unconformity, which may be indicative of such compressional tectonic activity. Seismic interpretation in the offshore area immediately to the north of the Absheron Peninsula also suggests that inversion of Cretaceous extensional half-grabens was caused by local Eocene compressional activity. In general, however, widespread compressional deformation in the Greater Caucasus occurred from Pliocene time onwards, and involves the deformation of the Maykop Formation and the Productive Series (Allen et al. 2003).

Pliocene – Recent sequence Pliocene fluvio-deltaic and lacustrine sequence A major change in sedimentation occurred at the end of the Miocene when the marine Maykop and Diatom suites were succeeded by the fluviolacustrine Productive Series. This change is marked by a basin-wide series of unconformities that results from a major base-level fall and the complete isolation of the Caspian from the global ocean between approximately 6 and 5.3 Ma (Kerimov et al. 1991; Jones & Simmons 1996; Reynolds et al. 1998). This event is thought to be coincident with and connected to the Messinian Salinity Crisis in the Mediterranean Basin (Abdullayev et al. in press). The base-level fall is thought to have led to the integration of drainage systems from the Russian Platform into the Palaeo-Volga system. Interpretation of central Caspian seismic data indicates that the Palaeo-Volga river system developed a canyon 600 m deep in places and 20 km wide. The onlap of the earliest Productive Series sediments, above the unconformity, occurs 700 km to the south of the preceding Pontian shelf margin and marks the extent of the base-level fall. The isolation of the Caspian from the global oceans meant that Caspian base level was controlled by the balance between sediment and water input and evaporation; the pronounced base-level shift seen at the base of the Productive Series suggests that, during this period of isolation, the Caspian was for the most part an underfilled lake.

The Productive Series is well known from wells and exposures onshore in the Absheron Peninsula and from well and seismic data offshore. It comprises a wedge of fluvio-deltaic and lacustrine sediments that are typically between 5 and 7 km in thickness in the South Caspian Basin (Fig. 7), but which thin and progressively onlap onto the Upper Miocene age unconformity surface across the central Caspian region (Fig. 8). Abdullayev et al. (in press) divide the Productive Series into three major units. The lowermost interval, including the Kalin and PK sandstones and the lowermost Kirmaky Shale sequence, is interpreted as a lowstand fluvio-deltaic unit, backstepping to the north in a transgressive episode that culminates in a maximum flooding event (upper Kirmaky Shale). The overlying NKP Sandstone represents a highstand sequence and consists of extensive braided fluvial sediments in the north transitioning into fluvially dominated delta-top sediments in the basin centre to the south. This lowermost Productive Series unit represents an interval of progradational development of sand-rich facies across the basin. The middle part of the Productive Series is composed of the NKG Shales and the sand-rich Fasila and Balakhany formations. It is bound by an unconformity surface at the bottom, and by a major flooding event at the top. It represents sedimentation in a transgressive cycle that is accompanied by an overall fining-upwards trend. Lithologically, the unit is composed of intervals of thick channel sands interbedded with grey lacustrine shales (in the Fasila and Balakhany formations). These sands represent periods of widespread growth of fluvial and delta-plain and delta-top facies during a lowstand. This middle Productive Series unit is predominantly aggradational in type, suggesting that there may have been a long-term balance between sediment supply, water input and subsidence during this period. The upper part of the Productive Series can be divided into two units. The lower part, the Sabunchy Formation, contains stacked channelized sands interbedded with relatively thick shale intervals. The thicker sands record mouth bars and distributary channels (Reynolds et al. 1998). Offshore the Sabunchy Formation is comprised almost entirely of shale, and reflects a landwards shift in facies with respect to the underlying Balakhany suite. This relates to a further decrease in sediment input from the Palaeo-Volga during an overall reduction in depositional gradient. The overlying Surakhany Formation is about 500– 600 m in the Absheron Peninsula outcrops and reaches 2– 2.5 km in thickness in the offshore South Caspian. It is a predominantly fine-grained succession, with thick shales and isolated channelized sand

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Fig. 7. Thickness map of the Productive Series. This is a depth isochore map constructed between the Top and Base Productive Series depth surfaces.

Fig. 8. Depositional model for the Productive Series. The location of this north–south schematic diagram is shown in the inset map and runs from the north of the Central Caspian Basin to the South Caspian Basin. The Productive Series is essentially a large lowstand wedge that thickens southwards into the South Caspian Basin.

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bodies. The uppermost part of Surakhany also contains 20–40 m-thick evaporite beds that are thought to increase in thickness southwards towards the basin centre. The Surakhany Formation represents a continuation of the overall fining-upwards cycle that resulted from the progressively decreasing coarse sediment input from the Palaeo-Volga system. There was also an increased input from other clastic sediment point sources during this period; specifically from the Palaeo-Kura on the western side of the SCB and the Palaeo-Amudarya in the east. Sediment input was at its lowest in the central parts of the SCB towards the end of the Surakhany deposition, and thick evaporites accumulated in the playa lake environment present here at this time.

Late Pliocene and Pleistocene marine sequence The Productive Series is overlain by the Akchagyl Formation, which records a major regional transgression and a marine flooding event, and marks a return to open marine conditions in the SCB for the first time since the Miocene. The Akchagyl package itself is a regionally extensive, condensed

section and is downlapped by highstand, Pleistocene shelf margin complexes. These shelf margin complexes prograded into a deep-water lake (c. 1000 m) that is similar to the present depositional setting of the Caspian Sea (see Abdullayev 2000). The compressional episode responsible for the formation of the major offshore fold structures seen in the SCB was the Plio-Pleistocene event, which started towards the end of the deposition of the Productive Series. The post-Productive Series marine sediments have onlapped these Late Pliocene anticlines, and the thinning of these sediments over the Pliocene structures is very evident on the Late Pliocene–Recent isopach map (Fig. 9). In summary, when viewed at large scale, the Productive Series represents a lowstand systems tract (LST), separated by the Messinian unconformity from the highstand systems tract (HST) of the Pontian sediments below. This large-scale LST exhibits progradational facies geometries, then aggradational facies geometries and, finally, backstepping geometries. It is capped by the Akchagyl Shale, a marine flooding event, followed by another HST represented by the Absheron Formation.

Fig. 9. Thickness map of the Late Pliocene–Recent sequence. This is a depth isochore map constructed between the present-day Caspian sea floor and the Top Productive Series depth surfaces.

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Subsidence modelling Subsidence mechanisms There appears to be a broad consensus that the SCB originated as a back-arc basin during the Mesozoic – Palaeogene Period. Brunet et al. (2003) described a number of reasons that suggest the SCB opened some time between the Callovian and the Middle Cretaceous after an Early(?) –Middle Jurassic rifting event, and consider that regional evidence, indicating increased subsidence rates in the Middle Jurassic, suggest that ocean crust may have begun to form during Callovian –Oxfordian time. The subsidence curves that can be constructed by decompaction of the Jurassic –Miocene sedimentary sequence in the SCB appear to show an exponentially decreasing pattern of subsidence similar to that expected in cooling oceanic crust (Parsons & Sclater 1977). From the end of the Miocene onwards the rate of sediment deposition within the South Caspian Basin

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showed a marked increase (Fig. 10), and the thick Pliocene sequence seems unlikely to be the consequence of simple post-Jurassic oceanic thermal subsidence within a basin connected to the global marine base level. There is, however, no evidence within the sedimentary fill of the basin for a renewal of extension in the Tertiary and there is no evidence of, and too little time for, a Tertiary thermal event whose cooling could have generated an enhanced rate of subsidence in the Pliocene. Thus, in order to understand more clearly the subsidence history of the South Caspian Basin, and in particular the rapid deposition during the Pliocene, combined structural –stratigraphic basin modelling was carried out on the regional cross-section shown in Figure 2. The modelling consisted of 2D flexural backstripping (Kusznir et al. 1995; Roberts et al. 1998) and whole-lithosphere forward modelling based on a simplified (pure shear only) version of the flexural –cantilever model (Kusznir et al. 1991, 1995). The modelling process as applied to

Fig. 10. Burial history curves from the South Caspian Basin. The locations of the colour-coded curves are indicated on the inset map.

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the South Caspian cross-section consisted of four stages. 1.

2.

3.

4.

Flexural backstripping (decompaction and unloading) without reverse thermal subsidence modelling. This reveals water-loaded tectonic subsidence. Flexural backstripping including reverse thermal subsidence modelling. This attempts a full palaeogeographic–palaeobathymetric restoration of the post-rift history of the basin. Whole-crust (and lithosphere) extensional forward modelling, with thermal subsidence. The model assumes an instantaneous rifting event in the Late Jurassic. This process attempts to replicate the thickness of the Jurassic and younger basin fill in a simple time-dependent forward model. Backstripping using the new constraints of the forward model. This checks for consistency between the two modelling procedures (Kusznir et al. 1995).

The initial cross-section input into the modelling is shown in Figure 11. The main objective of the subsidence modelling was to investigate whether the thick Mesozoic –Tertiary sedimentary sequence (20 kmþ) could be accommodated as the result of subsidence (syn- and post-rift) above oceanic or ocean margin crust of Late Jurassic age, or whether additional contributions to the total

subsidence–depositional history of the South Caspian should be considered. The modelling presented here uses a principal value for effective elastic thickness (Te) of 3 km. This constrains the flexural response to loading – unloading. Such a value is generally considered typical for a standard rift basin (Roberts et al. 1998; White 1999). Sensitivity tests using a higher Te of 10 km, more appropriate for a foreland basin, and a lower Te of 0 km (no lithosphere strength) have also been run. Changing Te in this way changes local structural geometries within the models, but does not impact the first-order results about the magnitude of regional subsidence in the South Caspian Basin. The first step in the modelling process was to flexurally backstrip and decompact the sedimentary sequence to top basement at a notional 145 Ma, without incorporating reverse thermal subsidence modelling. This provides a useful initial marker in unravelling the full subsidence history as it provides a measure of total water-loaded subsidence without reference to a specific driving mechanism. This reveals that, after accounting for the effects of flexural loading and sediment decompaction, the total water-loaded subsidence of the top basement in the South Caspian Basin averages between 5.5 and 6.5 km (line co-ordinates 450–650 km in Fig. 12). The full section was then backstripped once more, this time with reverse thermal subsidence

Fig. 11. North– south cross-section used as the input model for subsidence modelling. The section runs from the Central Caspian area to the Azerbaijan part of the South Caspian area (the location of the section is shown in Fig. 1).

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Fig. 12. Sections illustrating the effects of the sediment unloading and decompaction process. The upper section shows the impact of unloading and decompaction at 5.5 Ma (around the time of deposition of the lowermost Productive Series units). The lower section shows unloading of the entire sedimentary sequence and corresponds to the bathymetry expected at 145 Ma.

incorporated in the modelling. Figure 13 shows the profile of the Late Jurassic (145 Ma) beta stretching-factor used to constrain the thermal input to the model. Across the northern part of the section (Central Caspian Basin) a relatively low beta factor (a beta factor of 2) has been applied, while across the central part of the section (the compressional Absheron Ridge –Caucasus area) a notional beta of 5 has been applied. This rises to an ‘oceanic’ beta value of 100 across the South Caspian Basin. This means that the maximum possible thermal input has been applied across the South Caspian during the backstripping process.

Figure 14a–c show the results (at three stages) of backstripping the South Caspian part of the line, with thermal input from the beta profile in Figure 13. The top basement restoration (Fig. 14c) can be compared with the restoration incorporating no thermal input in Figure 12b. With thermal subsidence now accounted for in the model, the typical residual bathymetry across the South Caspian is 2 –3 km (cf. 5.5–6.5 km in Fig. 12b). Nearly 4 km of the total water-loaded subsidence is recovered by incorporating ‘oceanic’ thermal subsidence. The bathymetric range of 2–3 km corresponds with that expected above newly-formed ocean crust

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Fig. 13. The profile of beta factors used in the model to represent a thermal subsidence event at 145 Ma. The factors range from 2, used in the Central Caspian area, to values of 100 in parts of the South Caspian Basin

(typically c. 2.5 km) and is consistent with the idea that the total South Caspian sediment thickness could be accommodated by thermal subsidence of oceanic crust. Figure 14c also shows a local area of deeper restored bathymetry (4–6.5 km), bathymetry in this part of the model is probably not realistic. It

is likely that the southern margin of the Absheron Ridge has overthrust the subducted South Caspian crust here and may have resulted in significant local tectonic thickening in the South Caspian sedimentary sequence immediately to the south. This thickening and overthrusting are not captured in the current model and the subsidence model

Fig. 14. Section, over the South Caspian Basin area only, showing the effects of sediment unloading and decompaction on a section that has a Late Jurassic thermal input. Sections (a), (b) and (c) show the effects at 5.5, 30 and 145 Ma, respectively, with an effective elastic thickness (Te) of 3 km. Section (d) shows the effect at 145 Ma of using a Te of 10 km.

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makes no allowance for such tectonic effects. Realistic modelling would require an accurate and detailed geological interpretation of the basin margin; however, this is the thickest and deepest part of the sedimentary sequence, as well as probably being the most structurally complex part, and the quality of the current seismic data makes interpretation of this deep section difficult. Figure 14d shows the results of assuming a stronger crust with an effective elastic thickening of 10 km. This section shows similar bathymetry to Figure 14c over central and southern parts of the model, but tends to accentuate the area of unrealistically deep bathymetry adjacent to the Absheron Ridge. The next step in the modelling was to attempt to produce a whole-crust –whole-lithosphere forward model capable of generating sufficient subsidence to accommodate the observed stratigraphic thickness of the South Caspian Basin between the Late Jurassic (145 Ma) and the present day. The model was generated using a modification (simplification)

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of the flexural –cantilever model (Kusznir & Egan 1989; Kusznir et al. 1995), which models lithosphere extension as pure shear only without reference to specific fault geometries. This model differs from a 2D McKenzie (1978) model in that the flexural response to loading is still incorporated. Figure 15a shows the crustal structure and subsidence of the forward model at the syn-rift stage, with no sediments added. The northern (left) side of the model shows an initial crust of 35 km thickness thinned to less than 2 km. The southern (right) side of the model shows 35 km crust thinned to approximately 11 km. The beta profile associated with this extension is shown below (Fig. 15b), rising from about 3 in the south to move than 20 in the north. The northern part of the model is almost oceanic (the model itself is not actually formulated to generate new ocean basins at infinite stretching), while the southern part of the model shows what is best considered a highly extended continental margin. At beta factors greater than 2 –3 igneous addition to the crust would be expected,

Fig. 15. (a) Diagram showing the amount of thinning applied to the continental crust in the forward model over the South Caspian Basin section. The red area shows the initial thickness and the green area shows the final thickness of continental crust component. At beta factors greater than 5 there will be significant new volcanic addition to this thinned crust. (b) Diagram showing the variation in the beta factors used in the forward modelling over the South Caspian Basin section.

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and at factors of move than 10 there should be significant addition leading to oceanic crust formation, Figure 16a shows the result of allowing the syn-rift model in Figure 15 to thermally subside for 145 Ma to the present day, with the addition of sediment up to a present-day South Caspian bathymetry of 0.5 km. The thickness of the synrift layer in the model is entirely notional, and simply draws attention to the fact that this forward model accommodates both syn-rift and post-rift subsidence in its predictions. Figure 16b shows the same forward model of total sediment thickness compared with a template

of the present-day total stratigraphic thickness that was initially input to the backstripping (Figs 2a and 16c). There is a good fit between the model and the template. This means that the total sediment thickness interpreted in the South Caspian Basin is compatible with subsidence of oceanic (or near oceanic) crust in the north adjacent to highly stretched continental crust in the south. This is not proof of oceanic crust in the northern part of the South Caspian, but it shows that the total sediment thickness geometry of this particular stratigraphic model is compatible with this hypothesis and additional subsidence mechanisms are not

Fig. 16. The diagrams shows the result of forward modelling with thermal subsidence using the beta factors shown in Figure 15b. (a) shows the thickness of sediment in blue and yellow (these are notional post-rift and syn-rift phases generated in the model). The initial thickness of continental crust prior to forward modelling is shown in red and the final thickness is shown in green. (b) shows the forward model total sediment thickness with a template (red crosses) of the present-day total stratigraphic thickness superimposed. (c) shows the present-day total stratigraphic thickness.

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required to explain the observed thickness of the full Mesozoic –present stratigraphy. In the final stage of the modelling the beta profile produced by the synthetic forward model (Fig. 15b) was used to backstrip the South Caspian section once more (Fig. 17). The northern part of the South Caspian was backstripped with effectively oceanic beta factors, while the southern part of the South Caspian was backstripped with the lower beta factors of highly stretched continental crust. The beta factor of 3 used in the south of the section generates 66% of the maximum possible thermal subsidence, while a beta factor of 20 in the north generates 95% of the maximum possible thermal subsidence (because the magnitude of thermal subsidence relates directly to the thinning factor (1 2 1/ beta) rather than to the stretching factor (beta)

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itself: McKenzie 1978). The backstripped restoration at top basement, for Te ¼ 3 km (Fig. 17b), shows a restored bathymetry across the South Caspian in the range 2.5–3 km (similar to Fig. 14c). This bathymetric range is compatible either with the bathymetry at a new ocean ridge or the water-loaded synrift subsidence at a highly stretched continental margin (see forward model, Fig. 15a). A conceptual model compatible with these observations is that the South Caspian was stretched as a continental margin passing north into new oceanic crust (or proto-oceanic crust). Both parts of the basin initially subsided to the same water-loaded syn-rift depth (c. 2.5 km), but greater stretching and thermal subsidence in the northern oceanic area carried this part of the South Caspian to a deeper present-day depth following thermal subsidence for 145 Ma.

Fig. 17. The results of sediment unloading and decompaction using the thermal subsidence modelled in the forward modelling. (a) This shows the present-day stratigraphic section and the corresponding beta factors for the stretching of continental crust that were input into the backstripping step. (b) This shows backstripping of the full sedimentary section and shows a comparable Top Basement bathymetry to that seen in Figure 14c.

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Discussion The basin modelling suggests that the full sediment thickness of the South Caspian Basin can be explained by the process of continental margin extension and possible ocean basin formation in the Late Jurassic, followed by 145 Ma of thermal subsidence, and sediment deposition, loading and compaction. There is no requirement to invoke any post-Jurassic tectonic or thermal event across the South Caspian in order to explain the observed total sediment thickness of the basin (away from the loading effect of the Absheron Ridge). What the modelling cannot account for, however, is the marked increase in the rate of deposition observed during the Pliocene. To explain this we must consider the impact that the Pliocene base-level fall in the Caspian region had on deposition within the preexisting South Caspian Basin, which was created by Jurassic extension. We suggest that Pliocene deposition may have initiated in a temporarily land-locked basin with a base level approximately 1.5 km below contemporary global sea level (a surface air load that is equivalent to a water-loaded bathymetry of c. 2.5 km). Study of the onshore exposures around the basin indicates that sedimentation during the Mesozoic – Miocene was deep marine in type and would have occurred in substantial water depths (.2 km). In the offshore area there is evidence for a significant downwards shift in base level into the Pliocene, provided by the distribution, geometry and onlap patterns of the Pliocene Productive Series. This suggests that the Productive Series was actually deposited up to 2 km below previous global sea level, despite its fluvial– lacustrine character. The key to the greatly accelerated and rapid Pliocene depositional rates and sediment loading observed in the South Caspian Basin appears to be that a very large volume of sediment was delivered over a short time to a pre-existing topographic depression, constraining a basin with a depressed base level. The primary event responsible for this rapid increase in sedimentation appears to be the Messinian global sea-level change, which resulted both in the isolation of the Caspian Basin and in a pronounced base-level shift. The base-level shift resulted in the creation of a steep depositional gradient, the expanded hinterland (Russian Platform) of sediment supply to the PalaeoVolga system, and a loss of accommodation space and sedimentary bypass in the central Caspian area. The result was that the PalaeoVolga canyon system delivered an extreme quantity of sediment to the topographic depression of the South Caspian Basin. Deposition of the Productive Series appears to have occurred so rapidly that the lower part of the series rapidly prograded southwards over the entire SCB, and so

the Productive Series does not show the pronounced south to north thickening seen in the more slowly deposited Mesozoic– Miocene sequence (compare Figs 5 and 7). The modelling supports the idea that the pronounced thickening of the Mesozoic –Palaeogene section from south to north across the South Caspian Basin (Fig. 14b) is a result of the transition from thinned continental to oceanic crust across this region, and the resulting greater beta factor and subsidence expected in the northern area of proposed oceanic crust (Figs 15–17). Another contributory factor could be that the thickening is tectonic in nature and results from thrust repetition and folding in an accretionary prism. However, the main phases of South Caspian compression are late Tertiary in age and usually involve all the Pliocene and older strata; subduction is also considered to have been a Tertiary phenomenon and the fact that subduction-related seismicity does not extend deeper than 100 km has been thought to indicate that this is a young process that has not progressed far (Jackson et al. 2002). Egan et al. (2009) have also concluded from their modelling studies that the South Caspian subsidence profile can be best explained by a model using sediment-loaded, oceanic-type crust that has been involved in subduction beneath the central Caspian region. The subsidence modelling described above does not preclude the possibility that the South Caspian Basin has been influenced by subsidence in response to compression, shortening or subduction. The clear evidence of subduction and the tectonic deformation and overthrusting of the SCB along its western and northern margins may have contributed to the subsidence of the basin. However, the modelling does suggests that it is possible to account for the SCB subsidence primarily from the combined effects of the thermal subsidence of thinned continental and oceanic crust, and the sediment loading of this thin crust.

Conclusions The basin originated as a result of Middle– Late Jurassic extension in a back-arc setting to the north of a subducting Mesozoic Tethyan ocean and island arc. Modelling suggests that continental margin extension may have progressed sufficiently far to allow the formation of oceanic crust accompanied by sea-floor spreading within a limited region in the north part of the SCB. If the suggestion is correct that only a limited amount of subduction has so far occurred beneath the Absheron Ridge (e.g. 100 km), then the area of oceanic crust would have been relatively limited in width

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(e.g. 200 –400 km) before it passed into thinned continental crust in the south part of the basin. The exposed Mesozoic sequence onshore suggests that a shelf to basin transition probably characterized the northern margin of the oceanic basin. Deep-marine, predominantly argillaceous and carbonate, deposition appears to have persisted from the Mesozoic until the Late Miocene. The isolation of the Caspian region from the global ocean system from Messinian time onwards (5.3–6 Ma) marked a dramatic change in sedimentation. The pronounced base-level shift, elevated depositional gradients for the PalaeoVolga system and the migration of the shoreline 700 km southwards meant that unusually large volumes of sediment were delivered into the lake that existed in the SCB. The basin, floored by oceanic and thinned continental crust, began to subside more rapidly under the weight of this sediment. Sediment supply appears to have outpaced subsidence at this time and the Productive Series prograded across the entire SCB. In the later stages of Productive Series deposition the coarse clastic component was deposited progressively further northwards over the Central Caspian Basin; playa-lake conditions with finer-grained sediments were more common in central parts of the SCB. Towards the end of the Pliocene the South Caspian Basin became reconnected to the global ocean system, as evidenced by the regional marine flooding event represented by the Akchagyl Shale Formation. Finally, in the Pleistocene, shelf sands prograded out into the basin to produce the shelf to basin topography (with water depths ranging from 50 to .800 m) that is observed at the present day. This paper is a part of the regional study of the South Caspian Basin that has been ongoing at BP over the last decade, and has built on the collaborative efforts of many colleagues both within BP and elsewhere, including those from the Geological Institute of Azerbaijan (GIA) and State Oil Academy. The authors particularly want to thank our co-workers in BP who have contributed to the understanding of the geology of the South Caspian Basin, and in the preparation of this paper including A. Bowman, H. Doran, S. Duppenbecker, S. Grant, B. Husseynov, N. Piggott, A. Reynolds and R. Sultanov. The authors are also grateful to M.-F. Brunet and J. Granath for their constructive reviews and suggested improvements to the final paper. Special acknowledgement also goes to M. Joly for tireless effort and invaluable assistance in drafting the figures.

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The Cimmerian evolution of the Nakhlak – Anarak area, Central Iran, and its bearing for the reconstruction of the history of the Eurasian margin ANDREA ZANCHI1, STEFANO ZANCHETTA1, EDUARDO GARZANTI1, MARCO BALINI2, FABRIZIO BERRA2, MASSIMO MATTEI3 & GIOVANNI MUTTONI2 1

Dipartimento Scienze Geologiche e Geotecnologie, Universita` di Milano-Bicocca, Piazza della Scienza 4, Milano, 20126, Italy (e-mail: [email protected]) 2

Dipartimento di Scienze della Terra, Universita` di Milano, Via Mangiagalli 34, 20133 Milano, Italy 3

Dipartimento di Scienze Geologiche, Universita` Roma TRE, Largo San Leonardo Murialdo 1, 00146 Roma, Italy Abstract: New structural, sedimentological, petrological and palaeomagnetic data collected in the region of Nakhlak–Anarak provide important constraints on the Cimmerian evolution of Central Iran. The Olenekian– Upper Ladinian succession of Nakhlak was deposited in a forearc setting, and records the exhumation and erosion of an orogenic wedge, possibly located in the present-day Anarak region. The Triassic succession was deformed after Ladinian times and shows south-vergent folds and thrusts unconformably covered by Upper Cretaceous limestones following the Late Jurassic Neo-Cimmerian deformation. Palaeomagnetic data obtained in the Olenekian succession suggest a palaeoposition of the region close to Eurasia at a latitude around 208N. In addition, the palaeopoles do not support large anticlockwise rotations around vertical axes for central Iran with respect to Eurasia since the Middle Triassic, as previously suggested. The Anarak Metamorphic Complex (AMC) includes blueschist-facies metabasites associated with discontinuous slivers of serpentinized ultramafic rocks and Carboniferous greenschistfacies ‘Variscan’ metamorphic rocks, including widespread metacarbonates. The AMC was formed, at least partially, in the Triassic. Its erosion is recorded by the Middle Triassic Ba¯qoroq Formation at Nakhlak, which consists of conglomerates and sandstones rich in metamorphic detritus. The AMC was repeatedly deformed during post-Triassic times, giving origin to a complex structural setting characterized by strong tectonic fragmentation of previously formed tectonic units. Based on these data, we suggest that the Nakhlak–Anarak units represent an arc–trench system developed during the Eo-Cimmerian orogenic cycle. Different tectonic scenarios that can account for the evolution of the region and for the occurrence of this orogenic wedge in its present position within Central Iran are critically discussed, as well as its relationships with a presumed ‘Variscan’ metamorphic event.

The Cimmerian orogeny, affecting the southern Eurasian margin between Turkey and Thailand, is related to the collision of several microplates, most of which detached from northern Gondwana in the Early Permian during the opening of the Neotethys Ocean (Sengo¨r 1979; Stampfli et al. 1991; Dercourt et al. 2000; Stampfli & Borel 2002; Brunet et al. 2003; Torsvik & Cocks 2004; Angiolini et al. 2007). The Cimmerian mountainbuilding event, resulting in the closure of the various branches of the Palaeotethys Ocean, is well documented along the northern margin of Iran, which has always been considered as one of the main Cimmerian blocks. Several authors located the main Palaeotethys suture (Alavi 1991;

Boulin 1991; Ruttner 1993) in the region around Mashhad, as first suggested by Sto¨cklin (1974). More recent reconstructions (Alavi 1991; Ruttner 1993; Alavi et al. 1997) identified an accretionary wedge in this area related to the northward subduction of a Palaeotethyan branch below Eurasia (Fig. 1). The possible continuation of the suture along the northern margin of the present-day Alborz chain (Alavi 1996) has been discussed in several papers (Zanchi et al. 2006, 2009; Berra et al. 2007) that also suggest the occurrence of an important ‘Variscan’ sensu latu tectonometamorphic event in the allochthonous nappes emplaced on the northern margin of Iran during the Eo-Cimmerian collision (Zanchetta et al. 2009).

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 261–286. DOI: 10.1144/SP312.13 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Tectonic scheme of Iran with the main tectonic subdivisions. White, Cimmerian blocks; black, Mesozoic ophiolites along the Main Zagros and other sutures; modified from Angiolini et al. (2007). AA, Araxian–Azarbaijanian Zone; Ag, Aghdarband Basin; Ak, Anarak; GKDF, Great Kavir–Doruneh fault system; Hj, Hajiabad metamorphic complex; KDF, Kopeh Dagh Foredeep; KF, Kalmard Fault; Ks, Kor-e-Sefid metamorphic complex; MZT, Main Zagros thrust; Nk, Nakhlak Basin; Ny, Neyriz ophiolites; PB, Posht-e Badam; PTSZ, Palaeotethys Suture Zone; Sg, Saghand; SSZ, Sanandaj– Sirjan Zone.

The Palaeotethys collision zone is also exposed eastward along the Paropamisus and Northern Pamir ranges (Boulin 1988, 1991). The Gondwanan affinity of the blocks forming Iran is supported by several lines of evidence (Sto¨cklin 1974; Berberian & King 1981), based on the similarity of their late Proterozoic ‘Pan-African’ basement, the affinity of the late Pre-Cambrian and Cambrian successions across the Zagros suture, and the lack of a ‘Variscan’ deformation (Saidi et al. 1997). The Gondwanan affinity of Iran is also demonstrated by U–Pb geochronological data on granitoids belonging to the crystalline basement of Sanandaj– Sirjan, central Iran and the Alborz Mountains, which have given late Neoproterozoic–Early Cambrian ages (Hassanzadeh et al. 2008). Their presence in Iran is explained by the development of an active continental margin in the Peri-Gondwana region at

the end of the Neoproterozoic, as also suggested by the geochemical affinity of these rocks (Ramezani & Tucker 2003). The final collision of the Iranian composite plate with Eurasia occurred during the Late Triassic, and is marked by a regional unconformity sealed by the deposition of the Upper Triassic– Jurassic Shemshak Group (Fu¨rsich et al. 2009; Zanchi et al. 2009). U –Pb zircon ages of 220–210 Ma (Horton et al. 2008) from this clastic succession provide control on the time of the collision. Other authors (Boulin 1988, 1991; Golonka 2004) proposed a different scenario, hypothesizing that Iran was composed of different microplates (Alborz –north Iran and central Iran–Lut) that detached from Gondwana at different times and progressively collided with the composite southern Eurasian active margin during the Mesozoic.

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

The Nakhlak –Anarak region, located in the middle part of central Iran south of the Great Kavir Fault (Fig. 1), is thus a key area in understanding the Late Palaeozoic–Triassic evolution of Iran. In fact, the peculiar arc-related Triassic succession of Nakhlak shows contrasting features with the dominant carbonate platform facies widespread from the Alborz Mountains to the northern margin of the Jaz Murian depression (Alavi et al. 1997; Seyed-Emami 2003). The Olenekian –Ladinian succession consists of pelagic sedimentary rocks including Rosso ammonitico facies, volcanic arenites, alluvial fan conglomerates and turbidites deposited in a volcanic arc setting along an active margin (Balini et al. 2009). A few kilometres to the south of Nakhlak, the Anarak Metamorphic Complex (AMC) is exposed. It consists of an intricate rock association including metapelites, metabasites and marbles associated with slivers of ultramafic rocks (Sharkovski et al. 1984), interpreted as an accretionary wedge active from Late Palaeozoic– Triassic times (Bagheri & Stampfli 2008). The AMC is sharply interrupted westward by the Upper Cretaceous ‘Coloured Me´lange’, forming a more recent ophiolitic ‘ring’ that surrounds the internal portion of central Iran. The occurrence of this peculiar rock association within continental Iran poses several questions regarding its evolution and especially on the number of Cimmerian (Palaeotethys) sutures (single rather than multiple) between Eurasia and Iran (Sto¨cklin 1974; Sengo¨r 1979, 1990; Berberian & King 1981; Alavi 1991; Stampfli & Borel 2002; Bagheri & Stampfli 2008). Previous authors (Davoudzadeh & Weber-Diefenbach 1987; Soffel et al. 1996; Alavi et al. 1997) explained the anomalous position of the Nakhlak–Anarak region in terms of sparse palaeomagnetic data that apparently showed a post-Triassic 1358 anti-clockwise rotation of central Iran. That rotation was supposed to be responsible for displacing a large fragment of the Palaeotethys suture from the present-day Afghan– Iranian border region. However, new palaeomagnetic information obtained from the Olenekian succession of Nakhlak (Muttoni et al. 2009) challenges that interpretation and suggests that no large rotations have occurred in the area since the Triassic. In this paper, we describe the structural setting of the Nakhlak–Anarak region and outline aspects of the evolution of the region that remain problematic.

How many blocks form Iran? A basic problem in the literature is the definition of the different Late Palaeozoic–Middle Triassic domains that comprised Iran. The idea that Iran consists of more than one block is directly related

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to its complex setting, chiefly due to the occurrence of several ophiolitic belts that are assumed to record the opening and closure of several oceanic basins, and to the activity of large active intracontinental faults that split the area in several partially independent blocks. The evolution of Iran through time is complex owing to the continuous aggregation and differential motion between several blocks. According to the first reconstructions presented by Sto¨cklin (1974) and Berberian & King (1981), three main regions can be recognized during Palaeozoic –Triassic times: North Iran, Central Iran and the Sanandaj–Sirjan Zone. Following Berberian & King (1981) and the main subdivisions adopted in the Geological Survey of Iran’s 1:250 000 maps of Iran, north Iran includes the Alborz region, its boundary with central Iran running just to the south of the mountain belt. The Alborz belt is a Late Tertiary intracontinental belt (Allen et al. 2003) reactivating a Late Triassic orogen which itself was produced by the collision of Iran with the southern margin of Eurasia (Zanchi et al. 2009). The late Precambrian –Triassic succession is very similar to the sequences found in the other regions of Iran and, at least for a large part of the Palaeozoic, shows a Gondwanan affinity (Sto¨cklin 1974; Berberian & King 1981; Wendt et al. 2005; Angiolini et al. 2007). Central Iran occupies the internal part of Iran, and is a very complex and poorly known area, characterized by desert areas with poor exposures. The most peculiar feature is the occurrence of an Upper Mesozoic ophiolitic ‘ring’, the so-called ‘Coloured Me´lange’, which delimits its most internal part (Figs 1 and 2). Central Iran is also affected by a complex system of active intracontinental strike-slip faults (Fig. 1) causing an intensive north –south dextral shearing of the whole area (Walker & Jackson 2004). The left-lateral Great Kavir –Doruneh fault system to the north, crossing the northern part of central Iran, currently bounds the deformational system to the north. The deformation is accommodated within central Iran by N –S-trending dextral faults, which separate the Yazd, Tabas and Lut blocks. These faults are possibly inherited from Precambrian times (Berberian & King 1981) and were active during the Palaeozoic evolution of the region, as testified by facies and thickness variations across these structures (Wendt et al. 2005). The boundary between North and Central Iran coincides with minor facies changes in the Palaeozoic –Triassic successions. In addition, Leven & Gorgij (2006) and Gaetani et al. (2009) show a striking similarity between the PermoCarboniferous stratigraphic evolution of North and Central Iran, suggesting the absence of biotic barriers and thus their contiguity. Minor differences

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Fig. 2. General structural map of part of Central Iran, obtained from available maps (Geological Survey of Iran 1976, 1977, 1984a– c, 1987, 1989; Bagheri & Stampfli 2008) and from Ramezani & Tucker (2003). The square shows the location of Figure 3. Jn-Sr, ZK-Sr, ophiolitic fragments of Jandaq-Arusan and Zeber Kuh; Ch, Chapedony Complex (Eocene). Ca, Cambrian; T, Late Triassic; J, Jurassic; K, Cretaceous; P, Palaeocene; Eo, Eocene refer to radiometric ages of granitoids; P-B, Posht-e-Badam.

occur in the Abadeh region, located along the eastern border of the Sanandaj–Sirjan Zone. North Iran thus represents the northern margin of the Central Iran block at least for the Palaeozoic. The same conclusion is reached by Wendt et al. (2002, 2005), who

extensively reviewed the Palaeozoic successions of Iran. According to their work, most of the area, with a few exceptions in the Talesh Mountains (see also Zanchi et al. 2009 for a discussion) and ¯ ghdarband basin), southern east of Mashhad (A

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

Alborz, as well as the Yazd, Tabas and Lut blocks, belongs to the stable northern margin of Gondwana. No clear separation is evident during most of the Late Palaeozoic. The same authors exclude strongly subsided Devonian basins that were later affected by Variscan orogenic movements (‘mobile zones’) along the western margin of central Iran and in the southern Sanandaj–Sirjan Zone, as did Davoudzadeh & Weber-Diefenbach (1987). These authors, in fact, recognize the occurrence of isolated blocks within central Iran that were affected by a Variscan metamorphism: the Anarak metamorphic rocks, some of the metamorphic rocks exposed in the Saghand area, which have given a Rb–Sr age of 315 + 5 Ma (Crawford 1977), and possibly also the Zeber Kuh beds (Fig. 2) and the Carboniferous phyllites located NE of Tabas and SW of Kashmar. Two hypotheses are discussed by Davoudzadeh & Weber-Diefenbach (1987). † Occurrence of multiple Palaeotethyan basins separating blocks with different Palaeozoic histories. In their opinion, the Palaeotethys suture located along the northern margin of Iran represents the remnants of the main ocean, whereas a narrow trough was located south of Jaz Murian and the Sanandaj–Sirijan Zone; another trough was located along the Doruneh–Great Kavir fault system and is testified by the occurrence of Variscan metamorphic events. † The Palaeotethys suture was unique, and metamorphics of the Zeber Kuh– Anarak area were displaced from their original position previously located between Mashhad and Herat during a presumed Meso-Cenozoic anticlockwise rotation of the Central Iran microplate. Bagheri & Stampfli (2008) document the occurrence of a Late Palaeozoic –Triassic active margin in the area south of the Great Kavir Fault between Anarak and Jandaq, implying a sharp separation between the regions of Central Iran located north and south of this fault. These authors present new palaeontological and Ar– Ar age determinations documenting the evolution of a long-lived active margin between 330 and 230 Ma. They interpret the arc–trench successions as fragments of the Palaeotethys suture zone originally located E–NE of Mashhad and tectonically transported to its present position after the Cimmerian orogeny during the counter-clockwise rotation of central Iran. A further problem in understanding the structure of Central Iran is caused by the findings of Davidof & Arefifard (2007), who identified cold-water fusulinid assemblages in the Posht-e-Badam area along the Kalmard Fault (Fig. 2), west of Tabas. These are typical cool-water low-diversity peri-Gondwana assemblages, which are distributed from Oman to Karakoram, Central Pamir and the Baoshan block

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in China. This might indicate that the block lay far south during the Early Permian with respect to North and Central Iran, which show typical warmwater tropical fusulinids (Leven & Gorgij 2006; Angiolini et al. 2007). The Sanandaj–Sirjan Zone (SSZ) is a narrow belt of strongly deformed and metamorphosed rocks sandwiched between the Zagros Orogenic Belt and Central– North Iran (Berberian & King 1981). The SSZ is partially delimited on both sides by supra-subduction ophiolites formed in intra-oceanic island arcs (Ghasemi & Talbot 2006). It mainly consists of Palaeozoic–Mesozoic metamorphic rocks intruded by several generations of Meso-Cenozoic intrusives. Its history culminated in the Eocene Urumieh–Dokhtar magmatic arc that records the main Neotethys slab break-off. Evidence of the origin of the SSZ and of its evolution is controversial, mainly owing to strong tectonic disturbance and metamorphism that often hampers the recognition and dating of the protoliths. A major branch of the Permian Neotethys oceanic basin separated the SSZ block from the present-day Zagros region, which consists of deformed sedimentary rocks entirely belonging to the Arabian margin. Some authors (Sto¨cklin 1974; Berberian & King 1981; Ghasemi & Talbot 2006) interpret the SSZ as a block lying along the southern margin of Iran. Recent stratigraphic studies by Wendt et al. (2005) south of Kashan indicate that the SSZ was part of Central Iran during the Palaeozoic, showing very similar stratigraphic units. Recent data concerning the study of the protoliths of the Mesozoic metamorphic rocks exposed in the Muteh area, NW of Isfahan, also suggest that they represent a Palaeozoic succession consistent in age and composition with the coeval units of Central Iran (Rachidnejad-Omrani et al. 2002). In contrast, Ghasemi et al. (2002) identified in the southern part of the region, west of Hajiabad, a metamorphic complex consisting of garnet– gneiss amphibolite and anatectic granites that show Late Carboniferous ‘Variscan’ Ar –Ar ages ranging from 330 to 300 Ma. A possible Late Palaeozoic event followed by an Eo-Cimmerian deformation is also reported by Sengo¨r (1990 and references therein) in the Khor-e Sefid Mountains east of the Neyriz ophiolites. Sheikholeslami et al. (2003), based on several new K –Ar radiometric ages and on structural analyses performed in this area, favour the occurrence of a unique Eo-Cimmerian metamorphic event, although older K –Ar ages of 310 + 9 and 331 + 5 Ma on single mineral phases belonging, respectively, to a metagabbro and to an anatectic granite were also found. According to other interpretations (Sengo¨r 1990; Bagheri & Stampli 2008), the SSZ was probably located more westward, and its present position

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was achieved after the Triassic due to large left-lateral strike-slip orogen-parallel tectonic transport. However, detailed field structural data (Mohajjel & Fergusson 2000) describe important dextral transpression resulting from convergence during Tertiary and possibly also during Cretaceous times. In addition, Ghasemi & Talbot (2006) suggested at least 100 km of dextral shearing along recent faults between Arabia and the Zagros region. Sengo¨r (1990) hypothesized that the SSZ was part of a Carboniferous magmatic arc (the Podataksasi Zone) formed along the northern margin of Gondwana due to a presumed southward subduction of the Palaeotethys Ocean. According to this author, this presumed southwards subduction might also explain the opening of the Neotethys Ocean in a back-arc position between the Arabia, now represented by the Zagros belt, and the Cimmerian blocks. This hypothesis has been rejected by most of the authors for lack of geological evidence. The SSZ is currently divided at the latitude of Golpayegan into a northern and southern portion with different features. The northern SSZ was only affected by an intra-continental rift filled with thick Triassic– Jurassic deposits, whereas the southern SSZ was separated from central Iran along the Nain –Baft Ocean from the Triassic onwards. Both, however, were separated from the Arabian Plate by the opening of Neotethys since the Early Permian. That separation is recorded by Upper Permian –Triassic ophiolites of the Sikhoran Complex (Geological Survey of Iran 1993; Ghasemi et al. 2002; Sheikholeslami et al. 2003; Ghasemi & Talbot 2006), exposed east of Hajiabad (Fig. 1). They include mafic–ultramafic layered intrusive rocks with a tholeiitic affinity that also record an Eo-Cimmerian deformation unconformably covered by Jurassic sediments. According to Berberian & King (1981), the SSZ has remained the active margin of the Iran Plate since the end of the Triassic, and associate it with a subduction shift following the closure of the Palaleotethys Ocean to the north. Inversion of the oceanic basin began in the southern SSZ during Late Jurassic–Early Cretaceous (Neo-Cimmerian event), whereas in the northern part of the SSZ compressional deformation started during Cretaceous times.

Stratigraphic setting of Nakhlak The record of a Triassic active margin crops out at Nakhlak in Central Iran (Fig. 3), in contrast with the carbonate-platform dominated sedimentation (Elika¯ Formation, Shotori Formation, etc.) that characterizes most of the Iranian region during the Early– Middle Triassic (Alavi et al. 1997; Seyed-Emami 2003; Balini et al. 2009). The approximately

2400 m-thick succession (Olenekian–Ladinian) of Nakhlak includes three main stratigraphic units. The base of the Triassic succession is not exposed; it is separated by an important low-angle fault zone from coarse-grained hornblende – gabbro in close associations with small patches of strongly serpentinized and altered ultramafic rocks (Fig. 4). The Triassic at Nakhlak commences with the Olenekian–Anisian Alam Formation (1150 m thick), which consists of fine volcanic arenites, nodular limestones and marls with turbidites at the base. It was deposited in a marine setting and records a succession with alternating deepening and shallowing trends. The sedimentation was mostly volcaniclastic, but with three calcareous-dominated intervals. The upper part of the formation displays a shallowing-upwards trend passing into transitional–lagoonal sedimentation (Unionites assemblages) with increasing intercalations of fluvial conglomerates. The ?Late Anisian –Ladinian Ba¯qoroq Formation, 870 m thick, represents an alluvial fan deposited in a fluvial environment. The unit thins out northward and changes composition, passing from conglomerates to sandstones. It documents the erosion of upper-crustal metasedimentary rocks. Detailed petrographic analyses of the Ba¯qoroq sandstones and clasts will be discussed later in order to compare their composition with the metamorphic rocks exposed to the south in the Anarak Massif. The Late Ladinian, more than 370 m thick forms the Ashin Formation. It consists of fine-grained turbidites, and documents a rapid evolution from the continental and fluvial environment of the Ba¯qoroq Formation to a deep-water turbiditic system. Its age might reach up to the base of the Carnian (Alavi et al. 1997). The sandstone composition suggests a renewed volcaniclastic input, alternating with debris from metasedimentary rocks. An Upper Cretaceous – Palaeogene succession, including massive limestones at the base, and followed by marly limestones and red sandstones at the top, unconformably covers the deformed Triassic units.

Structural setting of Nakhlak Structural analyses have been performed in the Mesozoic successions of Nakhlak in the whole area of exposure (Fig. 4). The overall tectonic setting of the Nakhlak structure is shown in a general geological cross section based on our original field observations (Fig. 4). Structural investigations agree with the geological maps of Iran (Geological Survey of Iran 2003), and allow definition of the well-bedded Upper Cretaceous limestones that unconformably cover the deformed

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Fig. 3. Geological map of the Nakhlak– Anarak region, simplified from Geological Survey of Iran 1:250 000 and 1:100 000 maps of the region (Geological Survey of Iran 1984a, b, 2003), with the location of the two cross-sections of Figures 4 and 8. The locations of studied samples are also shown.

Triassic sequence (Figs 5A, B and 6A, D), contrary to what has been suggested by other authors (Alavi et al. 1997). Small reverse faults reactivating the discontinuity are observed only at local scale (Fig. 6B) and do not deform the contact between these units. Our fieldwork shows that the Nakhlak structure is much more complex than previously described. The Triassic succession is affected by WNW –ESEtrending folds and reverse faults. In the southernmost part of the area, the Triassic succession is thrusted onto poorly exposed amphibole-metagabbros and

serpentinites. These rocks are interpreted as a dismembered ophiolitic complex (Nakhlak-type ophiolites of Bagheri & Stampfli 2008), although no clear evidence for an oceanic lithosphere was given. The Upper Cretaceous beds also unconformably cover this mafic to ultramafic complex, indicating that they were juxtaposed to the Triassic successions before their deposition. The southern part of the Nakhlak area shows the most complete succession that consists of the Alam, Ba¯qoroq and Ashin formations. A well-developed axial plane cleavage occurs in the southern part of

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Fig. 4. Idealized geological cross section at Nakhlak, based on our original geological and structural observations; trace of the section in Figure 3. The topography is poorly constrained; vertical and horizontal scales are the same ones. Stereographic Schmidt’s projections, lower hemisphere, refer to the measured structural elements. Faults are represented as cyclographic projections with black dots representing striations with relative sense of motion; small empty circles are poles to bedding; grey circles are poles to axial plane fracture cleavage; triangles are poles to axial planes; and fold axes are black small circles.

the area in the Alam Formation. A pressure-solution cleavage is present in the limestone beds, whereas a fracture cleavage formed in the marly layers of the upper part of the formation. Axial plane cleavage also strikes WNW –ESE and dips to NNE. In the central part of the Nakhlak area, the Ashin Formation shows complex disharmonic and conical folds with steeply plunging axes, possibly related to bedding-parallel detachments along shaly layers. Owing to these reasons, these folds cannot be used to determine the direction of tectonic transport. The succession is interrupted northwards by an important WNW–ESE south-verging reverse fault. Tight folds in the footwall in the Ashin Formation show a strong dispersion of their axis trend (site Nakhlak 6; Fig. 4). This suggests that the folds formed before faulting and that they were reoriented due to the fault motion. In the hanging wall, the Ba¯qoroq and Ashin formations are completely overturned and show a small synformal structure parallel to the fault strike (Fig. 4). Overturned beds of the Ashin Formation are indicated by the orientation of flute casts; the overturned structure had been recognized several years ago by Holzer & Ghasemipour (1969). Present-day thickness of the Ashin Formation may reflect tectonic overthickening of the unit along both sides of the fault. To the north the Ashin Formation is crossed by strike-slip faults along which folds with vertical axes formed (Fig. 5B, 5C). In the northern part of the area, the

Alam Formation overthrusts the Ashin Formation towards the south along another north-dipping high-angle reverse fault. The Upper Cretaceous limestones unconformably cover most of the described structures. This suggests that folding and thrusting occurred between the Late Triassic and the end of the Early Cretaceous, and that these tectonic structures can be related to the Cimmerian orogeny sensu latu. The Upper Cretaceous beds are weakly deformed. The limestones are affected by a N –Strending asymmetric fold exposed along the east side of the area. The small N –S trending low-angle thrust faults that cut the unconformity in the southernmost part of the area can be interpreted as secondary shear structures related to the growth of the main fold. An important N– S-trending strike-slip fault crosses all of the described structures. Mesoscopic structures record a dextral strike-slip motion of the fault (Fig. 6E). The strong hydrothermal circulation that occurs along the fault is possibly related to the formation of the galena –cerussite ore deposits (Holzer & Ghazemipour 1969) of the Nakhlak mine. The fault shows an important vertical offset of the Cretaceous beds (several hundred metres), which are in contact with Palaeocene red conglomeratic layers. This fault belongs to the system of large N – S- to NNE–SSW-trending right-lateral faults that cross central Iran (Walker & Jackson

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Fig. 5. Field views of the Nakhlak structure. Angular unconformity between the deformed Triassic succession and the Upper Cretaceous limestones in the southern (A) and northern part (B) of the area. Tectonic repetitions of the Ba¯qoroq Formation are evident below the unconformity in both views. (C) close view of the complex fold with vertical axes probably related to Tertiary strike-slip faults superposed to the Mesozoic tectonic structure.

2004). The fault extends between the active Deh Shir and Anar faults, which show a similar trend and kinematics (Walker & Jackson 2004). This N–Strending right-lateral fault might be also related to the formation of the previously described N–S flexure, which can be interpreted as an in-line forced fold growing in a rigid basement before the final activation of the dextral shear zone. Examples of this kind are known along large strike-slip structures (Woodcock & Schubert 1995).

Petrography of Triassic Nakhlak sandstones A detailed study of the petrography of the Nakhlak succession is reported in Balini et al. (2009) and is here summarized in its main results. Sandstones of the lower part of the Alam Formation are volcanic arenites dominated by microlithic–felsitic lithic fragments and plagioclase, and containing quartz, granophyric–granitoid rock fragments and

chessboard albite. Such composition documents provenance from a magmatic arc dissected to various degrees, with sporadic contribution from encasing metamorphic wallrocks (‘Magmatic Arc Provenance’ of Dickinson 1985; Marsaglia & Ingersoll 1992; Garzanti et al. 2007). Relatively deep erosion levels within the arc massif are indicated by coarse-grained basal layers (‘Dissected Arc’ subprovenance), followed by overwhelming volcanic detritus indicating a major phase of intermediate volcanism (‘Undissected Arc’ subprovenance). Overlying sandstones include mica and metasedimentary rock fragments, pointing to an incision phase with subordinate supply from metamorphic wallrocks. A second volcanic cycle is suggested by overlying sandstones, initially derived largely from plagioclase-rich tuffs and next mainly including detritus from relatively mafic products. The unit is capped by sandstones relatively rich in quartz, chessboard albite, K-feldspar and

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Fig. 6. Field views at Nakhlak. (A) Stratigraphic angular unconformity with a pronounced erosion channel between the Alam Formation and the Upper Cretaceous; (B) local tectonic reactivation of the same surface unconformity of (A); (C) fracture cleavage (S1) in marly layers of the Alam Formation (S0); (D) panoramic view of the unconformity between the Triassic Alam Formation and the Upper Cretaceous limestones in the southern part of the Nakhlak area; (E) dextral shearing along the strike-slip fault F.4 of Figure 4 indicated by oblique pressure-solution cleavage; and (F) clasts of metamorphic rocks in the upper bed of the Ba¯qoroq Formation (gneiss, greenschists, quartzites).

granophyric–granitoid rock fragments, suggesting a phase of ceased volcanism and renewed erosion. Similar composition characterizes the basal layer of the overlying Ba¯qoroq Formation, whereas sandstones throughout the main body of the unit consist of quartz, metamorphic rock fragments, feldspars, granitoid rock fragments and micas, with only a few volcanic–subvolcanic rock

fragments. Dominant medium-rank –high-rank schistose–gneissic rock fragments (metamorphic index (MI) 338 + 15: Garzanti & Vezzoli 2003) indicate provenance from low-grade metasedimentary rocks (Fig. 6F). Such remarkably homogeneous quartzo-lithic metamorphiclastic composition documents a drastic change in provenance, revealing rapid erosion of a metamorphic

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complex (‘Recycled Orogen Provenance’ of Dickinson 1985; ‘Axial Belt Provenance’ of Garzanti et al. 2007). Occurrence of serpentinite clasts suggested by Alavi et al. (1997) is not confirmed by our study. The Ashin Formation includes mainly quartzolithic metamorphic clastic sandstones similar to those of the Ba¯qoroq Formation, intercalated with volcanic arenites. Volcanic arenites consist of virtually pure volcanic detritus, containing monocrystalline quartz with straight extinction and embayed or pyramidal outlines, and microlitic–vitric and pyroclastic grains, testifying to a phase of renewed intermediate–felsic explosive volcanism. Absence of sediments with mixed volcanic–metamorphic signatures (volcanic arenites invariably contain negligible amounts of metamorphic grains, and quartzo-lithic sandstones invariably include negligible amounts of volcanic detritus) indicate two different sources situated in geographically distinct areas. Conglomerates are common in the upper part of the Alam Formation (Late Bythinian–Late Anisian), the Ba¯qoroq (Lower Ladinian) and the lower part of the Ashin Formation (Late Ladinian). Pebbles of quartz, metamorphic rocks and granitoids occur at the base of the Ba¯qoroq Formation, and become gradually more abundant upwards. Volcanic rocks are also common. Pebbles of sedimentary rocks are found only in the lower part of the Ba¯qoroq Formation. Field analysis of pebble composition is summarized in Figure 7.

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Provenance interpretation The petrography of Triassic Naklakh sandstones, interpreted in the framework of available stratigraphic and geological information, support the following considerations: † the succession was deposited in a rapidly subsiding sedimentary basin during Early–Middle Triassic times; † dominant magmatic arc provenance for Alam sandstones and some Ashin sandstones indicate deposition in an arc-related, supra-subduction setting; † acidic–intermediate magmatism occurred in several distinct pulses, and was probably characterized by orogenic affinity varying irregularly in time from latite–andesite (lower Alam Formation) to basaltic andesite (upper Alam Formation) to dacite–rhyodacite (Ashin Formation); † quartzo-lithic metamorphic clastic signatures for Ba¯qoroq sandstones and many Ashin sandstones may indicate provenance from a metamorphic axial zone of an orogenic prism, exhumed and eroded at Middle Triassic times (Late Anisian –Early Ladinian); † active volcanic and metamorphic axial sources of detritus co-existed at Late Ladinian times, at opposite sides of the basin. Based on such reconstruction, the Nakhlak succession is interpreted as deposited in a forearc

Fig. 7. Cumulative bar diagrams for the observed clast compositions in the Triassic of Nakhlak. Petrographic composition of the pebbles of the conglomerates of the upper part of the Alam, Ba¯qoroq and lower part of the Ashin formations.

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basin. A significant orogenic event is suggested at Late Anisian times, when low-grade metamorphic rocks were exhumed and eroded. At Late Ladinian times, volcanism resumed in the magmatic arc.

The Anarak Metamorphic Complex and the metamorphic units of central Iran The Anarak Metamorphic Complex (AMC), firstly described by Sharkovski et al. (1984), forms an east– west-trending mountain ridge that separates the Triassic of Nakhlak to the north from a continuous non-metamorphic Palaeo-Mesozoic sedimentary succession to the south. The succession to the south belongs to the Yazd block of Central Iran (Fig. 3). The sequence shows features similar to other Palaeozoic exposures described in North and Central Iran (Wendt et al. 2005; Leven & Gorgij 2006). The AMC trends E –W and is sharply juxtaposed westward against the ‘Coloured Me´lange’ of the Nain –Baft Mesozoic ophiolites (Ghasemi & Talbot 2006) by the western termination of the Great Kavir Fault. To the east of Anarak, the AMC crops out discontinuously below the MesoCenozoic cover for about 150 km through Central Iran between Jandaq, Anarak and Khur (Fig. 2). It is associated with medium- to high-grade metamorphic rocks and ophiolitic remnants, exposed just south of the Great Kavir Fault (Bagheri & Stampfli 2008) (Fig. 2). The AMC is subdivided into several subunits (Sharkovski et al. 1984), namely the Morgha¯b, Cha¯h Gorbeh, Kabudan, Patya¯r and La¯k Marble, which all show different metamorphic evolutions. The Cha¯h Gorbeh unit includes phyllites, mica schists and greenschists interlayered with thick marble layers and calcareous schists. Glaucophaneand lawsonite-schists are described in this unit (Sharkovski et al. 1984); blueschists also occur in the Patya¯r and Kabudan units, which have a similar composition and are exposed near Khur, 50 km to the east. The La¯k Marble includes a thick succession of meta-carbonates with a few metapelitic and quartzitic intercalations. Sharkovski et al. (1984) describe this unit as a large nappe exposed in isolated klippen overlying the other metamorphic units. A few remnants of archaeocyathids (Paranacyathus sp., Coscinocyathus sp., Agastrocyathus sp., Dicthyocyathus sp.; Geological Survey of Iran 1984a) give an Early–?Middle Cambrian age to the La¯k Marble, whereas crinoids and gastropods dubitatively suggest a younger age (Sharkovski et al. 1984). Mica schists, gneisses, orthogneisses and amphibolites form the Palha¯vand Gneiss, which is separated by the Biabanak Fault from the Palaeozoic succession of the Yazd block exposed south of Anarak (Fig. 3).

Previously reported single phases and whole rock K –Ar radiometric ages (Sharkovski et al. 1984) are scattered between 420 and 208 Ma, with the main cluster between 300 and 375 Ma. The dates have been interpreted as the result of superposed Variscan and Late Triassic Cimmerian metamorphism. Triassic ages (222 + 10 and 208 + 8 Ma) have been reported by the same authors from the Kuh-e Dom Complex, which contains Palaeozoic fossils and is located NW of the Great Kavir Fault and of the Nain ophiolites (Fig. 2). Reyre & Mohafez (1970) obtained a Rb–Sr 203 + 13 Ma radiometric age from metamorphic rocks of the AMC around Anarak, which also suggests a Late Triassic metamorphism. Evidence of possible Precambrian protholiths in the AMC is given by a Rb–Sr age determination of 845 Ma (Reyre & Mohafez 1970). Additional stratigraphic constraints on the age of deformation of the AMC occur in the Khur area. Here the Shemshak Group overlies the AMC and related units affected by low-grade metamorphism with sharp nonconformity, suggesting an important Eo-Cimmerian tectono-metamorphic event (Sharkovski et al. 1984). In the whole region, a successive NeoCimmerian event is marked by the deposition of the Low Cretaceous Cha¯h Palang Conglomerate lying with a strong angular unconformity on older units, including the Shemshak Group. A very lowgrade metamorphic overprint on the Shemshak and Middle–Late Jurassic granitic rocks (Geological Survey of Iran 1972, 1976) around Torud, to the north and SW of Yazd (Fig. 2), suggests an even later thermal event. Elongated slivers of mafic and ultramafic rocks (‘Anarak-type ophiolites’ of Bagheri & Stampfli 2008), up to 5 km wide and 20 km long, are tectonically intercalated within the metamorphic rocks of the AMC. Ultramafic rocks occur especially along the northern slopes of the Anarak Massif. Small patches of serpentinites also crop out just to the north of the Biabanak Fault in the southern part of the area (Fig. 3). Most of the rocks can be classified as serpentinized harzburgite and olivine orthopyroxenite, occasionally with layered structures (Sharkovski et al. 1984). Undeformed plagiogranite (trondhjemite), quartz-diorite and tonalite occur as well. Bagheri & Stampfli (2008) report a 262+1 Ma U– Pb zircon age for a trondhjemite intruded in the Cha¯h Gorbeh unit, and consider that it predates the blueschists-facies metamorphism of the unit. Bagheri & Stampfli (2008) proposed a significantly new scenario, which implies the occurrence of a polyphase accretionary wedge that underwent metamorphism during Carboniferous times (‘Variscan event’), and also subsequent deformation during Permian and Triassic times in the region

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

south of the Great Kavir Fault. The La¯k Marble, the Patya¯r unit and the ophiolitic remnants (‘Anaraktype ophiolites’) in this reconstruction represent the relicts of a seamount characterized by an oceanic basement (ultramafic and metavolcanic rocks of the ‘Anarak-type ophiolites’) overlain by reefal carbonates (La¯k Marble). Metadolomites, metacherts and schists of the Patya¯r Unit represent the former lagoonal deposits and the successive pelagic sedimentary cover deposited after the drowning of the atoll. The same authors also question previous findings of Cambrian archaeocyathids (Sharkovski et al. 1984), although no alternative indication of the age of the reefal carbonates of the Anarak Seamount is given. According to Bagheri & Stampfli (2008), the Morgha¯b schists and the blueschists of Cha¯h Gorbeh unit form an accretionary wedge located along the southern active margin of the Eurasian Plate (Turan domain). The Anarak Seamount (La¯k Marble and Patya¯r unit) and the nearby Kabudan Guyot were partly subducted, metamorphosed and accreted into the wedge during either Permian or Triassic time. New radiometric Ar –Ar ages (Bagheri & Stampfli 2008) obtained on white micas constrain the presence of an older ‘Variscan’ sensu latu metamorphism in the Morgha¯b Schists unit (333 –320 Ma), also exposed east of Jandaq (Fig. 2), and of a younger event poorly dated between 280 and 232 Ma in the Cha¯h Gorbeh

273

unit, respectively from Ar–Ar dating of crossite and stilpnomelane. A new element introduced by Bagheri & Stampfli (2008) concerns the age and significance of the Dosha¯kh and the Bayazeh units, previously interpreted as a greenschists-facies Precambrian or Lower Palaeozoic metamorphic basement of the Yazd block (Sharkovski et al. 1984). According to Bagheri & Stampfli (2008), the Dosha¯kh unit has given a Middle Permian–Middle Triassic age based on conodonts, and it is now interpreted as a part of the accretionary wedge developed after the collision of the Anarak seamount (Bagheri & Stampfli 2008). The Bayazeh Flysch forms a distal turbidite unit related to the same PermoTriassic wedge, recording the final closure of the subduction-related basin.

New geological observations in the Anarak area The Anarak Massif was investigated along the main road that crosses the high mountain chain north of Anarak, reaching the locality of Cha¯h Gorbeh (Figs 3, 8 and 9A), in order to define the main structural features of the Anarak Metamorphic Complex (AMC) and its relationships with the MesoCenozoic units cropping out to the north. The Cha¯h Gorbeh unit consists here of an association of marbles, metapelites and metabasites which are

Fig. 8. Geological cross section at Anarak; trace of the section in Figure 3. The topography is poorly constrained; vertical and horizontal scales are the same ones. F1, older faults possibly related to the growth of the Eo-Cimmerian accretionary wedge; F2, late Tertiary faults. Stereoplots as in Figure 4. Convergent arrows correspond to the horizontal direction of the s1 axis; divergent arrows to the horizontal direction of the s3 axis. Stress axes are represented as stars with five (s1), four (s2) and three branches (s3) obtained with inversion of fault populations using the inversion methods proposed by Angelier (1984, 1990). The small sections in the rectangle is a lateral variation of the main one observed south of Kuh-e Cha¯h Gorbeh.

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A. ZANCHI ET AL.

Fig. 9. Field views of the Anarak Metamorphic Complex. (A) Northward view across the Anarak Massif along the main road; compare with the section of Figure 8; (B) thrust contact between the serpentinitic slivers occurring in the AMC and the Oligocene sedimentary deposits; (C) blueschists of the Cha¯h Gorbeh unit, possibly showing relicts of pillow lava structures; (D) blueschist boudins in a greenschist matrix, Cha¯h Gorbeh unit; and (E) pseudomorph aggregates on plagioclase phenocrystals, Cha¯h Gorbeh unit.

juxtaposed to strongly serpentinized ultramafic rocks by a steep north-dipping fault (Figs 8 and 9B). Both units overthrust a folded terrigenous succession mainly consisting of Eocene to (poorly constrained) Lower Miocene beds, including red marls and sandstones with discontinuous conglomeratic layers. Faults developed along the contact suggest a NE– SW direction of horizontal compression, consistent with the trends of fold axes developed within the Tertiary sediments. The northern

boundary of the Tertiary outcrops is also faulted, showing reverse faults with the same trends and conjugate strike-slip faults overprinting the contact with a well-preserved granitoid intrusive. A different structure was observed along another cross-section located to the east of the main road, where the Cha¯h Gorbeh unit is exposed. Along this section, a thin slice of blueschists is sandwiched between two serpentinitic layers that show northdipping tectonic boundaries. Blueschists occur in

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

metabasites, possibly representing a succession of meta-lava flows (Fig. 9C, D).

Petrography of the Anarak Metamorphic Complex We focused on two key rock associations that can give important insights into the evolution of the AMC: blueschists of the Cha¯h Gorbeh unit and ultramafic rocks. Blueschists and greenschists of the Cha¯h Gorbeh unit. Blueschists occur in the Cha¯h Gorbeh unit, forming small decimetric boudins possibly representing deformed pillow lavas; boudins are embedded in a foliated matrix consisting of epidote–amphibole– chlorite schists (Fig. 9C, D). Blueschists are very fine grained (crystal size ,2 mm) and display a S-tectonic texture with a pervasive foliation marked by the preferred shape orientation of amphibole and chlorite. They contain amphibole, chlorite, plagioclase, white mica, quartz, titanite, rutile, apatite with subordinate calcite, hematite and sulphides. A pervasive foliation (S1) is defined by blue-lavender amphibole and chlorite; white mica is aligned parallel to the S1 only in a few samples, whereas it commonly forms spherulitic aggregates bordered by a thin titanite corona (Fig. 10A). White mica is oriented in both random and radial directions within the spherulitic aggregates (Fig. 10B). Deflection of the S1 foliation around the aggregates indicates that they predate the foliation. A pervasive post-S1 crystallization of randomly oriented blue–green amphibole and chlorite occurs in several samples, probably indicating a transition from blueschist- to greenschist-facies conditions.

275

The equilibrium phase assemblage for the blueschist pressure –temperature (P –T ) conditions includes amI þ chlI þ wmcaII þ qtz þ ttn þ plI þ hem (all mineral abbreviations after IUGS recommendations). AmI has a composition varying from Mg-riebeckite to glaucophane, depending on the Al/Fe3þ ratio (Table 1). A general trend from Mg-riebeckitic toward glaucophanic composition from core to rim (i.e. increase in the Al/ Fe3þ ratio) was observed within larger amphibole crystals (Fig. 11A, B). The K content is close to zero, whereas low amounts of Ti and Mn have been detected (0.02 atoms per formula unit (apfu)) displaying no correlation with Al/Fe3þ and XMg (Table 2). AmII, growing statically with no preferred orientation together with chlII, has an actinolitic composition. White mica, wmcaII, aligned parallel to the S1 foliation has a phengitic composition with Si comprised between 3.5 and 3.6 apfu and XMg close to 0.68 apfu (Fig. 11C and Table 2). Significant amounts of Ti (0.04 apfu) and Cr (0.02 apfu) have been detected in several crystals. White mica, wmcaI, forming the pre-S1 spherulitic aggregates, have Si contents around 3.4 apfu, slightly lower than wmcaII. Plagioclase in textural equilibrium with the syn-S1 phase assemblage displays an albitic composition (Ab98) close to the end-member. Thin rims of Ca-rich plagioclase (Ab95An05), plII, where observed to grow on plgI where the static crystallization of post-S1 amII and chlII is more intense. Chlorite shows no significant compositional variations as both chlI, grown synkinematically to S1, and chlII, grown statically with amII, have similar compositions to XMg close to 0.70 apfu. As previously described, titanite forms thin coronae, no more than 0.2–0.3 mm in width,

Fig. 10. Photomicrographs of blueschists from the Cha¯h Gorbeh unit of the Anarak Metamorphic Complex. Am, blue-amphibole; ch, chlorite; ttn, titanite; wmca, white mica. I, II, II refer to the different phase assemblages; see text for discussion. (A) Spherulithic aggregates of WmcaI rimmed by titanite coronae within a fine-grained matrix made by acicular bluish amphibole, chlorite and plagioclase; (B) WmcaI aggregates embedded within a fine-grained matrix in which the S1 foliation is visible in the right part of the image.

276

Table 1. Mineral chemistry of the amphiboles in the blueschists of the Anarak Metamorphic Complex, Cha¯h Gorbeh Unit. Location of samples IC33 and IC34 given in Figure 3 IC33 ampI

IC33 ampI

IC33 ampI

IC33 ampI

IC34 ampI

IC34 ampI

IC34 ampI

IC34 ampI

IC33 ampII

IC33 ampII

IC33 ampII

IC34 ampII

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO NiO MgO CaO Na2O K2O

55.86 0.07 2.7 0 11.73 7.16 0.13 0.01 11.4 2.55 5.96 0.04

57.45 0.04 3.24 0.05 11.23 7.67 0.15 0.04 10.93 1.06 6.8 0.02

57.27 0.06 3.21 0.03 11 7.75 0.11 0.01 11.21 1.02 6.67 0.03

57.53 0.01 3.77 0.06 10.22 8.26 0.12 0.04 10.74 0.77 6.85 0.01

56.41 0.11 2.63 0.08 12.62 7.96 0.1 0.05 10.34 0.88 6.8 0.04

56.75 0.15 3.07 0.01 11.88 8.24 0.12 0.01 10.02 1.21 6.86 0.03

56.36 0.11 2.46 0.04 13.49 7.54 0.07 0 10.42 1.01 6.82 0.02

56.15 0.14 2.66 0.04 14.28 7.1 0.11 0.05 10.28 0.88 6.93 0.01

51.34 0.05 2.02 0.01 11.41 2.45 0 0 16.12 10.46 1.87 0.03

53.49 0.04 2.06 0.04 7.19 5.77 0.01 0.01 15.73 10.46 1.83 0.02

52.72 0.02 2.15 0.01 8.21 5.7 0.01 0.01 15.75 10.56 1.61 0.03

50.95 0.01 1.96 0.01 10.77 2.99 0.02 0.01 16.03 10.78 1.6 0.03

Total

97.61

98.68

98.37

98.38

98.02

98.35

98.34

98.63

95.76

96.65

96.78

95.16

Si Ti Al Cr Fe3þ Fe2þ Mn Ni Mg Ca Na K Cation XMg NaB

7.957 0.007 0.453 0.000 1.257 0.853 0.016 0.001 2.420 0.389 1.646 0.007

8.056 0.004 0.535 0.005 1.185 0.899 0.017 0.005 2.284 0.160 1.849 0.003

8.050 0.006 0.532 0.003 1.164 0.911 0.013 0.001 2.349 0.153 1.818 0.006

8.075 0.001 0.624 0.006 1.079 0.969 0.015 0.004 2.247 0.116 1.864 0.003

8.022 0.012 0.441 0.009 1.351 0.947 0.012 0.006 2.192 0.135 1.875 0.007

8.034 0.016 0.512 0.001 1.266 0.976 0.015 0.001 2.114 0.183 1.883 0.005

7.996 0.011 0.411 0.005 1.441 0.895 0.008 0.000 2.204 0.154 1.876 0.005

7.951 0.015 0.444 0.004 1.521 0.841 0.013 0.005 2.170 0.133 1.903 0.002

7.46 0.01 0.35 0.00 1.25 0.30 0.00 0.00 3.49 1.63 0.53 0.01

7.68 0.00 0.35 0.00 0.78 0.69 0.00 0.00 3.37 1.61 0.51 0.00

7.59 0.00 0.37 0.00 0.89 0.69 0.00 0.00 3.38 1.63 0.45 0.01

7.46 0.00 0.34 0.00 1.19 0.37 0.00 0.00 3.50 1.69 0.45 0.01

15.007 0.534 1.611

15.003 0.523 1.840

15.006 0.531 1.818

15.003 0.523 1.864

15.007 0.488 1.865

15.005 0.485 1.817

15.005 0.486 1.846

15.002 0.479 1.867

15.006 0.693 0.372

15.004 0.696 0.39

15.006 0.682 0.371

15.006 0.693 0.309

A. ZANCHI ET AL.

Sample Phase

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

277

Fig. 11. Compositional variations of (A) and (B) amphibole, (C) white micas and (D) titanites from the blueschists of the Cha¯h Gorbeh unit. AmI are substantially Mg-riebeckits and display a variation trend towards glaucophanic compositions from core to rim. Micas parallel to the S1 foliation (wmcaII) have a higher Si content with respect to wmcaI. Titanites show a weak compositional zoning with Ti/(Al þ Fe3þ) ratio increasing towards the outer rim.

around wmcaI aggregates. A loosely defined compositional zoning of ttn coronae going from the inner rim, in contact with wmcaI, to the outer rim, in contact with the syn-S1 phases, was observed (Fig. 11D); zoning is individuated by an increase of Ti (from 0.9 to 1 apfu) and decrease of Al þ Fe3þ from 0.12 to 0.02 apfu. Estimate of the peak metamorphic conditions for the blueschists of the Cha¯h Gorbeh unit is not straightforward, as the syn-S1 phase assemblage, in the absence of both epidote and lawsonite, is loosely constrained in the P– T field. An upper pressure limit is posed by the absence of omphacitic clinopyroxene, which constrains the peak pressure for the syn-S1 assemblage below 1.2 GPa. In addition, a lack of aragonite in the S1 paragenesis, indicating that calcite was the stable carbonate phase during peak conditions, constrains the upper pressure boundary to 0.8– 0.9 GPa. An upper temperature limit is also provided by the absence of garnet (Apted & Liou 1983; Guiraud et al. 1990), which is stable above 400–450 8C for pressure conditions ranging between 0.4 and 0.9 GPa.

The absence of lawsonite could be tentatively used as a lower temperature boundary, confining the peak phase mineral assemblage with Na-amphibole, albite, white mica, chlorite and titanite at a temperature above 300 8C for a 0.4 –0.9 GPa pressure range. The pressure peak recorded by the syn-S1 phase assemblage was followed by a decrease in pressure and an increase in temperature, suggested by the crystallization of actinolitic amphibole and chlorite and the growing of thin rims of Ca-rich plgII on plI. This suggests that the blueschists of the Cha¯h Gorbeh unit were subducted into an accretionary wedge at a depth of at least 15–20 km. Serpentinized peridotites. Ultramafic rocks have been sampled along the main road from Anarak to the north. Here they are almost completely serpentinized, preserving only sparse relicts of primary phase assemblages (Fig. 12A). Serpentinites display a schistose texture, with isolated domains showing a preserved massive texture. Clinopyroxene, olivine and spinel relicts have been found in these domains,

278

Table 2. Mineral chemistry of white micas, chlorite, titanite and plagioclase within the blueschists of the Anarak Metamorphic Complex, Cha¯h Gorbeh Unit. Location of samples IC33 and IC34 is given Figure 3 IC33 wmII

IC33 wmII

IC33 wmI

IC33 wmII

IC34 wmII

IC33 chlI

IC33 chlI

IC33 ttn

IC34 ttn

IC34 ttn

IC33 plgI

IC33 plgI

IC34 plgI

IC34 plgI

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO NiO MgO CaO Na2O K2O

51.35 0.29 22.44 0.15 0 3.98 0.04 0 4.86 0 0.06 10.99

51.59 0.27 21.64 0.23 0 4.91 0.06 0 5.84 0 0.04 10.27

48.4 0.8 18.63 0.28 0 8.33 0.07 0 8.53 0 0.06 10.02

52.01 0.39 20.62 0.21 0 5.13 0.04 0 6.17 0 0.05 10.53

52.62 0.32 21.99 0.15 0 4.05 0.05 0 4.85 0 0.05 10.81

29.79 0.03 18.13 0.11 0 16.72 0.32 0.03 22.39 0.12 0.06 0.1

30.15 0.04 17.99 0.14 0 17.1 0.29 0.07 22.57 0.13 0.07 0.09

31.68 37.02 1.43 0 0.24 0.44 0.03 0 0.05 27.88 0.15 0.05

31.85 37.37 1.2 0.01 0.16 0.62 0.01 0 0.05 27.79 0.16 0.05

31.54 38.17 0.5 0.05 0.33 0.35 0.05 0 0.03 27.98 0.07 0.05

69.46 0.00 19.48 0.00 0.00 0.16 0.02 0.03 0.02 0.00 11.56 0.04

68.94 0.03 19.29 0.00 0.00 0.41 0.00 0.00 0.04 0.04 11.66 0.04

69.56 0.02 19.38 0.01 0 0.5 0 0 0.08 0.06 11.78 0.03

69.47 0.02 19.51 0.04 0 0.4 0.01 0 0.03 0.05 11.73 0.02

Total

94.16

94.85

95.12

95.15

94.89

87.8

88.64

98.97

99.27

99.12

100.78

100.45

101.42

101.28

Si Ti Al Cr Fe3þ Fe2þ Mn Ni Mg Ca Na K

3.516 0.015 1.811 0.008 0.000 0.228 0.002 0.000 0.496 0.000 0.008 0.960

3.512 0.014 1.736 0.012 0.000 0.280 0.004 0.000 0.593 0.000 0.005 0.892

3.381 0.042 1.534 0.016 0.000 0.487 0.004 0.000 0.888 0.000 0.008 0.893

3.540 0.020 1.654 0.011 0.000 0.292 0.002 0.000 0.626 0.000 0.007 0.914

3.565 0.016 1.756 0.008 0.000 0.229 0.003 0.000 0.490 0.000 0.007 0.934

6.001 0.005 4.304 0.018 0.000 2.817 0.055 0.005 6.722 0.026 0.023 0.026

6.022 0.006 4.235 0.022 0.000 2.857 0.049 0.011 6.720 0.028 0.027 0.023

1.032 0.907 0.055 0.000 0.006 0.012 0.001 0.000 0.002 0.973 0.010 0.002

1.036 0.914 0.046 0.000 0.004 0.017 0.000 0.000 0.002 0.968 0.010 0.002

1.032 0.939 0.019 0.001 0.008 0.010 0.001 0.000 0.002 0.981 0.004 0.002

3.018 0.000 0.998 0.000 0.000 0.006 0.001 0.001 0.001 0.000 0.974 0.002

3.003 0.001 0.990 0.000 0.000 0.015 0.000 0.000 0.003 0.002 0.985 0.002

3.001 0.001 0.985 0.000 0.000 0.018 0.000 0.000 0.005 0.003 0.985 0.002

3.002 0.001 0.994 0.001 0.000 0.015 0.000 0.000 0.002 0.002 0.983 0.001

A. ZANCHI ET AL.

Sample Mineral

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

279

Fig. 12. (A) Photomicrograph of serpentinized peridotite from the Cha¯h Gorbeh unit showing bastite pseudomorphs after pyroxenes; (B) bastites within a matrix displaying mesh to interpenetrating textures; (C) back scattered electrons/scanning electron microscope (BSE/SEM) image of an ameboid spinel porphyroclast and clinopyroxene relicts within a tremolite þ antigorite matrix; and (D) BSE/SEM image of complex chemical zoning within a spinel porphyroclast. ant, antigorite; cpx, clinopyroxene; sp, spinel; trm, tremolite.

whilst only chromian spinel porphyroclasts are present in the schistose ones. Schistosity is defined by the preferred orientation of antigorite, replacing bastites of lizardite and chrysotile on primary olivine and pyroxene. Mesh and interpenetrating textures are common in several samples (Fig. 12B). Preserved spinel grains display an amoeboid habit with complex compositional zoning (Fig. 12C, D). In many cores an Al-rich spinel is preserved (sp I, dark grey) with Cr# (Cr# ¼ molar[100*Cr/(Cr þ Al)]) 0.14–0.16, Fe3þ/(Al þ Cr þ Fe3þ) , 0.1 and Ti content close to zero, representing, possibly, a primary phase (Dick & Bullen 1984). Successive re-equilibrations are indicated by overgrowth rims progressively richer in Fe and depleted in Cr (Table 3). The surrounding darkgrey matrix mainly consists of antigorite and tremolite with minor amounts of chlorite. Rare olivine relicts preserved at the core of pseudomorphs with mesh texture have a Fo94 composition with small amounts of Cr (0.04 Cr2O3 wt%), Mn (0.20 MnO wt%) and Ni (0.11 NiO wt%).

Clinopyroxene relicts, observed in one sample, display irregular shapes, varying with the degree of substitution by tremolite and antigorite. Their chemical composition is close to the diopside endmember, with amount of Na and Al close to zero (Table 4). Tiny intergrowths of tremolite and a sodic–calcic amphiboles are observed around some clinopyroxene relicts. The presence of such reaction rims may indicate an increase in pressure, consistent with the metamorphic evolution suggested for the blueschists of the Cha¯h Gorbeh unit.

Discussion Petrographic and geological data on the Nakhlak – Anarak region provide several constraints for the reconstruction of the Triassic geodynamic evolution of Central Iran. They suggest that the Nakhlak succession was deposited in a forearc setting to the north of a subduction zone. Intermittent stages of volcanic activity alternating with uplift and erosion may indicate oblique convergence and

280

A. ZANCHI ET AL.

Table 3. Mineral chemistry of spinels contained in the ultramafics of the Anarak Metamorphic Complex. Location of samples CI5 and CI6 is given in Figure 3 Sample Phase

CI5 sp1

CI5 sp1

CI6 sp1

CI6 sp2

CI6 sp2

CI5 sp3

CI6 sp3

SiO2 TiO2 Al2O3 Cr2O3 Fe2O3 FeO MnO NiO ZnO MgO CaO Na2O K2O

0.03 0.04 52.87 15 0.06 11.16 0.1 0.27 0 18.48 0 0.01 0

0.04 0.03 51.88 15.32 0.2 11.47 0.13 0.29 0 17.98 0 0.02 0

0.05 0.04 52.76 14.89 0.67 10.73 0.12 0.32 0 18.79 0 0.01 0

0.21 0.09 0.13 7.13 60.63 29.36 1.06 0.12 0 0.45 0.04 0 0

1.63 0.53 0.2 9.25 55.36 28.07 1.77 0.28 0 1.99 0.06 0.09 0

0.66 0.14 0.09 3.79 64.2 29.55 0.58 0.1 0 0.99 0.04 0.09 0

0.83 0.13 0.13 4.5 62.99 29.37 0.67 0.16 0 1.05 0.19 0.08 0

Total

98.03

98.03

98.03

99.22

99.23

100.22

100.1

Si Ti Al Cr Fe3þ Fe2þ Mn Ni Zn Mg Ca Na

0.001 0.001 1.677 0.319 0.001 0.251 0.002 0.006 0.000 0.741 0.000 0.001

0.001 0.001 1.664 0.330 0.004 0.261 0.003 0.006 0.000 0.729 0.000 0.001

0.001 0.001 1.667 0.316 0.014 0.241 0.003 0.007 0.000 0.751 0.000 0.001

0.008 0.003 0.006 0.217 1.756 0.945 0.035 0.004 0.000 0.026 0.002 0.000

0.061 0.015 0.009 0.276 1.569 0.884 0.057 0.009 0.000 0.112 0.002 0.007

0.025 0.004 0.004 0.114 1.831 0.937 0.019 0.003 0.000 0.056 0.002 0.007

0.031 0.004 0.006 0.135 1.795 0.930 0.022 0.005 0.000 0.059 0.008 0.006

K Fe3þ/(Al þ Cr þ Fe3þ) Cr#

0.000 0.001 0.160

0.000 0.002 0.165

0.000 0.007 0.159

0.000 0.887 0.974

0.000 0.847 0.969

0.000 0.940 0.966

0.000 0.927 0.959

strike-slip tectonics in the arc massif. The interpretation of the major erosional event documented by the Ba¯qoroq Formation is fundamental for the reconstruction of the evolution of Central Iran. Based on present distribution of rock outcrops in Central Iran, we note that lithologies of the metamorphic rocks exposed south of Nakhlak are compatible with features of metamorphic detritus supplied to the Ba¯qoroq Formation. If the source of metamorphic detritus in the Ba¯qoroq Formation is, indeed, the Anarak complex, it may be envisaged as: (a) a young metamorphic complex newly produced by Early –Middle Triassic collision tectonics; or (b) an older metamorphic basement involved in a new orogenic cycle. The new geochronological ages and reconstruction obtained by Bagheri & Stampfli (2008) indicate that a ‘Variscan’ event is recorded between the Great Kavir and the Biababak faults, and that a Permo-Triassic accretionary wedge was active in the Anarak region to the south of Nakhlak.

The occurrence of high-pressure mineralogical assemblages in the Cha¯h Gorbeh unit, already described in the literature (Sharkovski et al. 1984) but successively ignored by several authors (Davoudzadeh & Weber-Diefenbach 1987; Soffel et al. 1996; Alavi et al. 1997), is demonstrated by our analysis, constraining the P –T evolution of the Cha¯h Gorbeh unit to a subduction-related margin (Fig. 13). A minimum age for the highpressure metamorphism is given by Ar –Ar stilpnomelane (233 + 2 Ma), post-dating the event (Bagheri & Stampfli 2008), although younger ages have been published. The exposure and unroofing of the accretionary wedge is recorded by the thick alluvial fans of the Ladinian Ba¯qoroq Formation and by the Dosha¯kh and Bayazeh Flysch (Bagheri & Stampfli 2008), which also contain upper Ladinian –middle Carnian conodont faunas. The Ba¯qoroq Formation is rich in clasts of marble and metapelitic rocks, very similar to the

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

281

Table 4. Mineral chemistry of clinopyroxenes contained in the ultramafics of the Anarak Metamorphic Complex. Location of samples CI5 and CI6 is given in Figure 3 Sample Phase

CI5 cpx

CI5 cpx

CI5 cpx

CI6 cpx

CI6 cpx

CI5 ol

CI5 ol

CI5 ol

SiO2 TiO2 Al2O3 Cr2O3 FeO* MnO NiO MgO CaO Na2O K2O

54.45 0 0.02 0.07 1.13 0.14 0.01 17.39 25.58 0.05 0.01

54.11 0.01 0.03 0.05 0.99 0.13 0 17.24 25.63 0.02 0.01

54.85 0.01 0.03 0 0.85 0.09 0.05 17.39 25.81 0.03 0

55.05 0.05 0.04 0.01 1.01 0.08 0.05 17.51 25.68 0.04 0.01

54.74 0.01 0.13 0.01 1.02 0.07 0.02 17.63 25.67 0.05 0

41.12 0 0 0.04 6.14 0.19 0.09 52.57 0.05 0 0

41.34 0 0.01 0.03 6.09 0.23 0.07 52.51 0.04 0.01 0

41.09 0 0 0.05 6.32 0.21 0.11 51.88 0.04 0.01 0.01

Total

98.89

98.22

99.12

99.49

99.18

100.33

99.72

100.2

Si Ti Al Cr Fe3þ Fe2þ Mn Ni Mg Ca Na K

1.998 0.000 0.001 0.002 0.001 0.035 0.004 0.000 0.951 1.006 0.004 0.001

1.998 0.000 0.001 0.002 0.000 0.031 0.004 0.000 0.949 1.014 0.001 0.001

2.004 0.000 0.001 0.000 0.001 0.025 0.003 0.002 0.947 1.010 0.002 0.000

2.003 0.001 0.002 0.000 0.004 0.025 0.003 0.002 0.950 1.001 0.003 0.001

1.998 0.000 0.006 0.000 0.000 0.026 0.002 0.001 0.959 1.004 0.004 0.000

0.987 0.000 0.000 0.001 0.000 0.123 0.004 0.002 1.882 0.001 0.000 0.000

0.992 0.000 0.000 0.001 0.000 0.122 0.005 0.001 1.878 0.001 0.001 0.000

0.993 0.000 0.000 0.001 0.000 0.128 0.004 0.002 1.870 0.001 0.001 0.000

XMg Cation

0.965 4.002

0.969 4.001

0.975 3.995

0.974 3.994

0.974 4.000

0.935 3.000

0.935 3.000

0.932 3.000

metasedimentary units of the Morgha¯b schists and Palha¯vand gneiss dated as Carboniferous. Uplift and erosion of the prism may be explained by the introduction of rigid bodies into the trench, as suggested by Bagheri & Stampfli (2008), who interpret a large part of the AMC as partially subducted seamounts. The youngest Triassic unit of the Nakhlak succession, the Ashin Formation, is rich in felsic volcaniclastic turbiditic sandstones indicating a reprisal of the arc activity in the Late Ladinian– ?Carnian. This might be ascribed to a transtensional stage of renewed subsidence and volcanism in an arc –trench setting associated with oblique convergence. The persistence of distinct and unmixed metamorphic and volcanic provenances during deposition of the Ashin Formation is thus explained by renewed volcanic activity in the north and continuing erosion of the accretionary prism in the south. These data point to the occurrence of a collision event between an active margin and the Yazd block, although its precise age has not yet been defined so far. Conversely, the Late Carboniferous radiometric

dates obtained in the Jandaq and Anarak region by Bagheri & Stampfli (2008) pose several questions as to their interpretation: no clear relationships between deformational structures, mineral assemblages and their pressure–temperature –time (P– T –t) evolution has been established by the authors. This implies that their radiometric data may be mixed ages not directly relatable to single metamorphic events. In addition, the occurrence of sillimanite –K-feldspar gneisses of the Jandaq belt, together with blueschists within the Variscan accretionary wedge complex, is fairly uncommon and needs more investigation, as these peculiar high-temperature metamorphic rocks can be more easily found in different geodynamic settings. Late Palaeozoic metamorphic rocks dated to the same time interval in the Talesh Mountains (Zanchetta et al. 2009) and in Afghanistan (Boulin 1988, 1991) generally show, in fact, a highpressure imprint of blueschist–eclogitic conditions. The occurrence of the Variscan – Eo-Cimmerian wedge within central Iran has been explained by the displacement of a large fragment of the Palaeotethys suture once located between Mashhad and

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Fig. 13. Tectonic scheme of Iran with the main tectonic subdivisions, modified from the Geological Survey of Iran (1989). The location of the Eo-Cimmerian and Variscan units is highlighted. Green, Mesozoic ophiolites along the Main Zagros Thrust and other sutures. AA, Araxian–Azarbaijanian zone; Ab, Abadeh; Ag, Aghdarband Basin; AMC, Anarak Metamorphic Complex; Ba, Baft; Es, Esfahan; GKDF, Great Kavir–Doruneh fault system; Gol, Golpayegan; GS, Gorgan Schist; Ka, Kashmar; Kd, Ko-e-Dom; KDF, Kopeh Dagh Foredeep; KF, Kalmard Fault; Kn, Kashan; Ks, Kor-e-Sefid metamorphic complex; Hj, Hajiabad Metamorphic Complex; Ma, Mashhad; MZT, Main Zagros Thrust; Nk, Nakhlak Basin; Nn, Nain; Ny, Neyriz ophiolites; PB, Posht-e Badam; PTSZ, Palaeotethys Suture Zone; Sa, Sabzevar; SG, Shanderman and Gasht metamorphic complexes; Si, Sirjan; SSZ, Sanandaj – Sirjan Zone; Ta, Tabas; To, Torud; Yz, Yazd. Radiometric ages from the literature are reported for Variscan and Cimmerian metamorphosed units; see the text for references. WR, whole rock.

Afghanistan along the southern margin of the Turan domain (Bagheri & Stampfli 2008). According to their interpretation, the Late Cretaceous opening of the Nain–Sabzevar Ocean in a back-arc position was a consequence of the Neotethys subduction below the Sanandaj–Sirjan Zone. It separated parts of the accretionary complex from the Turan domain, and these portions were successively transported within Central Iran during the closure of the basin.

That interpretation follows from the idea, already proposed by several authors, that large counter-clockwise rotation (1358) of the whole Central Iran Microplate accompanied these processes (Wensink 1979; Davoudzadeh et al. 1981; Davoudzadeh & Weber-Diefenbach 1987; Soffel et al. 1996). However, several inconsistencies derive from the mechanisms of emplacement of the Jandak–Anarak terranes and especially the

THE CIMMERIAN EVOLUTION OF CENTRAL IRAN

supposed counter-clockwise rotations of the whole Central Iran region south of the Great Kavir– Doruneh fault system. New palaeomagnetic results and several geological data collected so far strongly question the amount of rotation postulated by most authors. Palaeomagnetic data (Muttoni et al. 2009) advise caution with regard to the hypothesis of Soffel et al. (1996) that a 1358 counter-clockwise block rotation has occurred between the internal part of Central Iran and Europe since the Middle Triassic, an hypothesis that is based on sparse data from the Triassic sequence of Nakhlak. Multiple remagnetization events pervasively overprinted the volcaniclastic sandstones and siltstones at Nakhlak, whereas magnetization components of presumed primary age were observed only in a single nodular limestone interval. These primary components, when compared with reference European directions, suggest an Eurasian affinity, and no significant vertical axis rotations of the Nakhlak structure since the late Early Triassic (site IR04 in Muttoni et al. 2009). Undoubtedly, more data are required to unravel in detail the complex tectonic history of Iran, but we stress that our new palaeomagnetic data do not confirm a 1358 counter-clockwise rotation of the internal part of Central Iran with respect to Eurasia since the Middle Triassic (Soffel et al. 1996). Similar conclusions were reached also by Besse et al. (1998). A further geological argument against a large counter-clockwise vertical axis rotation of Central Iran is given by Wendt et al. (2005), who conclude that regional facies analysis of Upper Silurian– Lower Carboniferous successions is not consistent with a large 1358 anticlockwise rotation of central Iran. Similarly, the present distribution of the units within the Anarak area indicates that a considerable rotation of the interior of Central Iran never occurred. In fact, the orientation of the three main Eo-Cimmerian elements in the area: the arc region of Nakhlak, the accretionary wedge, and the foreland area of the Yazd block, is consistent with their original palaeogeography in an arc –trench system. The northern boundary of the Variscan units of Central Iran also poses several questions worthy of discussion. Bagheri & Stampfli (2008) recognized that the Variscan wedge is juxtaposed against the Airekan terrane south of the Doruneh fault south of Sabzevar, and that it consists of deformed peraluminous Cambrian granitic rocks with a U –Pb zircon age of 549 + 15 Ma. These granitic rocks match the radiometric ages of other granitic intrusions forming the Gondwanan basement of Central Iran, such as near Posht-e-Badam (Ramezani & Tucker 2003) and the gneissic granites (Hassanzadeh et al. 2008) of the Shotori Kuh, 50 km

283

NE of Torud on the north side of the Great Kavir Fault. In spite of this strong affinity, Bagheri & Stampfli (2008) hypothesize that the Airekan terrane is part of the Hun block, which detached during the Lower Palaeozoic from Gondwana and accreted to Eurasia before Variscan times. This interpretation is inconsistent with available information (Garzanti & Gaetani 2002; Natal’in & Sengo¨r 2005 and references therein), as the so-called ‘Turan domain’, from which the Airekan terrane may derive, includes high-grade metamorphic rocks with a strong Late Palaeozoic metamorphic imprint. Several authors have noted a strong affinity between the arc-related successions of the Aghdarband Triassic basin (Ruttner 1991, 1993) on the southern margin of Eurasia, east of Mashhad, and those of Nakhlak (Alavi et al. 1997). This supposed correlation has been used as the most persuasive argument to state that these localities were part of the same Palaeotethyan convergent margin, and that the Nakhlak succession was originally close to Aghdarband in the same trench–slope interval (Alavi et al. 1997). Balini et al. (2009) question the possible proximity of the two localities based on comparison among ammonoid faunas. Olenekian ammonoids from Aghdarband show a Periscaspian afinity, whereas the ones from Nakhlak show a typical Tethyan affinity: such an important palaeobiogeographic difference is not consistent with a proximity of the two areas.

Conclusions We interpret the Triassic history of the Nakhlak – Anarak region as the result of the evolution of an arc –trench system related to Eo-Cimmerian orogenic events. The Triassic volcaniclastic successions of Nakhlak records the evolution of a volcanic arc facing an accretionary wedge, the remnants of which can be identified in the blueschists, ultramafic rocks and metacarbonates of the Anarak Metamorphic Complex (AMC). Deformation of the Nakhlak succession is presumably related to the collision of the arc– trench system with the Yazd block possibly during Late Triassic times, or to the subsequent Neo-Cimmerian event that strongly affected Central Iran between the end of the Jurassic and the beginning of the Cretaceous. Intensive compressional phenomena affected the study area during the Tertiary, evidenced in folding and thrusting of the Eocene –Oligocene terrigenous successions in the Anarak region and in northvergent reverse faults dismembering the AMC. Considerable strike-slip movements occurred along a major N– S-trending dextral strike-slip fault that displaced the eastern portion of the Nakhlak area.

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This structure is part of the fault system that has affected Central Iran since the late Tertiary (Walker & Jackson 2004), causing differential block rotations along vertical axes. New palaeomagnetic data obtained in the Triassic succession of Nakhlak indicate that there is no clear evidence of significant counter-clockwise rotations, and that palaeomagnetic information should be interpreted with extreme caution because of recent overprint affecting the succession. In any such case the latitude implied by the palaeomagnetic data suggests an Eurasian position for Nakhlak at the beginning of the Triassic. Several contradictory and as-yet unresolved aspects of the Late Palaeozoic–Triassic evolution of Iran have been discussed in this paper, based on our field data and the analysis of the literature. Although several lines of evidence may suggest that large displacements of entire crustal blocks occurred during the amalgamation of Iran, the exact times, mechanism and the supposed amount of displacements of these terranes are still obscure. The interpretation of the Palaeozoic– Triassic evolution of Iran is also related to the understanding of the significance of the Variscan and Eo-Cimmerian metamorphism elsewhere in Iran, presently at large distances from the southern part of the Sanandaj– Sirjan Zone and its coeval metamorphic belts of Central and North Iran. Unravelling the palaeogeographical and structural relationships among these different crustal blocks is a difficult challenge for future geological research, and possibly represent a Mesozoic equivalent of the present-day ‘Indonesian-type’ orogenic scenario. This work was funded by the MEBE programme (Proposal 3.30: ‘Tectonic evolution of the Yazd, Tabas and Lut blocks (Central Iran) by means of palaeomagnetic, structural and stratigraphic data’; leader M. Mattei) in close collaboration with the Geological Survey of Iran. Dr M. R. Ghasemi, Dr A. Saidi and his colleagues of Geological Survey of Iran of Teheran are warmly thanked for continuous help and assistance in the field. The manuscript benefited from the reviews by A. Baud, M. Berberian and B. Natal’in. The three editors of the special volume, M.-F. Brunet, J. Granath and M. Wilmsen, are warmly thanked for their suggestions and accurate final revision of the manuscript.

Appendix Quantitative mineral analyses were obtained with a JEOL 8200 Superprobe equipped with WDS spectrometers at the University of Milano. Operating conditions were 15 kV and 15 nA. Natural silicates were used as standards and the resulting data were corrected through a ZAF correction procedure. Mineral formulae of amphibole from the blueschists of the Cha¯h Gorbeh unit were recalculated on the basis of

23 oxygens and 15 cation þ K. Plagioclase formulae were recalculated on the basis of 8 oxygens per formula unit. White micas formulae result from considering all Fe as Fe2þ on the basis of 11 O. Cations in titanites were obtained on the basis of 3 total cations and 10 charges minus OH groups. Chlorites formulae were calculated on the basis of 28 oxygens. Olivine analyses were calculated considering all Fe as Fe2þ and cations sum equal to 3. Spinel formulae are expressed considering 3 cations and 8 charges. Clinopyroxenes analyses were recalculated on the basis of 6 oxygens and considering Fe3þ ¼ Acmite.

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donne´es structurales et datations 40K– 40Ar sur les roches me´tamorphiques de la re´gion de Neyriz (zone de Sanandaj– Sirjan, Iran me´ridional). Leur inte´reˆt dans le cadre du domaine ne´o-te´thysien du Moyen-Orient. Comptes Rendus Geoscience, 335, 981–991. S OFFEL , H., D AVOUDZADEH , M., R OLF , C. & S CHMIDT , S. 1996. New paleomagnetic data from Central Iran and a Triassic paleoreconstruction. Geologische Rundschau, 85, 293– 302. S TAMPFLI , G. M. & B OREL , G. D. 2002. A plate tectonic model for the Palaeozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrones. Earth Planetary Sciences Letters, 196, 17–33. S TAMPFLI , G. M., M ARCOUX , J. & B AUD , A. 1991. Tethyan margins in space and time. Palaeogeography, Palaeoclimatology, Palaeoecology, 87, 373– 409. S TO¨ CKLIN , J. 1974. Possible ancient continental margins in Iran. In: B URK , C. A. & D RAKE , C. L. (eds) The Geology of Continental Margins. Springer, Berlin, 873–887. T ORSVIK , T. H. & C OCKS , R. M. 2004. Earth geography from 400 to 250 Ma: a palaeomagnetic, faunal and facies review. Journal of the Geological Society, London, 161, 555– 572. W ALKER , R. & J ACKSON , J. 2004. Active tectonics and late Cenozoic strain distribution in central and eastern Iran. Tectonics, 23, TC5010; doi:10.1029/ 2003TC001529. W ENDT , J., K AUFMANN , B., B ELKA , Z., F ARSAN , N. & B AVANDPUR , A. K. 2002. Devonian/Lower Carbonifeous stratigraphy, facies patterns and palaeogeography of Iran. Part I. Southeastern Iran. Acta Geologica Polonica, 52(2), 129 –168. W ENDT , J., K AUFMANN , B., B ELKA , Z., F ARSAN , N. & B AVANDPUR , A. K. 2005. Devonian/Lower Carbonifeous stratigraphy, facies patterns and palaeogeography of Iran Part II. Northern and Central Iran. Acta Geologica Polonica, 55(1), 31– 97. W ENSINK , H. 1979. The implications of some palaeomagnetic data from Iran from its structural history. Geologie en Mijnbow, 58, 175–185. W OODCOCK , N. H. & S CHUBERT , C. 1995. Continental strike-slip tectonics. In: H ANCOCK , P. L. (ed.) Continental Deformation. Pergamon, Oxford, 251–263. Z ANCHETTA , S., Z ANCHI , A., V ILLA , I., P OLI , S. & M UTTONI , G. 2009. The Shanderman eclogites: a Late Carboniferous high-pressive event in the NW Talesh Mountains (NW Iran). In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 57– 78. Z ANCHI , A., B ERRA , F., M ATTEI , M., G HASSEMI , M. & S ABOURI , J. 2006. Inversion tectonics in Central Alborz, Iran. Journal of Structural Geology, 28, 2023– 2037. Z ANCHI , A., Z ANCHETTA , S. ET AL . 2009. The Eo-Cimmerian (Late? Triassic) orogeny in North Iran. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 31–55.

The Triassic stratigraphic succession of Nakhlak (Central Iran), a record from an active margin MARCO BALINI1, ALDA NICORA1, FABRIZIO BERRA1, EDUARDO GARZANTI2, MARCO LEVERA1, MASSIMO MATTEI3, GIOVANNI MUTTONI1, ANDREA ZANCHI2, IRENE BOLLATI1, CRISTIANO LARGHI1, STEFANO ZANCHETTA1, REZA SALAMATI4 & FATHULLAH MOSSAVVARI4 1

Dipartimento di Scienze della Terra, Universita` di Milano, Via Mangiagalli 34, 20133 Milano, Italy (e-mail: [email protected]) 2

Dipartimento di Scienze Geologiche e Geotecnologie, Universita` di Milano-Bicocca, Piazza della Scienza 4, Milano, 20126, Italy 3

Dipartimento di Scienze Geologiche, Universita` Roma TRE, Largo San Leonardo Murialdo 1, 00146 Roma, Italy

4

Geological Survey of Iran, Azadi Square, Meraj Avenue, 13185-1494, Tehran, Iran Abstract: An important, 2.4 km-thick Triassic succession is exposed at Nakhlak (central Iran). This succession was deformed during the Cimmerian orogeny and truncated by an angular unconformity with undeformed Upper Cretaceous sediments. This integrated stratigraphic study of the Triassic included bed-by-bed sampling for ammonoids, conodonts and bivalves, as well as limestone and sandstone petrographic analyses. The Nakhlak Group succession consists of three formations: Alam (Olenekian –Anisian), Ba¯qoroq (?Upper Anisian– Ladinian) and Ashin (Upper Ladinian). The Alam Formation records several shifts from carbonate to siliciclastic deposition, the Ba¯qoroq Formation consists of continental conglomerates and the Ashin Formation documents the transition to deep-sea turbiditic sedimentation. Petrographic composition has been studied for sandstones and conglomerates. Provenance analysis for Alam and most of the Ashin samples suggests a volcanic arc setting, whereas the samples from the Ba¯qoroq Formation are related to exhumation of a metamorphic basement. The provenance data, together with the great thickness, the sudden change of facies, the abundance of volcaniclastic supply, the relatively common occurrence of tuffitic layers and the orogenic calc-alkaline affinity of the volcanism, point to sedimentation along an active margin in a forearc setting. A comparison between the Triassic of Nakhlak and the Triassic succession exposed in the erosional window of Aghdarband (Koppeh Dag, NE Iran) indicates that both were deposited along active margins. However, they do not show the same type of evolution. Nakhlak and Aghdarband have quite different ammonoid faunal affinities during the Early Triassic, but similar faunal composition from the Bithynian to Late Ladinian. These results argue against the location of Nakhlak close to Aghdarband.

Central Iran is geologically a very complex area, characterized by a tremendous variety of rock types ranging from Precambrian to Miocene sedimentary rocks, Palaeozoic–Cenozoic ultramafic–acid igneous rocks and Palaeozoic –Mesozoic metamorphic rocks. Such an astonishing geological diversity is related to a very long history that started with the assembly of Gondwana in early Palaeozoic times and continued to the present-day collision of the Arabian Plate with Eurasia. The unravelling of this long and complex history is challenging. The understanding of the Cimmerian system (Sengo¨r 1984) is especially difficult as the Carboniferous– Jurassic rocks were often deformed,

eroded, covered and/or metamorphosed during more recent collisional events, and their boundaries were sometimes reactivated by more recent faults and thrusts. Despite many published contributions to the geology of Central Iran, no shared interpretation has emerged at the microplate and local scales. For instance, there is no consensus on the proposal by Davoudzadeh et al. (1981) and Soffel et al. (1996) that the present-day Central Iran microplate has rotated by 1358 counterclockwise since the Triassic. This model has usually been accepted (Davoudzadeh & Weber-Diefenbach 1987; Ruttner 1993; Alavi et al. 1997; Seyed-Emami 2003), but a

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 287–321. DOI: 10.1144/SP312.14 0305-8719/09/$15.00 # The Geological Society of London 2009.

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test carried out by Wendt et al. (2005) on the basis of the palaeogeographical distribution of Palaeozoic facies actually failed to confirm the model. At the local scale there are often different interpretations of the stratigraphic v. tectonic relationships of several sedimentary and metasedimentary units. For instance, in the key areas of Nakhlak and Anarak, very different pictures were provided not only for the complexly metamorphic Anarak Range (e.g. Sharkovski et al. 1984; Bagheri 2007; Bagheri & Stampfli 2008), but also for the nearby unmetamorphosed Nakhlak Range (e.g. Sharkovski et al. 1984; Alavi et al. 1997). In order to clarify the general setting of Central Iran and its related Cimmerian history we selected the Nakhlak–Anarak area, which has been known since the 1970s (Davoudzadeh & Seyed-Emami 1972) for the contrast of its peculiar 2.4 km-thick mixed siliciclastic, volcaniclastic and carbonatic succession with the surrounding Triassic successions (i.e. Shotori Range, Tabas: Sto¨cklin et al. 1965; see Seyed-Emami 2003 for a comprehensive summary). Based on lithological similarity, several authors suggested a correlation between the Nakhlak Triassic and an almost coeval succession exposed at Aghdarband (Koppeh Dag, NE Iran). Moreover, this correlation was used to support the 1358 counterclockwise rotation of Central Iran since the Triassic (Davoudzadeh et al. 1981). Despite its great significance it is worth noting that the geology and stratigraphy of the Nakhlak –Anarak area are not known in detail, as demonstrated by the rather different descriptions available in the literature (Davoudzadeh & Seyed Emami 1972; Sharkovski et al. 1984; Alavi et al. 1997; Vaziri 2001). The area was visited by a team of stratigraphers, structural geologists, palaeontologists and palaeomagnetists in 2003 and 2004. Most of the

stratigraphic data are presented here, while the geological, structural and the palaeomagnetic analyses are described in separate contributions (Muttoni et al. 2009; Zanchi et al. 2009).

The geology of the Nakhlak area: open questions The small village of Nakhlak lies near an old, but still active, lead mine located in a small NW– SE-oriented range, the Kuh-e-Qal’eh Bozorg, about 60 km NNE of the town of Anarak (Yazd block) in central Iran (Fig. 1a, b). The range is about 13 km in length, and the maximum elevation is 1439 m a.s.l. (above sea level) (Fig. 2). From this site, Davoudzadeh & Seyed Emami (1969, 1972) described a very thick Triassic succession consisting of the Alam, Ba¯qoroq and Ashin formations (Nakhlak Group), which were referred to Early Triassic –Late Ladinian by Tozer (1972) on the basis of ammonoids and daonellids. The Triassic succession exposed in the southwestern and central sectors of the Kuh-e-Qal’eh Bozorg (in the following text simplified as the Nakhlak range) was thrusted southwards onto ophiolites (Holzer & Ghasemipour 1973) and truncated with an angular unconformity by shallow-water limestones of Late Cretaceous age. Since its first description, the Nakhlak Group turned out to be totally different from the coeval successions of Central Iran (e.g. Shotori Range: Sto¨cklin et al. 1965) because of its great thickness, on the one hand, and its siliciclastic and volcaniclastic content, on the other. For these reasons the Triassic of Nakhlak was always taken into consideration by authors dealing with the Cimmerian history of Iran (Davoudzadeh et al. 1981; Ruttner 1993;

Fig. 1. Location maps of Nakhlak (central Iran). (a) Structural model of Iran, from Angiolini et al. (2007). (b) Road connections of the studied locality (co-ordinates system from GSI 1:250 000 geological map Anarak).

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Fig. 2. Panoramic view of the Kuh-e-Qal’eh Bozorg from the SW. The range is divided into two steep cliffs consisting of sandstones and limestones of Late Cretaceous age. The maximum elevation of the northern and southern cliffs is respectively 1439 and 1435 m a.s.l. The Triassic succession is exposed on the very gentle hills in the foreground on the left and central part of the picture. Dotted line shows the studied areas.

Soffel et al. 1996; Alavi et al. 1997; Saidi et al. 1997; Seyed-Emami 2003). However, despite its great importance, the knowledge of the geology and stratigraphy of Nakhlak area has been based on relatively few data, probably because the area was visited very few times after 1972 (Sharkovski et al. 1984; Vaziri 1996, 2001; Alavi et al. 1997). The few authors who studied the geology and stratigraphy of Nakhlak reported only general information on the stratigraphic succession such as lithology, few elements of lithofacies and qualitative descriptions of conglomerates. No information on sandstone petrography was provided, and dating was based on macrofossils that were not collected bed-by-bed. Most surprisingly, the picture of the Nakhlak area to emerge from our work is notably different. Two main points differ from the literature. The first is the geological setting of the Nakhlak range and the second is the lithostratigraphy of the lower part of the Nakhlak Group, in particular of its oldest unit, the Alam Formation. The geological setting of the Nakhlak range has been interpreted in three different ways (Fig. 3a–c). In their pioneering works, Davoudzadeh & Seyed-Emami (1969, 1972) interpreted the Nakhlak Group as a NE-dipping monocline truncated with angular unconformity by Cretaceous shallow-water limestones. Holzer & Ghasemipour (1973) and Sharkovski et al. (1984) recognized that the Triassic succession is thrusted onto ophiolites, and confirmed the angular unconformity between the Triassic succession and the Upper Cretaceous sediments. However, they illustrated a more complex setting of the Triassic succession (Fig. 3b), forming a NW –SE-oriented syncline. This reconstruction implies folding of the Triassic succession before the uplift and erosion of the Nakhlak Group

between the Late Triassic and the Late Cretaceous (Sharkovski et al. 1984, p. 122), i.e. in Cimmerian times. This tectonic setting and the related history were not confirmed by Alavi et al. (1997). They detected thrust faults within the Triassic succession and between the Nakhlak Group, the Cretaceous and the Palaeocene units (Fig. 3c), documenting that the tectonic setting of the area is at least partially post-Cimmerian. The second point of disagreement in the literature is the lithology of the lower part of the Nakhlak Group, i.e. of the Alam Formation. According to Davoudzadeh & Seyed-Emami (1972, fig. 4, pp. 10–14), the 873 m-thick Alam Formation consists of a wide variety of lithologies also containing some intervals dominated by limestones. In contrast, Alavi et al. (1997, fig. 3, p. 1568) described the 1350 m-thick Alam Formation as mostly consisting of volcanic sandstone lithologies with rare occurrence of detrital/shallow-water limestones.

Geological setting of the area The structural setting of the Nakhlak Range is much more complex than reported in the literature. The description is provided by Zanchi et al. (2009), and here only the main features are summarized. The entire Triassic succession is stacked on hornblende metagabbros and silicified ultramafics south of Nakhlak along a low-angle trending fault. These rocks were interpreted as ophiolites by Holzer & Ghasemipour (1973), Alavi et al. (1997) and Bagheri & Stampfli (2008). The Triassic succession is strongly folded and faulted. Tectonic repetitions occur especially in the northern part of the area (see also Zanchi et al. 2009; Fig. 4), but all the structures are

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Fig. 3. The different interpretations of the geology of Nakhlak area. (a) Monocline truncated by the Cretaceous with angular unconformity (from Davoudzadeh & Seyed-Emami 1972). (b) Syncline of Triassic sediments covered with angular unconformity by the Upper Cretaceous (from Sharkovski et al. 1984). (c) Monocline overthrusted by three structural units (Alavi et al. 1997, fig. 1B; the co-ordinate system is probably different in to that (b)). (d) Satellite image (ASTER) of the Nakhlak area, with lithostratigraphic boundaries of Triassic units (white dashed line), angular unconformity (yellow line), major faults (red lines) and position of the studied stratigraphic sections (white line). See the text for an explanation on the lithostratigraphy of the Cretaceous succession.

unconformably covered by the Upper Cretaceous limestones that seal this complex deformation with a sharp erosional surface. The Upper Cretaceous and Palaeocene beds are also deformed, showing an asymmetric anticline developed along an important north–south strike-slip dextral fault. This fault is possibly related to the Neogene evolution of central Iran (Walker & Jackson 2004), but it is also of great interest for mining because lead mineralization (cerussite and galena) of the active Nakhlak mine are concentrated in veins along the fault zone (Holzer & Ghasemipour 1973). Minor strike-slip faults orthogonal to the pre-Cretaceous deformational structures also deform the Triassic succession. The ‘Upper Cretaceous’ unit (Cenomanian– Turonian) was studied only by Holzer & Ghasemipour (1973) and, more recently, by Vaziri et al. (2005). This unit has never been formalized according to the rules of lithostratigraphic nomenclature, but it is mapped as a formation in the Geological

Survey of Iran (GSI) 1:100 000 Nakhlak sheet and 1:250 000 Anarak sheet. The lower part of the Upper Cretaceous is transgressive and consists of some tens of metres of sandstones and hybrid arenites followed by shallow-water rudist-rich limestones. It was deposited with a spectacular angular unconformity on the deformed Triassic succession (Fig. 4). The lower boundary of the Upper Cretaceous is erosional and not tectonic, as reported by Alavi et al. (1997). The sedimentary nature of this boundary is also indicated by the occurrence of polimictic conglomerates in pockets (Fig. 4d, e). Some reverse faults can be locally detected in the lower part of the Upper Cretaceous unit, but often they reactivate the lithological boundary of the conglomerate in pockets and the overlying sandstones rather than the unconformity, i.e. the base of the conglomerates. In both examples the displacement is only of a few metres (see also Zanchi et al. 2009, fig. 6A, B).

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Fig. 4. General and detailed views of the angular unconformity between the Triassic succession and the Upper Cretaceous in the southern cliff of the Kuh-e-Qal’eh Bozorg. (a) General view of the angular unconformity between the Alam Formation and the Upper Cretaceous. (b) Angular unconformity between the Ba¯qoroq Formation and the overlying Upper Cretaceous. (c) Small boulder of shales from the Alam Formation included in a pocket at the base of the Upper Cretaceous sandstones. (d) Detail of the contact between the Alam Formation and the Upper Cretaceous. (e) Pockets of conglomerates at the base of the Upper Cretaceous sandstones. Acronyms: cg, conglomerate; sd, sandstone.

Stratigraphic sections and sampling Owing to the gentle topography, the great stratigraphic thickness and the deformation, there is no single complete exposure of the Nakhlak Group. The composite section across the group (Figs 3d and 5) is based on

nine stratigraphic sections that are correlated on the basis of distinct lithological markers. In the studied part of the Nakhlak range, the total thickness of the Nakhlak Group is about 2400 m, and it is subdivided as follows (from bottom to top): Alam Formation (1150 m); Ba¯qoroq Formation

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Fig. 5. Composite stratigraphic section of the Nakhlak Group, with the stratigraphic position of the studied sections. The GPS co-ordinates (WGS84 system) of sections Nakhlak 10 and 7 are shown. For the sections within the Alam Formation the co-ordinates are shown in the detailed logs. The stratigraphic position of the Nakhlak 3 section with respect to the base of member E is measured with GPS and compass. The members of the Alam Formation are labelled following the stratigraphic succession (A–G); the limestone members are emphasized in yellow. The approach to the lithostratigraphy of the Alam Formation is explained in the text. The legend is for all the stratigraphic sections.

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(870 m); and Ashin Formation (.370 m). The stratigraphic sections were measured with great detail and sampled bed-by-bed for ammonoids, conodonts, bivalves, limestone and sandstone petrographic analyses, as well as for palaeomagnetic analyses. About 680 ammonoids were collected, but most of them are of too small a size or too fragmentary for a complete identification. Conodonts were never previously reported from the Nakhlak Group; 49 samples were processed in this study and 12 samples (24%) yielded conodonts. The petrographic composition of conglomerate pebbles was recognized in the field and verified using thin sections. Sandstone petrography has been determined for the first time.

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sandstones, conglomerates, tuffaceous conglomerates, tuffs, siltstones, green, red and grey shales– marls, as well as a wide variety of limestones. However, the very wide variety of rocks is organized into lithologically distinct intervals. Davoudzadeh & Seyed-Emami (1972) proposed the subdivision of the Alam Formation into six members that were further subdivided into 13 intervals. This detailed subdivision does not fit very well with the stratigraphic succession observed in our study area, thus we amend Davoudzadeh & Seyed-Emami’s lithostratigraphy and propose the subdivision of the formation into seven members (members A–G). The lack of toponyms in the study area prevents a complete formal definition of these members according to the rules of the International Stratigraphic Guide (Salvador 1994). However, we emphasize that the subdivision of the formation in members is justified because of the following reasons:

This 1150 m-thick unit shows a very wide lithological composition comprising sandstones, tuffaceous

† the Alam Formation is very thick and covers an interval of at least 6 Ma between the Olenekian

Stratigraphy of the Nakhlak Group

Fig. 6. (a) General view of the lower part of the Alam Formation at Nakhlak 1 section. (b) and (c) Details of member B ‘oolitic limestone’. (b) Oolitic grainstones with cross-bedding (south of section Nakhlak 2); the scale is 14 cm. (c). Oolitic grainstone with bioclasts of echinoderms, also as ooids’ core, gastropods and brachiopods; some ooids have slightly micritized rim (sample KN121, south of section Nakhlak 2).

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and the Anisian (radioisotopic data from ammonoid-controlled section from south China: Ovtcharova et al. 2006; Galfetti et al. 2007). The subdivision in members is worthwhile from the point of view of the description of the succession; † several palaeoenvironments are documented in the Alam Formation, thus the subdivision into members is useful; † the members have a distinct stratigraphic position within the Alam Formation; † five members show such a typical lithology that they can be recognized in the field even in small outcrops. The remaining three members can be recognized on the basis of the combination of lithology and stratigraphic position with respect to the other members. Member A (?lower Olenekian). The lower boundary of the Alam Formation is tectonic, as already described by Sharkovski et al. (1984) and Alavi et al. (1997). The lowest part of the Alam Formation is usually not very well exposed. The best outcrop in the studied part of the Nakhlak Range is shown in Figure 6a, and the log in Figure 7. Member A consists of 18 m of brownish siltstones and dolomites, overlain by 35 m of dark brown siltstones, volcanic sandstones and shales. The unit is probably equivalent to intervals 1 and 2 of member 1 of Davoudzadeh & Seyed-Emami (1972).

The unit does not yield any fossils and the early Olenekian age is suggested on the basis of its stratigraphic position. Member B ‘oolitic limestone’ (?lower Olenekian). This member is a very distinct lithological marker that can be recognized within the lower part of Alam Formation, and is probably equivalent to interval 3 of member 1 of Davoudzadeh & Seyed-Emami (1972). Member B is up to 38 m thick and is a little better exposed than is member A. In the study area it crops out at the Nakhlak 1 section (Figs 6a and 7) and south of the Nakhlak 2 section. The lithology consists of oolitic bioclastic grainstones forming 80 cm – 2 m-thick beds with planar bedding, planar lamination and cross-bedding (Fig. 6b, c). The boundaries are sharp, and can be traced on the first and on the last occurrences of oolitic grainstones. The member does not yield any well-preserved fossils; the possibly early Olenekian age is supposed on the basis of its stratigraphic position. Member C ‘varicoloured shales’ (middle –upper Olenekian). Member B is conformably overlain by another lithological marker, member C, consisting of an approximately 30 m-thick sandstonedominated lower part and an up to 127 m-thick green –red/purple shale-dominated upper part with relatively common ammonoids. This lithological marker can be easily followed for at least 2 km along strike, i.e. through all of the studied part of the Nakhlak range. This unit had already been recognized by Davoudzadeh & Seyed-Emami (1972, p. 14: interval 4 of the member 2), and is exposed along the Nakhlak 1 and Nakhlak 2 sections (Fig. 8). Boundaries. The lower boundary is rather sharp and is drawn at the top of the uppermost limestone of member B at the Nakhlak 1 section. The upper boundary is drawn at the base of the first red nodular limestone bed of member D at the Nakhlak 2 section.

Fig. 7. The lower part of the Alam Formation at Nakhlak 1 section, where the succession of members A to C is exposed. Sample KN121 is from an outcrop some tens of metres south of the base of the Nakhlak 2 section.

Lithofacies. Member C is divided into a lower part dominated by sandstones and an upper part dominated by shales. The lower part is documented at both the Nakhlak 1 and the Nakhlak 2 sections (Figs 7, 8 and 9a). At the Nakhlak 1 section it is 27 m thick, whereas at the Nakhlak 2 section it is at least 30 m thick. The lithology consists of alternations of green volcaniclastic sandstones and hybrid arenites with green shales. Thickness of the sandstone– hybrid arenite beds is from 15 cm to about 1 m. Shales usually occur in thinner interbeds and the sandstone– shale ratio is from 2:1 to 4:1 (Fig. 9a). The sedimentary structures are normal grading, planar and more rarely cross-laminations.

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Fig. 8. Nakhlak 2 section. The section consists of two segments that are correlated by the boundary between the lower and upper part of member C. This boundary is sharp, thus it is easy to detect in the field (see also Fig. 9a). The first segment is exposed on the right-hand side of a small valley. At the top of the segment the valley is divided into two branches. The second segment of the section is measured on the right-hand side of the east branch.

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Fig. 9. Lithofacies of the Alam Formation. (a) The boundary between the lower and upper part of member C (Nakhlak 2 section, I segment). (b) Shales with thin-bedded limestones and sandstones of the upper part of member C (Nakhlak 2 section, position shown in (a)); the scale is 100 cm. (c) Ammonoid Albanites osmanicus from member C (Nakhlak 2 section, sample NK136); the scale is 20 mm. (d) Outcrop view of member D (Nakhlak 2 section, interval between samples NK26 and NK27) showing the typical thin layering of nodular limestone, here with a thin sandstone bed (sd); the scale is 20 cm. (e) nodular ammonoid-bearing limestones of member D ‘nodular limestone’ (Nakhlak 2 section, sample NK32); the scale is 15 mm. (f) basal part of member G (Nakhlak 4 section); the scale (felt-tip) is 14 cm. (g) member G ‘grey shales’: detail of the conglomerates (Nakhlak 6 section, sample AL1); the scale is 60 cm. (h) member G: close-up view of a bivalve-bearing marly limestone (Nakhlak 6 section, sample NK190); the scale is 14 cm.

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The upper part is documented at the Nakhlak 2 section. The lithology consists of an alternation of green or red/purple shales –very-fine-grained siltstones and of very-thin-bedded limestones and sandstones (sandstone –limestone/shale ratio 1:2– 1:5) (Fig. 9a, b). The average thickness of the shale intervals is from 40 to 60 cm, whereas the sandstones are 3–14 cm thick (maximum 30 cm). The shales – siltstones are organized into reddish/purplish or green intervals from 40 cm to several meters thick. The upper 22–24 m of the member show a uniform green colour. Fossils and age. The upper part of this member is well known from the literature (Davoudzadeh & Seyed-Emami 1972; Tozer 1972) as ammonoid bearing, but actually the ammonoids are quite rare. Most of them were collected following the beds along strike or from lateral outcrops correlated to the Nakhlak 2 section. A selection of the most significant ammonoid samples is shown in Figure 8. Two conodont samples were taken from the lower part of the member, but they were barren. The ammonoids can be found directly interbedded within the shales (Fig. 9c) or on the surface of the beds. The first type of preservation mostly leads to incomplete specimens, consisting of a body chamber with only part of the phragmocone, while the specimens preserved on the bed surface are more complete, but difficult to extract from the matrix. The taxonomic analysis of the new collection confirms the faunal list provided by Tozer (1972), with the only exception of Procolumbites karataucikus Astakhova (Fig. 10b). This species is identified for the first time at Nakhalk, but a specimen conspecific with this taxon had already been collected by Tozer, even if it was attributed to Paragoceras mediterraneum (Arthaber) (Tozer 1972, plate 1, fig. 2a, b). The bed-by-bed collection allows the first recognition of the distribution of the taxa within member C. Two main points can be emphasized. The first is that the ammonoids were found through all of the upper part of member C (Figs 8 and 10). The second is that there is not a single fauna in this interval, as some of the genera are confined to the lower, middle or upper part of the interval. The ammonoid assemblages (Fig. 10) are dominated by Albanites (A. arbanus [Arthaber] and A. osmanicus [Arthaber]), which represents more than 50% of the collected specimens. In decreasing order of frequency, the second group of ammonoids are the Columbitidae, represented by Columbites, Subcolumbites and Procolumbites in stratigraphic order. Tirolites is documented only in the upper part of the member, together with Procolumbites and the long-ranging Stacheites (S. undatus [Astakhova] and S. concavus Shevyrev).

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Columbitidae and Tirolites allow the chronostratigraphic attribution of member C to the midOlenekian (early– mid-Spathian sensu Tozer 1981; late early Spathian sensu Ovtcharova et al. 2006 and Galfetti et al. 2007). More accurate attribution is difficult because, at present, there is not one global ammonoid scale for the Olenekian stage, but some regional scales consisting of informal subdivisions (beds) not yet fully described and formalized according to the rules of the International Stratigraphic Guide (Salvador 1994). Following Tozer’s (1981) ammonoid scale the fossiliferous interval can be assigned to the Columbites and Subcolumbites beds, while following the scale recently outlined in south China (Ovtcharova et al. 2006; Galfetti et al. 2007) it can be correlated with the Tirolites/Columbites and the Procolumbites beds. It is worthwhile to note that while in member C the range of Tirolites overlaps the range of Procolumbites, this overlap does not seem to be documented in China. The fauna of member C is also coeval with the ammonoid fauna of Mangyshlak (Shevyrev 1968; Balini et al. 2000). However, the ammonoid assemblages of member C are much more similar to the Tethyan fauna of Kc¸ira in Albania (Arthaber 1908, 1911; Germani 1997) and Chios (Renz & Renz 1948; Mertmann & Jacobshagen 2003) than to the fauna of Mangyshlak. Member D ‘nodular limestone’ (upper Olenekian). Member C is conformably overlain by a very typical unit (member D) consisting of very-thinbedded pink nodular limestones with ammonoids, and rare thin-bedded sandstone intercalations with a thickness of about 80 m. This member can be followed along strike in the studied area of the Nakhlak range, but it is also documented in the northern part (north of elevation 1439 m). This member is equivalent to unit 5 of member 2 described by Davoudzadeh & Seyed-Emami (1972), but it was not recognized by Alavi et al. (1997). The member is best exposed at the Nakhlak 2 section (Fig. 8). Boundaries. The lower boundary is traced at the base of the first red nodular limestone layer. The upper boundary is located on the top of the last red nodular limestone bed. Lithofacies. The lithology is quite monotonous and consists of thin-bedded nodular limestones in beds from 1 to 3 –5 cm thick, sometimes amalgamated into 20 –30 cm-thick beds (Fig. 9d, e). Nodules are light pink (yellow weathering) and very small (5–10 mm). They are bounded by red– purple abundant shales/shaly marls. Some 3 –10 cm-thick graded sandstone beds also occur (Fig. 9d).

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Fig. 10. Olenekian ammonoids from member C ‘varicoloured shales’. Albanites arbanus (Arthaber), NK86-2 (MPUM 9831) (a) lateral view, (a0 ) ventral view and (a00 ) section. Procolumbites karataucikus Astakhova, NK150-1 (MPUM 9832) (b) lateral view, (b0 ) section and (b00 ) ventral view. Albanites osmanicus (Arthaber), NK157-15 (MPUM 9833) (c) lateral view, (c0 ) ventral view and (c00 ) section. Paragoceras mediterraneum (Arthaber), NK157-12 (MPUM 9834) (d) section and (d0 ) lateral view. (e) Columbites ventroangustus Shevyrev, NK145-1 (MPUM 9835), lateral view.

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Fossils and age. Ammonoids occur frequently (Figs 8 and 9e), especially in the upper half of the member. However, the very small size (diameter about 15 mm, rarely up to 25 mm), together with the preservation of their tests, often prevents complete classification. The ammonoid assemblages are dominated by leiostraca, long-ranging ammonoids typical of the red nodular ‘Ammonitico Rosso’-like limestones. In this case Procarnites and Isculitoides are the most common genera. Dagnoceras sp. indet. has also been identified in sample NK29ter. Isculitoides and Dagnoceras are typical of the Spathian, but their range is not yet well calibrated. Taking into account also the age of the underlying member C, a late Olenekian age can be suggested for member D. The assignment to the Olenekian is also confirmed by the occurrence of the Olenekian conodont Neospathodus homeri (Bender) in sample NK22. Member E (upper Olenekian –lower Bithynian). Member E is about 330 m thick and shows a rather wide lithofacies variability both along and across strike. The interval is equivalent to member 3 of Davoudzadeh & Seyed-Emami (1972). In all the studied sections the lower part of the member is not well exposed. The thickness of the member has been estimated by GPS and compass measurements of the stratigraphic position of the Nakhlak 3 section with respect to the base of the member (Figs 5 and 11). Boundaries. The lower boundary of member E is drawn at the top of the last red nodular limestone bed of member D. In the central part of the study area member E is conformably covered by member F. In the eastern part of the study area, member F interfingers with the upper part of member E. The latter is covered by member G and the boundary is traced at the top of the last planar limestone bed. Lithofacies and fossils. Member E is dominated by siliciclastic and volcaniclastic sediments with rare limestones in the lower part (the Nakhlak 2 section: Fig. 8). Limestones become gradually more frequent in the upper part (the Nakhlak 3 section: Fig. 11). From east to west, i.e. from the Nakhlak 2 and 3 sections to the Nakhlak 5 section (Fig. 12), the succession becomes more siliciclastic,

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with a strong reduction in the carbonates. The following description of the succession is mostly based on the Nakhlak 2 and 3 sections. The distribution of fossils is described together with the lithofacies because it is strictly connected with the distribution of the limestone intervals. In the lower part of the member there are paraconglomerate beds with erosional bases ranging from 3 to 5 m thick. They yield blocks with low roundness and sphericity, and with an average maximum size of between 10 and 30 cm (cobbles), but sometimes up to 1 m (boulders), within a tuffaceous sandy matrix. The cobbles and boulders consist of grey–green sandstones, limestones and pink nodular limestones, which can be referred to members E, D ‘nodular limestone’ and, possibly, the basal part of member C of the Alam Formation. The middle part of the member (lower part of the Nakhlak 3 section: Fig. 11) consists of dark grey– green, black weathered volcaniclastic sandstones, often organized in thickening- and coarseningupwards cycles (Fig. 13). The limestones are rare in this part (NK45–NK47; NK48; NK50; NK53bis-53) and consist of thin-bedded light-brown–light-green bioclastic mudstones with rare ammonoid crosssections. Samples NK402 and NK48 (Fig. 14a–e) provided conodonts Neospathodus homeri (Bender) and N. gondolelloides (Bender), which are Olenekian in age. Chiosella timorensis (Nogami), marker of the Olenekian–Anisian boundary has been identified in sample NK53 together with Neogondolella regale (Mosher), Ng. sp. A (Fig. 14f-g), occurring also in sample NK53bis and the ammonoid Leiophyllites sp. The upper part of the member (Nakhlak 3 and Nakhlak 9 sections: Fig. 11) is characterized by an increase in wavy- (pseudonodular) to planar-bedded bioclastic packstones with crinoids, brachiopods, gastropods and rare volcanic quartz (samples NK55– NK68) (KN61–KN74 at Nakhlak 9). Sometimes the bivalves Plagiostoma cf. striatum and Neoschizodus sp. (sample NK431 and NK433, Fig. 15) also occur. The uppermost part of the member consists of wackestones with filaments, some crinoids and brachiopods (from sample NK72bis to NK75 at Nakhlak 3; sample KN82 at Nakhlak 9). Two Anisian conodont assemblages have been identified (Fig. 11). Sample NK63 yielded Neogondolella cf. regale and transitional Ng. regale-Paragondolella bulgarica. The second fauna is documented in

Fig. 10. (Continued) Subcolumbites europaeus (Arthaber), NK153-10 (MPUM 9836) (f) lateral view, (f0 ) ventral view and (f00 ) section. Tirolites sp., NK215-12 (MPUM 9837) (g) lateral view and (g0 ) ventral view. (h) Stacheites undatus (Astakhova), NK159-21 (MPUM 9838), lateral view. Stacheites concavus Shevyrev, NK168-1 (MPUM 9839) (i) ventral view and (i0 ) lateral view. All the specimens are housed at the Museo di Paleontologia, Dipartimento di Scienze della Terra, Milano University; the inventory number is shown in brackets. All the specimens are whitened with ammonium chloride. The scale bar for all specimens is 1 cm.

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Fig. 11. Stratigraphic logs of sections Nakhlak 3 and 9 showing the upper part of member E and the basal part of member F. The position of the two sections is shown in Figures 13b, 16a and 17. The black star shows the marker level that has been used to calibrate with GPS and compass the Nakhlak 3 section with respect to the top of member D.

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Aghdarbandites ismidicus (Arthaber) (Fig. 18a) and Nicomedites, together with the conodonts Neogondolella regale (Mosher), P. bulgarica (Budurov & Stefanov) and transitional N. regale– N. constricta cornuta. Acrochordiceras was found in sample NK106 and Kocaelia toulai (Arthaber) occurs in sample NK452 (Fig. 18b). The brachiopod Tetractinella is also rather common.

Fig. 12. Upper part of the Nakhlak 5 section. The section is 250 m thick and consists of a monotonous succession of sandstones that are exposed on the right-hand side of a wide valley located in the western part of the study area (see Figs 2 and 3d). Here only the upper part is shown. The three limestone intervals are coeval with member F. The age of this part of the section is Bithynian on the basis of lithological correlations with the Nakhlak 3, 4 and 9 sections.

sample NK435 and consists of transitional forms P. bulgarica–P. hanbulogi while NK436 consists of Ng. cf. constricta cornuta. Both faunas can be attributed to the late Bithynian. The uppermost part of member E includes the basinal facies coeval with member F ‘carbonate mounds’ (Nakhlak 9 and 3 sections: Fig. 16a; Nakhlak 4 section: Fig. 17). Lithology is dominated by an alternation of grey– light green shales –marls with ammonoid-bearing bioclastic mudstones. Sample NK77 yielded the Bithynian ammonoids

Age. The late Olenekian–late Bithynian age of member E is well defined by the integration of ammonoid and conodont data. The lowest fossiliferous interval (NK402 and NK48) is late Olenekian in age. The occurrence of Chiosella timorensis (Nogami) in sample NK53 marks the base of the Aegean, while the Aghdarbandites ismidicus Zone (late Bithynian) is documented by its index fossil, as well as by Kocaelia toulai (Arthaber), which is an additional marker of this zone (Fantini Sestini 1988, 1990). Member F ‘carbonate mounds’ (Bithynian). Member F consists of at least two coalescent carbonate bodies (Fig. 16) that reach a maximum thickness of about 50 m and have a present-day lateral length of about 500 m. This rock body had already been mapped as a distinct unit by Holzer & Ghasemipour (1973). It is equivalent to part of Davoudzadeh & Seyed-Emami’s (1972) member 4, and it was also recognized by Alavi et al. (1997). Boundaries. The upper part of the underlying member E shows an increasing amount of carbonate debris (mainly skeletal), but makes a sharp boundary with the bottom of member F. Member F interfingers with member E and is capped by the member G. The upper boundary is sharp, and the flat top of the mound is characterized by abundant ammonoids and crinoids.

Fig. 13. Outcrop views of the upper part of member E. (a) Coarsening- and thickening-upwards cycles in the upper part of the Nakhlak 5 section. The beds on top of the slope are the limestone interval with samples NK162 and NK163 (see Fig. 12). (b) Lower part of the Nakhlak 3 section: the dashed line shows the trace of the section, the dotted line marks the base of the conglomerates at the base of the section; in the background is the II carbonate mound of member F ‘carbonate mounds’ and the upper part of the Nakhlak 9 section.

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Fig. 14. Conodonts from the Alam Formation. Neospathodus gondolelloides (Bender), NK48 (a) upper view and (a0 ) lateral view. (b) Neospathodus gondolelloides (Bender), NK402, lateral view. Neospathodus gondolelloides (Bender), NK402 (c) upper view, (c0 ) lateral view and (c00 ) oblique/lower view. Neospathodus homeri (Bender), NK402 (d) lateral view and (d0 ) lower view. (e) Neospathodus gondolelloides (Bender), NK48 oblique/lower view. Neogondolella sp. A, NK53 (f) upper view and (f 0 ) lower view. Chiosella timorensis (Nogami), broken specimen, NK53 (g) upper view, (g0 ) lateral view and (g00 ) lower view. The scale bar for all specimens is 200 mm.

Lithofacies. The mounds are characterized by a massive portion that passes laterally to clinostratified (angles evaluated in about 108–158), bedded bioclastic limestones that cover the underlying siliciclastics and carbonates. Two mounds can be distinguished and separated by intercalation of finegrained sandstones in the marginal part. They seem to be amalgamated in the innermost part, where the massive portions are coalescent. The carbonate mounds pinch out laterally, and represent isolated carbonate bodies that interfinger with the prevailing volcanoclastic succession of member E. Whereas the lateral transition between the mounds and the siliciclastics is gradual, the sharp topmost boundary of the mounds is indicative of a rapid change in the depositional environment. Different facies assemblages have been recognized, on the basis of the sedimentological structures and of the microfacies association (Fig. 16b–g). The inner part of the mounds is characterized by calcimicrobial limestones. They consist of massive clotted

framestones, characterized by open-space structures filled with fine-grained internal sediments or spar calcite. Bioclasts are generally rare and consist of large isolated sponges (up to 5 cm in diameter), ostracods, serpulids and bryozoans. Bryozoans and serpulids grow on calcimicrobial limestones, thus indicating hard substrate. Close to the margins of the mounds, bioclastic limestones (rich in crinoids, echinoids, bivalves, gastropods, brachiopod, bryozoans and green algae) or intraclastic packstones prevail. Intraclasts are locally present, together with extrabasinal volcanic fragments. Reworking of the bioclasts is generally high, as documented by mechanical erosion and by rounded shape. Basinwards, the dominant facies is represented by hybrid limestones. They consist of wackestones with a calcareous matrix and abundant fine-grained quartz-rich volcanic material, with sparse bioclasts. Bioclasts are represented by thin-shelled bivalves, brachiopods, crinoid ossicles, ostracods and bryozoans. Hybrid limestones with lithic fragments are common.

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Fig. 15. Bivalves from the Alam Formation, members E and G ‘grey shales’. (a) Plagiostoma cf. P. striatum, left valve, NK431 (MPUM 9842-1). (b) Plagiostoma cf. P. striatum, right valve, NK433 (MPUM 9843). (c) Neoschizodus sp., right valve, NK433 (MPUM 9844). (d) Unionites sp., left valve, NK454, specimen from (e). Marly limestone showing monotypic assemblage consisting of Unionites sp., NK454 (MPUM 9840) (e) lower surface and (e0 ) upper surface. (f) Plagiostoma cf. P. striatum, right valve, NK431 (MPUM 9842-2). (g) Unionites sp., NK454, specimens from (e). (h) Unionites sp., NK454 (MPUM 9841). All the specimens are housed at the Museo di Paleontologia, Dipartimento di Scienze della Terra, Milano University; the inventory number is shown in brackets. All the specimens are whitened with ammonoium chloride. The scale bar for all specimens is 10 mm.

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Fig. 16. Member F ‘carbonate mounds’. (a) View of the mounds from the south. Note the transition from the massive carbonates (CM; inner part of the mounds) to the basinal facies (BF; member E) across clinostratified low-angle slope facies (SF). In the background is the Upper Cretaceous succession that unconformably covers the Alam Formation. The Nakhlak 9 section (Fig. 11) was measured along the right slope. (b) Detail of the massive limestones of the inner part of the mound. Note the presence of light-grey cavities filled by internal sediments. (c) Clinostratified bioclastic (crinoidal) limestones of the slope facies. (d) Bioclastic packstone from the slope facies of the mound. Skeletal grains are mainly represented by gastropods and crinoids. Note that several skeletal grains are rounded, suggesting an intense reworking of the bioclasts on the sea bottom. Sample NK55. (e) Bioclastic packstone with extrabasinal grains (quartz). Transition from slope to basinal facies. Sample NK61. (f) Massive wackestone with bryozoans in life position. In the top left-hand corner note the cavity filled with light-coloured internal sediments. Sample KN109. (g) Intra-bioclastic packstone at the transition between the massive mound and the slope facies. Some clasts are encrusted by organisms (arrow). Most of the intraclast consists of clotted micrite, probably derived by the inner part of the mound. Sample NK109.

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Fig. 17. The Nakhlak 4 section showing the upper part of member E and the lower part of member G ‘grey shales’. The field photography shows the position of the section with respect to the II carbonate mound and the upper part of the Nakhlak 3 section.

The geometry of the mound and the presence of shallow-water structures (i.e. hummocky crossstratification; Fig. 11) indicate that the carbonate production occurred above storm-weather wave base. The top of the mound was likely within the lower photic zone, suggesting that the microbial

carbonate was probably produced by a lightsensitive community. The Anisian carbonate mounds of Nakhlak represent a particular carbonate factory that closely resembles the Early Triassic mounds of Southern China (Lehrmann 1999). They were, however,

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Fig. 18. Bithynian ammonoids from member E of the Alam Formation. Aghdarbandites ismidicus (Arthaber), NK77-37 (MPUM 9850) (a) lateral view and (a0 ) ventral view. Kocaelia toulai (Arthaber), NK452-5 (MPUM 9851) (b) lateral view and (b0 ) ventral view. All the specimens are housed at the Museo di Paleontologia, Dipartimento di Scienze della Terra, Milano University; the inventory number is shown in brackets. All the specimens are whitened with ammonoium chloride. The scale bar for all specimens is 1 cm.

developed in different bathymetric conditions; subtidal for Nakhlak and peritidal for southern China. Fossils and age. Member F is quite fossiliferous, but most of the fossils can be recognized only in thin sections. Nevertheless, some age-diagnostic ammonoids and conodonts occur below and above the member, as well as in coeval basinal facies, firmly fixing the chronostratigraphic position of the member. The base of the member is Bithynian, probably close to the boundary between the Nicomedites osmani and Aghdarbandites ismidicus zones. The top is late Bithynian (Aghdarbandites ismidicus Zone) on the basis of the ammonoids found at the very base of the member G. Brachiopods were found in the limestonedominated lithofacies of the top of the underlying member E, within the clinostratified facies of the slope of the mounds and in the neritic limestones coeval with the mounds. The genus Tetractinella is rather common. Member G ‘grey shales’ (upper Bithynian– Pelsonian or Illiryan). The upper part of the Alam Formation (Nakhlak 6 section, Fig. 19) is about 420 m thick and is dominated by light-grey to rare light-green shales and marls. The abundance of shales and marls makes this interval very easily recognizable. The member includes a thick interval

attributed by Davoudzadeh & Seyed-Emami (1972, figs 3 and 4) to the upper part of member 4, members 5 and 6 of the Alam Formation, and the lower part of member 1 of the Ba¯qoroq Formation. Boundaries. The lower boundary with member F is traced immediately above the hardground at the top of the carbonate mounds. Where member G overlies member E, its boundary is traced at the base of the 65–78 m-thick subunit consisting of grey shales with marly limestone lenses. The upper boundary with the Ba¯qoroq Formation is located at the base of the first massive red conglomerate bed. Lithofacies. The member can be subdivided into three parts. The lowest interval is about 65 m thick at the Nakhlak 4 section and consists of light-grey to light-green shales with marly limestone lenses (Fig. 9f), sometimes with tuffitic layers. This facies yielded ammonoids and is also documented at the Nakhlak 6 section (about 78 m thick). The intermediate interval is about 200 m thick and is characterized by three main intercalations of 1–4 m-thick lenticular –planar-bedded conglomerates, as well as by quite common sandstones and microconglomerates. The conglomerates are clastsupported (Fig. 9g), and the clasts, consisting of sedimentary rocks, are well rounded with good

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Fig. 19. The Nakhlak 6 section, showing the top of member F, member G and the base of the Ba¯qoroq Formation (B.F.).

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sphericity. In this intermediate part of the member, some bivalve-bearing marls and limestones also occur (Fig. 9h). The upper part of the member shows an increasing amount of fine-grained sandstones, as well as sandstones and microconglomerates. Fossils and age. Ammonoids have been found in the lower part of this member, while bivalves are relatively common in the intermediate and upper part. The lower part of the member at the Nakhlak 4 section (Fig. 19) yielded mostly long-ranging leiostraca ammonoids of the family Gymnitidae, particularly G. religiosus Diener and G. asseretoi Tozer, which are not age-diagnostic. A pelagic leiostracadominated ammonoid fauna is also documented at the Nakhlak 6 section, immediately above the hardground on top of member F. A rich fauna consisting of Kocaelia, Sturia, Leiophyllites and Monophyllites was identified in sample NK453, and the conodont Paragondolella bulgarica (Budurov & Stefanov) was found in the first limestone level (sample NK455). Kocaelia is typical of the Aghdarbandites ismidicus Zone (late Bithynian: Fantini Sestini 1988, 1990; cf. Krystyn & Tatzreiter 1991), thus fixing the late Bithynian age of the base of the member, which is also consistent with the presence of P. bulgarica. The marls and limestones of the intermediate interval of the member yielded an extremely abundant bivalve fauna consisting only of Unionites (Fig. 15e, g, h), but are barren of conodonts (samples NK188 and NK189). They are, therefore, referred to a brackish–freshwater lagoonal environment. The bivalves occur in 5–20 cm-thick layers dominated by specimens with closed valves (autochthonous assemblages), and in comparably thick levels as openarticulate–butterfly specimens mostly concentrated in nests (parautochthonous assemblages: Fig. 9 g). Autochthonous and allochthonous oligotypic Unionites fauna are common in the upper part of the member. Age-diagnostic fossils are lacking in the intermediate and upper part of the member. Thus, the age of the top of the member and top of the Alam Formation is not biochronostratigraphically calibrated.

Ba¯qoroq Formation (?Late Anisian – Ladinian) This weathering-resistant unit is normally very well exposed. The unit was sampled along a single section (Nakhlak 10) located in the eastern part of the working area (Fig. 3d). At this site the overall thickness of the Ba¯qoroq Formation is about 870 m.

Lithology. The Ba¯qoroq Formation (Nakhlak 10 section; Figs 5 and 20) is dominated by red conglomerates and microconglomerates, within which red–yellow sandstones can also be found, especially in the upper part of the unit. The general trend of the unit is slightly fining-upwards. Conglomerates are normally organized in 1–20 m-thick intervals (in the lower part, exceptionally, up to 60 m thick) separated by microconglomerates or sandstones intercalations up to 5 m thick. In the uppermost part of the unit these intercalations are made of lightgrey–light-green siltstones. The conglomeratic intervals are chaotic (lowermost part of the unit), but more commonly they show lenticular- to cross-bedding. Clasts are rounded to well rounded, and discoidal– subspherical. Grain size of the conglomerates covers the range between pebbles and cobbles, but in the chaotic intervals in the lowermost part of the unit small boulders are rather common. Fossil content. The unit is reported as nonfossiliferous in the literature. A conodont sample from a bioclastic limestone layer located in the lowermost part of the formation (12 –15 m from the base) was barren. In the absence of biostratigraphic data, the ?Late Anisian –Ladinian age of the unit is defined on the basis of its stratigraphic position between the Alam and the Ashin formations.

Ashin Formation (Late Ladinian) This unit crops out in the tectonically disturbed part of the Nakhlak range, and only one section (Nakhlak 7) was measured. In this section about 370 m of the formation are exposed. However, the stratigraphic boundary with the Upper Cretaceous is not preserved as the Ashin Formation is locally in tectonic contact with the Ba¯qoroq Formation so that the true thickness of the Ashin Formation is unknown. Lithology. The formation can be subdivided into three parts (Fig. 5). In the lower part the lithology consists of shales and very-fine-grained siltstones, with intercalation of up to 4– 5 m-thick white – grey quartz-rich conglomeratic layers. The general trend of this part of the formation is fining- and thinning-upwards, and the conglomeratic layers are gradually replaced by sandstones that become thinner bedded. In the middle part of this interval there are tuff levels, testifying to syn-sedimentary volcanic activity. The middle part of the formation, starting approximately from sample NK176 (about 75 m above the base of the Nakhlak 7 section), consists of an alternation of fine-grained sandstones in 5–10 cm-thick beds and shales in 10–100 cm-thick beds. The sandstones show the typical sedimentological features of

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Fig. 20. Composite section of the Nakhlak Group showing the palaeobathymetry, the position of sandstone and conglomerate samples, the petrographic composition of the conglomerates of the upper Alam (sites AL2– AL1), Ba¯qoroq (sites BQ10– BQ1) and lower Ashin (sites AS3– AS1) formations, and the composition of sandstones plotted in a Q– F– L diagram (quartz–feldspar–aphanitic lithic grains). The asterisk marks the last occurrence of pebbles of oolitic limestone. GPS co-ordinates of the 15 sampling sites for conglomerates are listed in the Appendix.

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turbidites, such as scour and tool marks, parallel-, cross- and convolute-lamination. The upper part of the Ashin Formation, starting about 270 m above the base, continues the finingand thinning-upwards trend, as the lithology consists of an alternation of 0.5 –1 cm-thick very-fine-grained sandstones or siltstones with 1–5 cm-thick shales. Occasionally, there are limestones up to 5 cm in thickness. Fossil content. The Ashin Formation is known for a rather poor body fossil fauna and more diverse ichnofossil assemblages. Tozer (1972) described the ammonoids Megaphyllites sp. indet., Proarcestes sp. indet. and Arpadites cf. A. szaboi, as well as the bivalve Daonella lommeli from the middle part of the unit. More recently, Vaziri (1996) reported Romanites simionescui from the lower part of the formation. During the fieldwork neither ammonoids nor bivalves were found. The softbottom deep-water trace fossils Agrichnia, Pascichnia and Repichnia (Nereites ichnofacies), as well as iso-oriented crinoid stems, are quite common, especially in the middle part of the Ashin Formation (see also Vaziri & Fu¨rsich 2007). The fine-grained and often siliciclastic lithology is not the best for conodont analysis. Only one conodont sample was taken from a limestone level in the upper part of the section (sample NK178, about 310 m from the base), but it was barren. In the absence of new data, the chronological assignment of the Ashin Formation relies on data from the literature. Not one of the ammonoids reported from the Ashin Formation is a shortranging form: Arpadites is a Ladinian genus, and Romanites simionescui in uncondensed successions is found only in the Upper Ladinian (see Krystyn & Tatzreiter 1991). The bivalve Daonella lommeli is very common in the Upper Ladinian, but it can also be found in the lowermost Carnian. We assign the Ashin Formation to the Late Ladinian, while the age of the top of the formation remains an open problem.

Sandstone and conglomerate petrography Sandstone petrography was studied in thin sections. In each section, 400 points were counted by the Gazzi–Dickinson method (Ingersoll et al. 1984). Traditional ternary parameters and plots (Dickinson 1985) were supplemented, specifically as lithic grains are concerned, by an extended spectrum of key indices. Metamorphic rock fragments were classified according to both composition and metamorphic rank, mainly inferred from degree of recrystallization of mica flakes. Average rank for each sample was expressed by the ‘metamorphic index’ (MI), which varies from 0 in detritus from

sedimentary and volcanic cover rocks to 500 in detritus from high-grade basement rocks (Garzanti & Vezzoli 2003) (see Table 1).

Alam Formation Mostly upper medium to coarse-grained sandstones of the Alam Formation (10 samples) are volcanic arenites, dominated by microlitic–felsitic lithic fragments and plagioclase (Fig. 21); common quartz, granophyric–granitoid rock fragments, chessboard albite, micas and minor metasedimentary rock fragments occur locally (Q ¼ quartz 25 + 19, F ¼ feldspar (KF ¼ K-feldspar; P ¼ plagioclase) 42 + 20, Lv ¼ volcanic grains 31 + 18, Lm ¼ metamorphic grains 2 + 3; P/F ¼ plagioclase/feldspar 88 + 16). Such composition documents provenance from a magmatic arc dissected to various degrees, with sporadic contribution from metamorphic wallrocks (‘Magmatic Arc Provenance’ of Dickinson 1985; Marsaglia & Ingersoll 1992; Garzanti et al. 2007). The coarse-grained basal layers of the unit (NK169, member A, Nakhlak 1 section) include abundant quartz, chessboard albite and granophyric– granitoid rock fragments, indicating relatively deep erosional levels within the arc massif (‘Dissected Arc’ subprovenance). Slightly higher in the succession (NK170 and NK171, member A, Nakhlak 1 section), overwhelming plagioclase and microlitic–felsitic volcanic grains suggest a major phase of intermediate (latite–andesite?) active volcanism (‘Undissected Arc’ subprovenance). The very-fine-grained sandstones of member C (NK126 and NK3, Nakhlak 2 section) point to an incision phase with subordinate supply from metamorphic wallrocks (quartz, micas, and quartz þ mica lithic fragments). A second volcanic cycle is suggested by the overlying sandstone layers, principally derived from plagioclase-rich tuffs (NK10, member C, Nakhlak 2 section), followed by sandstones (NK43 and NK44: member E, Nakhlak 3 section) dominated by microlitic volcanic rock fragments and plagioclase, derived from relatively mafic products (basaltic andesite?). The unit is capped by sandstones (NK164: top of member E; KN51: member G, Nakhlak 10 section) relatively rich in quartz, chessboard albite, K-feldspar and granophyric–granitoid rock fragments, suggesting a phase of ceased volcanism and renewed erosion.

Ba¯qoroq Formation Fine- to medium-grained sandstones of the Ba¯qoroq Formation (seven samples þ basal sample: Fig. 21) have remarkably homogeneous quartzolithic metamorphiclastic composition (Q 63 + 4, F 14 + 4,

N

GSZ (mm)

GSZ ( f )

Q

KF

Pl

Lvf

Lvm

Ls

Lmlr

Lmf

Lmb

Lu

mica

HM

Total

MI

Ashin Formation (volcanic arenites)

2

504

5

203

Ba¯qoroq Formation (main body)

7

288

1 10

388 367

23 0 67 6 61 4 51 24 17

0 0 7 4 7 5 3 4 4

55 1 9 3 7 5 9 38 22

5 3 1 1 1 1 13 6 4

17 1 1 1 2 1 22 24 21

0 0 0 0 0 0 1 0 0

0 0 0 1 1 1 0 1 1

0 0 10 3 17 3 0 1 1

0 0 0 0 0 0 0 0 0

0 0 0 0 0 0 0 0 0

0 0 4 3 4 1 0 3 3

0 0 0 0 0 0 0 0 0

100.0

Ashin Formation (metamorphiclastic sandstones)

1.0 0.5 2.3 0.8 1.8 0.5 1.4 1.4 1.2

1 2 328 41 338 15 13 26 35

Ba¯qoroq Formation (base) Alam Formation

100.0 100.0 100.0 100.0

N, number of counted samples; GSZ, grain size; Q, quartz; KF, K-feldspar; PI, plagioclase; L, aphanitic lithic grains (Lvf, felsic volcanic; Lvm, microlithic volcanic; Ls, sedimentary; Lmlr, low-rank metamorphic; Lmf, felsic metamorphic; Lmb, metabasite; Lu, ultramafic); HM, heavy minerals; MI, metamorphic index. Mean and standard deviation are given for each stratigraphic interval.

THE TRIASSIC SUCCESSION OF NAKHLAK

Table 1. Detrital modes of Nakhlak sandstones. In each of 25 very-fine- to coarse-grained sandstone samples, 400 – 450 points were counted by Irene Bollati at the Milano-Bicocca University according to the Gazzi –Dickinson method (Ingersoll et al. 1984). A detailed classification scheme allowed us to collect quantitative information on metamorphic rank of rock fragments (MI; Garzanti & Vezzoli 2003). Mean grain size of studied samples was determined in thin section by comparison with standards. Compositional parameters are defined in Garzanti et al. (2007)

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Fig. 21. Conglomerate and sandstone petrography of the Ba¯qoroq Formation. (a) Epidote amphibolite pebble of the Ba¯qoroq Formation (KN35). The light-green– bluish-green fine-grained amphiboles individuate a pervasive foliation. Rutile/ilmenite and plagioclase are also visible. (b) Slightly deformed granite pebble. White mica flakes and quartz display undulose extinction (KN39). (c) Oolitic limestone pebble from the lower part of the formation. (d) Deeply altered pebble made of acidic volcanic with a porphyritic texture. Quartz and plagioclase phenocrysts, almost completely saussuritized, are visible. Sandstone petrography. (e) Feldspatholithic sandstone of the Alam Formation (NK171) derived from an undissected volcanic arc. (f) Quartzolithic micaceous sandstone of the Ba¯qoroq Formation (KN32) derived from an axial metamorphic belt. Both photographs were taken with crossed polars. The bar scale for (e) and (f) is 100 mm.

Lv 4 + 1, Lm 19 + 4; P/F 51 + 29), contrasting sharply with that of underlying Alam sandstones. Provenance from a metamorphic complex is indicated (‘Recycled Orogen Provenance’ of Dickinson 1985; ‘Axial Belt Provenance’ of Garzanti et al. 2007). Only the basal layer (KN49) contains

common volcanic rock fragments, suggesting either terminal volcanic activity or erosion of older volcanic and volcaniclastic rocks. Sandstones in the main body of the unit (KN45, KN43, KN36, KN34, KN33, KN32 and KN31) monotonously consist of quartz, metamorphic rock fragments,

THE TRIASSIC SUCCESSION OF NAKHLAK

feldspars, granitoid rock fragments and micas, in that order of abundance. A few volcanic– subvolcanic rock fragments are invariably present. Dominant medium- to high-rank schistose– gneissic rock fragments (MI 338 + 15; Garzanti & Vezzoli 2003) indicate provenance from low-grade metasedimentary rocks.

Ashin Formation Sandstones of the Ashin Formation have bimodal compositions (Fig. 21), mainly including very-fine to medium-grained quartzolithic metamorphiclastic sandstones such as those of the Ba¯qoroq Formation (KN173, KN174, KN28, KN175 and KN26; Q 70 + 8, F 17 + 7, Lv 2 + 2, Lm 11 + 3; P/F 62 + 16; MI 328 + 41), intercalated with upper medium- to coarse-grained volcanic arenites such as those of the Alam Formation (KN176, KN56; Q 23 + 0, F 55 + 1, Lv 22 + 2, Lm 0 + 0; P/F 100 + 0). Volcanic arenites consist of virtually pure volcanic detritus, including monocrystalline quartz with straight extinction and embayed or bipyramidal outlines, and microlitic –vitric and pyroclastic grains, testifying to a phase of renewed intermediate –felsic (dacite–rhyodacite?) explosive volcanism. Absence of sediments with mixed volcanic –metamorphic signatures (volcanic arenites invariably contain trivial amounts of metamorphic grains, and quartzolithic sandstones invariably include trivial amounts of volcanic detritus) indicate two sharply separate sources at the two opposite sides of the sedimentary basin.

Conglomerate analysis Ten sites for conglomerate analysis were selected in the field within the Ba¯qoroq Formation, and some additional sites were selected within the uppermost Alam Formation (two sites) and lower Ashin Formation (three sites). The composition of at least 100 pebbles was identified at each site, according to a classification of the pebbles in 10 classes, simplified later to six classes (Fig. 21). Furthermore, in order to verify the pebble description, in these sites we sampled pebbles of all the lithologies for thin section analysis (Fig. 20). The following remarks can be pointed out, summarizing the data (Fig. 21): † the conglomerates of the uppermost Alam Formation are composed only of clasts of sedimentary rocks; † sedimentary rocks are also an important component of the conglomerates of the lower part of the Ba¯qoroq Formation (sites BQ10–BQ7), but, from the base of the formation, the quartz pebbles become increasingly very common;

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† the most frequent type of sedimentary rocks in the sites BQ10– BQ7 are oolitic grainstones (Fig. 20c). Note that oolitic limestones are documented only within the member B ‘oolitic limestone’ of the Alam Formation (38 m thick at the Nakhlak 1 section). No oolitic limestones are reported from the nearby outcrops of Ordovician –Permian successions of sedimentary rocks east of Anarak (Sharkovski et al. 1984; Mattei et al. 2004); † within the Ba¯qoroq Formation, quartz is accompanied by metamorphic rocks and by granitoids; † ophiolitic fragments, quoted by Alavi et al. (1997), have not actually been recognized; † volcanoclastic –volcanic rocks are not documented within the conglomerates of the Ashin Formation.

Sedimentary evolution The stratigraphic evolution of the Nakhlak Group was previously described by Davoudzadeh & Seyed-Emami (1972), and more recently by Alavi et al. (1997) and Seyed-Emami (2003). These authors subdivided the evolution of the entire succession into three different stages, corresponding to the three lithostratigraphic units: Alam, Ba¯qoroq and Ashin formations. Alavi et al. (1997) described a first shallowing- and coarsening-upwards evolution in the Alam Formation, a major unconformity at the base of the fining-upwards Ba¯qoroq Formation, and another unconformity at the base of the Ashin Formation that was referred to as a turbiditic unit. Alavi et al. (1997) interpreted the Nakhlak succession as arc-related, and deposited between the trench and the trench–slope break. The subdivision of the Nakhlak Group into three major depositional cycles (lower marine, intermediate continental and upper marine) is confirmed by our data. The first cycle is documented by the Alam Formation (Olenekian–Anisian), while the third cycle, preserved only in the transgressive part, is represented by the Ashin Formation (Upper Ladinian). The intermediate continental episode is recorded by the Ba¯qoroq Formation (Lower Ladinian). We provide a more detailed and constrained reconstruction of the evolution of the succession that documents a more complex interaction between clastic –volcanoclastic supply and carbonate sedimentation, between regional subsidence and uplift, and between relative sealevel changes and syn-sedimentary tectonic and volcanic activity. In particular the evolution of the Olenekian –Anisian part of the succession (Alam Formation) notably differs from the interpretation available in the literature.

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The Alam Formation represents a general transgressive–regressive cycle containing short cycles of variation of siliciclastic and volcaniclastic supply. The base of the succession is not preserved, as it tectonically lies on metagabbros. The transgressive part of the lower cycle is documented by the interval from member A to member D (‘nodular limestone’), while the general regression is represented from member E to member G (‘grey shales’). The older unit (member A, ?lower Olenekian) was deposited in a coastal marine –transitional environment with input of fine-grained volcaniclastics with dissected –transitional arc provenance, probably under arid conditions (presence of dolostones). Member B (‘oolitic limestone’; about 38 m; ?lower Olenekian), deposited in an highenergy shallow-marine environment, documents a rapid transgressive trend with a reduction of volcaniclastic supply. The transition from member A to member B is sharp, as is the boundary between members B and C. The transgressive trend continues after the deposition of member B, and shows an increase in volcaniclastic supply that consists of green volcaniclastic sandstones, hybrid arenites and green shales of the basal part of member C (27 m; Olenekian). This unit was deposited under the influence of currents (cross-laminations) and was probably above the storm wave base. The abundance of shales and siltstones in the upper part of member C (‘varicoloured shales’, up to 127 m; middle–upper Olenekian) represents the deposition in slightly deeper water, probably below the storm wave base. Provenance analysis of sandstones from member C documents an increase of metamorphic content that reflects uplift and erosion of the nearby volcanic arc coeval with the deepening of the basin. The end of the uplift of the arc is marked by the decreased clastic input that preceeded the deposition of the overlying member D (‘nodular limestone’, about 80 m thick, upper Olenekian). This member was deposited in a deep-water environment under relative tectonic quiescence, with sparse input of volcaniclastic sandstones. The low-energy environment, the abundance of ammonoids and the typical nodular facies prove the maximum deepening of the basin. The transition from member D to the overlying thick member E (about 330 m; upper Olenekianlower Bithynian) records a renewed, abundant volcaniclastic input and syn-depositional tectonic activity leading to the deposition of the thick paraconglomeratic beds. The abundance of extrabasinal input and the thick paraconglomerates suggest an increase of sedimentation rate with a strong dilution of carbonate input by volcaniclastic sediments. The general trend of member E is regressive, with gradual reduction of the accommodation space

due to the infilling of the basin. Sea bottom was surely above the storm wave base and the reduction of bathymetry, coupled with the decreased volcaniclastic supply, favoured the onset of the carbonate mounds of member F (up to 50 m; Bithynian). The carbonate production of the mounds influenced sedimentation in the surrounding basinal areas, with deposition of an alternation of resedimented bioclastic limestones and fine-grained sandstones. The carbonate production suddenly ended in the late Bithynian. The driving mechanism for this crisis of the carbonate factory can be deduced from some observations. The boundary between the bioclastic limestones of member F and the tuffs and marls with pelagic ammonoids of the basal part of member G is very sharp and marked by a hardground at the top of the carbonate mounds. Transitional facies such as nodular – pseudonodular limestones with brachiopods in between members F and G are completely absent, indicating that the carbonate production ceased very rapidly. The occurrence of the hardground on top of the carbonate mounds indicates that the crisis of the carbonate factory cannot be ascribed to the siliciclastic input. Therefore, the most probable cause of the end of the carbonate production can be due to a very rapid sea-level rise (drowning unconformity). Considering the environmental setting, the evidence of volcanic activity after the drowning and the rapidity of the environmental change, we suggest that the drowning of the mound could very probably be tectonically controlled, before the deposition of the basal tuff of member G. Member G (‘grey shales’, upper Bithynian – Pelsonian or Illyrian) is characterized by a coarsening-upwards trend, coupled with a general regressive evolution. The maximum depth of the basin is documented by the leiostraca ammonoidbearing marls immediately overlying the top of the carbonate mounds (member F). The sandstones and conglomerate intervals of the middle and upper part of the members reflect the progradation of coastal to alluvial facies intercalated with freshwater –brackish water lagoonal facies with oligotypic Unionites fauna. The Ba¯qoroq Formation (up to 870 m thick; ?Upper Anisian– Ladinian), mostly consisting of red conglomerates and sandstones, is referred to a fluvial environment in a semi-arid climate. The unit continues the continental evolution of member G of the Alam Formation and shows a general fining-upwards trend with a reduction in coarse-grained massive conglomerates and an increase in sandstones. The fining-upwards trend and the clastic input of the Ba¯qoroq Formation are indicative of a tectonic uplift in the surroundings of the Nakhlak Basin. The uplift started during the

THE TRIASSIC SUCCESSION OF NAKHLAK

deposition of member G, as indicated by the petrographic composition of sandstones and conglomerates, which suggest erosion of the volcanic arc during an interval of scarce or absent volcanic activity. Sedimentary rock pebbles are abundant in the conglomerates of member G of the Alam Formation, while their frequency considerably decreases in conglomerates of the lower part of the Ba¯qoroq Formation (Fig. 21: samples BQ10–BQ7). The sedimentary pebbles disappear between sample BQ6 and BQ5, reflecting the end of the erosion of the sedimentary cover. The erosion of the volcanic rocks was, in part, coeval with the erosion of the sedimentary rocks. The frequency of volcanic pebbles is relatively high in samples from BQ10 to BQ7, then they disappear in the middle part of the formation, but slowly become more frequent in the upper part of the unit. The exhumation of the metamorphic basement, documented by quartz, granitoids and metamorphic rocks, started at the base of the Ba¯qoroq Formation (Upper Anisian or Lower Ladinian) and continued in the lower part of the Ashin Formation (Lower Ladinian). In the Late Ladinian a new tectonic event led to the very fast transgression of the Ashin Formation. The quartz-rich conglomeratic beds at the base of the unit were deposited in a fluvial–transitional environment, which became open marine in a few tens of metres, as documented by the occurrence of the ammonoid Romanites simionescui reported by Vaziri (1996). This fast environmental change is accompanied by a reprisal of volcanic activity documented by tuff layers up to a few metres thick, and continues with the deposition of the fine-grained turbiditic sandstones of the middle and upper part of the Ashin Formation. This part of the formation is interpreted to be deposited at a water depth of several hundreds of metres not only because of its sedimentological features, but also because of the rich deepwater ichnofacies (Vaziri & Fu¨rsich 2007). The petrographic analyses of sandstones fully support this picture. A Ba¯qoroq-like clastic supply from the metamorphic and intrusive rocks (quartzolithic metamorphiclastic sandstones), and a new source from volcanic centres (volcanic arenites), are emphasized. The latter type of arenite especially documents the presence of volcanoes close to the study area. The transition from continental conglomerates of the Ba¯qoroq Formation to the deep-water turbiditic facies of the Ashin Formation occurs in about 100 m of stratigraphic thickness. The bathymetric change documented in such a relatively thin thickness shows a relative sea-level rise of several hundreds of metres that cannot be explained by eustasy alone. Eustasy might have provided an additional contribution to the facies change, as the

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Ladinian was characterized by an eustatic sea-level rise (Haq et al. 1988), but most of the relative change must be ascribed to a rapid increase in subsidence. The reprisal of volcanic activity, documented by tuffs and by sandstone petrography, is also consistent with a sudden increase in subsidence of the area during the Late Ladinian.

The significance of the Nakhlak Group The new stratigraphically well-constrained data have a good potential for improving the knowledge of the geological and geodynamic setting of the Nakhlak area during the Triassic, although a new geodynamic model for Central Iran cannot rely only on these data. The relationship of the Triassic of Nakhlak with the Anarak metamorphic range, and the discussion of Bagheri & Stampfli’s (2008) model for the evolution of Central Iran, are treated in a separate contribution (Zanchi et al. 2009). The discussion of the 1358 counterclockwise rotation of Central Iran from a palaeogmagnetic viewpoint is treated by Muttoni et al. (2009).

Geodynamic setting The available stratigraphic, geological and petrographic data from the Triassic succession of Nakhlak support the following considerations: † a 2.4 km-thick succession was deposited in an irregularly subsiding sedimentary basin during Early– Middle Triassic times; † volcaniclastic and siliciclastic supply in the basin was high from Olenekian to Late Ladinian, with a few carbonate-dominated episodes restricted to short periods of reduced clastic supply; † sedimentation in the basin was strongly influenced by active syn-sedimentary tectonics documented by fast changes of the subsidence (i.e. the ‘drowning’ of the Ba¯qoroq Formation), as well as by sudden changes in facies (i.e. the drowning of member F of the Alam Formation) and by the debris flow at the base of member E of the Alam Formation; † the basin was close to an active volcanic area, as documented by tuffaceous intervals in the succession from the Olenekian to the Late Ladinian; † provenance analyses on Alam sandstones and some Ashin sandstones indicate supply from a magmatic arc; † magmatism occurred in several distinct pulses, and was characterized by orogenic calc-alkaline affinity including latite –andesite (lower Alam Formation), basaltic andesite (upper Alam Formation) and rhyodacite (Ashin Formation); † two different sources for volcanic and metamorphic detritus co-existed at Late Ladinian times (Ashin Formation);

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† the occurrence of ophiolitic rock fragments (serpentinites) in the Ba¯qoroq Formation mentioned by Alavi et al. (1997) is not confirmed. No fragments of serpentinites were recognized during the statistic counting of 400 grains per seven thin sections of sandstones and more than 100 grains of conglomerates per 10 sites; † the Triassic succession was deformed during the Cimmerian orogeny before the Late Cretaceous. The study area was not affected by postCretaceous thrusts; thus, there is no evidence for post-Cimmerian tectonics as reported by Alavi et al. (1997). These data suggest the location of the Nakhlak area along an active margin during the Triassic. The absence of proximal volcanic deposits excludes the possibility that deposition occurred in an intra-arc basin and strongly supports the inference that the Nakhlak succession represents a portion of a forearc basin related to a subduction zone that was active during the Early –Middle Triassic. The geology of the Anarak area, south of Nakhlak, is consistent with this reconstruction, and the Anarak Metamorphic Complex might represent the remnant of an accretionary wedge (Bagheri & Stampfli 2008; see Zanchi et al. 2009 for discussion).

Nakhlak with possible primary magnetization show no rotation (site IR02) or only moderate counterclockwise rotation (site IR04) since the Olenekian. In the light of these new data a more careful comparison of the Nakhlak and Aghdarband successions is warranted. Here we briefly review not only the similarities, but also the differences, between the two successions. For Aghdarband we used data from the literature and new data collected during two field seasons in the area (2005 and 2007).

General setting Even if the geodynamic position of Nakhlak is uncertain, the location of Aghdarband on the southern margin of the Laurasia is well accepted. This palaeoposition relies on the identification of an ophiolitic belt (Eftekharnezhad & Behroozi 1991) and an accretionary wedge (Alavi 1991; Ruttner 1993; Alavi et al. 1997) to the south, both related to the subduction of the Palaeotethys, as well as on the basis of faunal affinity of the Devonian–Carboniferous fauna recorded in units underlying the Triassic succession (Wendt et al. 2005).

Stratigraphic succession and correlation Comparison of Nakhlak and Aghdarband Since its first description (Davouzadeh & Seyed-Emami 1972) the Triassic Nakhlak Group has been compared and correlated with the Triassic succession exposed in the erosional window of Aghdarband (Ruttner 1984, 1991, 1993; Alavi et al. 1997), in the Koppeh Dag. These comparisons attempted to explain the striking difference in the Triassic of Nakhlak compared with the Triassic of the surrounding areas of Alborz, Central Iran and Sanandaj–Sirjan (see Seyed-Emami 2003 for a summary). The stratigraphic similarities between the two areas provided support to the model of 1358 counterclockwise rotation of Central Iran since the Triassic (Davoudzadeh et al. 1981; Davoudzadeh & Weber-Diefenbach 1987; Soffel et al. 1996). However, the lithological correlations emphasized the similarities and ignored the differences between Nakhlak and Aghdarband. Recent investigations on the distribution of Devonian–Lower Carboniferous facies in central and northern Iran question the 1358 counterclockwise rotation. Wendt et al. (2005, pp. 77–78) stressed a facies distribution that is almost consistent with the present-day position of Central Iran with respect to the Koppeh Dag. Muttoni et al. (2009) also advise caution in considering the counterclockwise model because the very few samples from the Triassic of

The Triassic of Aghdarband is known from several contributions (Ruttner 1984, 1991, 1993; Baud & Stampfli 1989; Baud et al. 1991a, b; Krystyn & Tatzreiter 1991; Alavi et al. 1997). The succession is illustrated in Figure 22, which is mostly based on Ruttner (1991, fig. 4). The succession shows some similarities with Nakhlak, as well as some differences. Only the lower –middle part of the succession, below the Cimmerian unconformity (Qara Geithan –Sina formations), has time equivalents in the Nakhlak section. Within this interval there is an important unconformity at the base of the Sina Formation that was not well detected by all the authors. This unconformity was first recognized by Ruttner (1984, 1991, 1993) and is confirmed by our investigations, while Baud et al. (1991b, fig. 6) identified an unconformity in a lower stratigraphic position, between the Sefid Kuh Limestone and the overlying ‘fossil horizon 1’, i.e. the Nazarkardeh Formation. The lower parts of the Aghdarband and Nakhlak successions are different in six major ways: † The 500 m-thick Qara Geithan Formation (Lower Triassic: Ruttner 1991 and Baud et al. 1991b; Upper Permian: Alavi et al. 1997), consisting of alluvial sandstones and conglomerates rich in granitoids, has no counterpart in the Nakhlak Group.

THE TRIASSIC SUCCESSION OF NAKHLAK

Fig. 22. Comparison of the Nakhlak Group (left) with the the Triassic succession exposed in the erosional window of Aghdarband (Koppeh Dag, NE Iran). The Aghdarband log is composite and based on Ruttner’s (1991) tectonic slice I and II, with projection of the Anabeh Conglomerate from tectonic slice III. The calibration of the succession is based on ammonoid and conodont data from Baud et al. (1991a, b), Krystyn & Tatzreiter (1991) and a new Olenekian ammonoid fauna. Dating of the Miankuhi Formation is based on plant megafossils (Boersma & van Konijnenburg-van Cittert 1991). Abbreviations: NF, Nazarkardeh Formation; FM, Faqir Marl Bed.

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† The overlying Sefid Kuh Limestone, a 200 m-thick Olenekian (Spathian: Baud et al. 1991a) carbonate ramp, shows a different evolution to that of the Alam Formation. The Sefid Kuh Limestone is dated with conodonts as middle –upper Olenekian (Spathian), and it is coeval with members C, D and with the lower part of member E of the Alam Formation. During this interval the general trend of the Nakhlak Basin was first deepening (members C and D), after which the basin was affected by tectonic instability and started a shallowing trend (lower part of member E: Fig. 21). The sudden drowning of the Sefid Kuh Limestone is Aegean in age (Baud et al. 1991a), coeval with a relatively monotonous and shallowing part of member E of the Alam Formation (middle part of the Nakhlak 3 section). † The development of the carbonate mounds of member F of the Alam Formation (Bithynian) at Nakhlak is coeval with the deposition of the basinal Nazarkardeh Formation. The latter is very similar to the upper part of member E of the Alam Formation, but the late Bithynian – Late Anisian evolution of the two areas is very different. The tectonically controlled drowning of member F of the Alam Formation is late Bithynian in age and is followed by a rather regular shallowing trend from marine –transitional– continental fluvial facies (member G ‘grey shales’) up to the base of the Ba¯qoroq Formation. † The Ba¯qoroq Formation, which documents the uplift and erosion of a metamorphic basement, is dated only by its stratigraphic position as ?Late Anisian –Ladinian in age at Nakhlak. However, some Late Anisian –Early Ladinian conodonts were reported from the basal part of the Sina Formation at Aghdarband (Ruttner 1991, p. 32), which demonstrates that between the Bithynian and the Late Anisian –Early Ladinian the Aghdarband succession was first uplifted and eroded, and then the basin suddenly subsided in the Late Anisian –Early Ladinian to accumulate marine sediments. This trend is reversed with respect to the sedimentary evolution of the Nakhlak Basin. † That rapid subsidence at Aghdarband leads to the deposition of the turbiditic volcanoclastic sandstones of the Sandstone Member of the Sina Formation. This member is almost coeval with the Ba¯qoroq Formation but the depositional setting is very different, the Ba¯qoroq Formation being deposited in a continental setting. † From a petrographic point of view, the two units are slightly different. The Sina Formation mostly consists of volcanoclastic sandstones with a minor amount of metamorphic fragments,

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while the metamorphic fragments are much more abundant in the Ba¯qoroq Formation. In contrast to differences in their early history, during the Late Ladinian both the Aghdarband and the Nakhlak successions show a similar deepening trend. At Nakhlak the deepening trend is documented by the transition of the Ba¯qoroq and Ashin formations, which was mainly controlled by tectonics with a possible additional eustatic component. At Aghdarband an almost coeval deepening is represented by the transition from the Sandstone to Shale members of the Sina Formation. The first member is referred to a sandstonedominated deep-marine environment, while the Shale Member represents a shale-dominated turbiditic system (Baud et al. 1991b).

Olenekian and Bithynian. In the Aghdarband succession no major unconformity seems recorded in this interval, but the basal unconformity of the Sina Formation is slightly younger. The easiest solution could be to locate Nakhlak on an active margin of a microplate approaching Aghdarband (i.e. the Turan Plate) during the Early Triassic. However, this model cannot be supported by only ammonoid data and would require a solution to the doubts on the 1358 counterclockwise rotation of Central Iran. What can be concluded at the present is that the Alavi et al. (1997) location of Nakhlak and Aghdarband on the same arc– trench system is not consistent with Triassic ammonoid palaeobiogeography.

Faunal similarity

Conclusions

Three ammonoid-bearing intervals can be recognized in the Aghdarband succession (Fig. 22). The oldest interval has recently been found in the Sefid Kuh Limestone and is Olenekian in age (Balini et al. in prep.). The second is Bithynian in age, while the third is Late Ladinian in age (Fossil horizon 1 and 2 of Krystyn & Tatzreiter 1991 and Ruttner 1991). Two Bithynian ammonoid zones are recorded in the Nazarkardeh Formation: the early Bithynian Nicomedites osmani Zone and the late Bithynian Aghdarbandites ismidicus Zone. The Late Ladinian ammonoids are concentrated in the Faqir Marl Bed, i.e. at the base of the Shale Member of the Sina Formation. The occurrence in both the Nakhlak and the Aghdarband successions of three coeval ammonoid faunas is unusual, but faunal composition is even more interesting. The Olenekian ammonoid fauna of member C of the Alam Formation is dominated by Albanites and Columbitidae, showing typical Tethyan affinity with Kc¸ira (Albania: Arthaber 1908, 1911; Germani 1997) and Chios (Greece: Renz & Renz 1948; Mertmann & Jacobshagen 2003). The few Olenekian ammonoids of the middle Sefid Kuh Formation consist of tirolitids of the group of T. rossicus Kiparisova (Balini et al. in prep.). This group of tirolitids is known only from Mangyshlak (Shevyrev 1968; Balini et al. 2000) and suggests a pericaspian affinity for the Olenekian of Aghdarband. This notably different faunal composition of the two sites does not persist in the Bithynian and the Late Ladinian, which are characterized by almost identical ammonoid fauna at Nakhlak and Aghdarband. The understanding of the palaeogeographic and palaeogeodynamic significance of the Triassic faunal affinities requires further investigation. The change in faunal affinity took place between the

The stratigraphic revision of the Triassic succession of Nakhlak based on the analyses of lithofacies, ammonoids, conodonts, bivalves, petrography of sandstones and conglomerates lead to the following results: † the stratigraphic evolution of the succession consisting of the Alam (Olenekian–Anisian), Ba¯qoroq (?Upper Anisian –Ladinian) and Ashin formations (Upper Ladinian) has been documented. The 2.4 km-thick succession is divided into two sedimentary cycles. Dominant magmatic arc provenance for Alam sandstones and some Ashin sandstones indicate deposition in an arc-related setting. Magmatism occurred in several distinct pulses, and was characterized by orogenic calc-alkaline affinity; † the sandstone and conglomerate petrography of the Nakhlak succession records two intervals (Olenekian and Early–Late Ladinan) of important metamorphic supply. In particular, the Ba¯qoroq Formation documents the exhumation of a metamorphic basement, as well as the erosion of a volcanic arc. The sedimentary succession around the base of the Ba¯qoroq Formation does not show any indication of deformation that would be necessary to provide a final demonstration of a collision; † the succession was deposited along an active margin, most probably in a forearc setting; † facies comparison demonstrates that the Aghdarband (NE Iran, Koppeh Dag) and Nakhlak successions were both deposited on an active margin, but lithological constrains are not sufficient to demonstrate their possible proximity. The interpretation of similarities and differences between the two localities is not unequivocal, especially with regard to the relative palaeogeodynamic position of the two areas;

THE TRIASSIC SUCCESSION OF NAKHLAK

† Nakhlak and Aghdarband have different ammonoid faunal affinities during the Early Triassic, and similar affinities from the Bithynian to the Late Ladinian; † Triassic ammonoid palaeobiogeographical data argue against the location of Nakhlak close to the active margin of Aghdarband, at least during the Early Triassic, and do not support the model suggested by Alavi et al. (1997, fig. 10), who placed Nakhlak and Aghdarband, respectively, towards the arc and towards the trench of the same trench –slope break. Fieldwork (2003: M. Balini, A. Nicora and R. Salamati; 2004: M. Balini, F. Berra, G. Muttoni, M. Levera, M. Mattei, A. Zarchi and F. Mossavvari) was carried out in the structure of two MEBE projects: ‘Stratigraphy of Selected Permian and Triassic Sections in Iran’ (leader M. Gaetani) and ‘Tectonic Evolution of the Yazd, Tabas and Lut Blocks (Central Iran) by Means of Palaeomagnetic, Structural and Stratigraphic Data’ (leader M. Mattei). We are deeply indebted to the Geological Survey of Iran and especially with M. Ghassemi, for the logistic and technical support. We would like to thank the GSI drivers, and, in particular Mr Takshin for the extremely efficient help and co-operation. The administration and the miners of the Nakhlak Mining Company and village showed us extreme kindness and hospitality. This manuscript has been improved by stimulating suggestions of the reviewers, A. Robertson and A. Baud. Warm thanks also to the editorial board, and especially to M.-F. Brunet, for support and stimulation. Very special thanks to F. S. Aghib, for her patience throughout the days, nights and weekends required for writing this paper.

Appendix 1 GPS co-ordinates (WGS84 system) of the 13 statistic sites for petrographic analysis of conglomerates within the Ba¯qoroq and Ashin formations. The two sites in the Alam Formation (AL1 and AL2) are located in the Nakhlak 6 section (Fig. 19). The sites AS1 –AS3 and BQ1 are located at very close distance, in the same stratigraphic section.

AS1–AS3 BQ1 BQ2 BQ3 BQ4 BQ5 BQ6 BQ7 BQ8 BQ9 BQ10

338330 2600 same as AS1 338330 2700 338330 1200 338330 0800 338330 0000 338320 5500 338320 5200 338320 4700 338320 4200 338320 4000

538500 0200 538490 5100 538500 0600 538500 0400 538500 0400 538500 0800 538500 0900 538500 1100 538500 0900 538500 1300

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L EHRMANN , D. J. 1999. Early Triassic calcimicrobial mounds and biostromes of the Nanpanjiang basin, south China. Geology, 27, 359–362. M ARSAGLIA , K. M. & I NGERSOLL , R. V. 1992. Compositional trends in arc-related, deep-marine sand and sandstone: a reassessment of magmatic-arc provenance. Geological Society of America Bulletin, 104, 1637– 1649. M ATTEI , M., B ALINI , M., B ERRA , F., M UTTONI , G. & Z ANCHI , A. 2004. Geodynamic evolution of the Yazd, Tabas and Lut blocks (Central Iran) by integration of paleomagnetic, structural and stratigraphic data. MEBE Report, 1 –18. M ERTMANN , D. & J ACOBSHAGEN , V. 2003. Upper Olenekian (Spathian) ammonoids from Chios (Lower Triassic, Greece): taxonomy and stratigraphic position. Rivista Italiana di Paleontologia e Stratigrafia, 109, 417– 447. M UTTONI , G., M ATTEI , M., B ALINI , M., Z ANCHI , A., G AETANI , M. & B ERRA , F. 2009. The drift history of Iran from the Ordovician to the Triassic. In: B RUNET , M.-F., W ILMSEN , M. & G RANATH , J. W. (eds) South Caspian to Central Iran Basins. Geological Society, London, Special Publications, 312, 7 –29. O VTCHAROVA , M., B UCHER , H., S CHALTEGGER , U., G ALFETTI , T., B RAYARD , A. & G UEX , J. 2006. New Early to Middle Triassic U–Pb ages from South China: Calibration with ammonoid biochronozones and implications for the timing of the Triassic biotic recovery. Earth and Planetary Science Letters, 243, 463– 475. R ENZ , C. & R ENZ , O. 1948. Eine untertriadische Ammonitenfauna von der griechischen Insel Chios. Schweizerische Pala¨ontologische Abhandlungen, 66, 3– 98. R UTTNER , A. W. 1984. The pre-Liassic basement of the eastern Kopet Dag Range. Neues Jahrbuch Geologie und Pala¨ontologie Abhandlungen, 168, 256–268. R UTTNER , A. W. 1991. Geology of the Aghdarband Area (Kopet Dag, NE-Iran). Abhandlungen der Geolologischen Bundesanstalt, 38, 7 –79. R UTTNER , A. W. 1993. Southern borderland of Triassic Laurasia in north-east Iran. Geologische Rundschau, 82, 110–120. S AIDI , A., B RUNET , M.-F. & R ICOU , L.-E. 1997. Continental accretion of the Iran Block to Eurasia as seen from Late Paleozoic to Early Cretaceous subsidence curves. Geodinamica Acta, 10, 198 –208. S ALVADOR , A. 1994. International Stratigraphic Guide, 2nd edn. Geological Society of America, Boulder, CO, 1 –214. S ENGO¨ R , A. M. C. 1984. The Cimmeride Orogenic System and the Tectonics of Eurasia. Geological Society of America, Special Paper, 195. S EYED -E MAMI , K. 2003. Triassic of Iran. Facies, 48, 91–106. S HARKOVSKI , M., S USOV , M. & K RIVYAKIN , B. 1984. Geology of the Anarak Area (Central Iran). Explanatory Text of the Anarak Quadrangle Map 1:250 000. Geological Survey of Iran, Tehran, 1 –143. S HEVYREV , A. A. 1968. Triassic ammonoids of southern U.S.S.R. Transactions of the Paleontological Institute, U.S.S.R. Academy of Sciences, 119, 1– 272 (in Russian).

THE TRIASSIC SUCCESSION OF NAKHLAK S OFFEL , H. C., D AVOUDZADEH , M., R OLF , C. & S CHMIDT , S. 1996. New palaeomagnetic data from Central Iran and a Triassic palaeoreconstruction. Geologische Rundschau, 85, 293– 302. S TO¨ CKLIN , J., E FTEKHAR -N EZHAD , J. & H USHMAND ZADEH , A. 1965. Geology of the Shotori Range (Tabas area, East Iran). Geological Survey of Iran Report, 3, 1– 69. T OZER , E. T. 1972. Triassic ammonoids and Daonella from the Nakhlak Group, Anarak region, Central Iran. Geological Survey of Iran Report, 28, 29–69. T OZER , E. T. 1981. Triassic Ammonoidea: Geographic and Stratigraphic Distribution. In: H OUSE , M. R. & S ENIOR , J. R. (eds) The Ammonoidea. Systematic Association, Special Volume, 18, 397–432. V AZIRI , H. 1996. Lithostratigraphy, biostratigraphy, and sedimentary environments of the Triassic rocks of the Nakhlak area in Central Iran. PhD thesis, Azad University, Iran. V AZIRI , H. 2001. The Triassic Nakhlak Group, an exotic ¨ . T., succession in Central Iran. In: AKINCI , O GO¨ RMU¨ S¸ , M., KUS¸ C¸ U , M., KARAGU¨ ZEL , R. & BOZCU , M. (eds) Proceedings of the 4th International Symposium on Eastern Mediterranean Geology, Isparta, Turkey, 53–68. V AZIRI , S. H. & F U¨ RSICH , F. T. 2007. Middle to Upper Triassic deep-water trace fossils from the Ashin

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An overview of the stratigraphy and facies development of the Jurassic System on the Tabas Block, east-central Iran ¨ RSICH1, KAZEM SEYED-EMAMI3 & MARKUS WILMSEN1,2*, FRANZ THEODOR FU MAHMOUD REZA MAJIDIFARD4 1

GeoZentrum Nordbayern der Universita¨t Erlangen-Nu¨rnberg, Fachgruppe Pala¨oUmwelt, Loewenichstrasse 28, D-91054 Erlangen, Germany

2

Senckenberg Naturhistorische Sammlungen Dresden, Museum fu¨r Mineralogie und Geologie, Sektion Pala¨ozoologie, Ko¨nigsbru¨cker Landstrasse 159, D-01109 Dresden, Germany 3

School of Mining Engineering, University College of Engineering, University of Tehran, P.O. Box 11365-4563, Tehran, Iran 4

Geological Survey of Iran, Box 131851-1494, Tehran, Iran

*Corresponding author (e-mail: [email protected]) Abstract: The Tabas Block of east-central Iran shows very thick and well-exposed Upper Triassic–Jurassic sequences, which are crucial for the understanding of the Mesozoic evolution of the Iran Plate. The succession is subdivided into major tectonostratigraphic units based on widespread unconformities related to the Cimmerian tectonic events. As elsewhere in Iran, there is a dramatic change from Middle Triassic platform carbonates (Shotori Formation) to the siliciclastic rocks of the Shemshak Group (Norian– Bajocian), reflecting the onset of Eo-Cimmerian deformation in northern Iran. Following the marine sedimentation of the Norian– Rhaetian Nayband Formation, the change to non-marine, coal-bearing siliciclastic rocks (Ab-e-Haji Formation) around the Triassic–Jurassic boundary is related to the main uplift phase of the Cimmerian orogeny. Condensed limestones of the Toarcian– Aalenian Badamu Formation indicate widespread transgression, followed by rapid lateral facies and thickness variations in the succeeding Lower Bajocian Hojedk Formation. This tectonic instability culminated in the middle Bajocian compressional– extensional Mid-Cimmerian event. The resulting Mid-Cimmerian unconformity separates the Shemshak Group from the Upper Bajocian– Upper Jurassic Magu (or Bidou) Group. The succeeding Late Bajocian– Bathonian onlap of the Parvadeh and Baghamshah formations (Baghamshah Subgroup) was caused by increased subsidence of the Tabas Block rather than a eustatic sealevel rise, followed by the development of a large-scale platform– basin carbonate system (Callovian– Kimmeridgian Esfandiar Subgroup). Block faulting starting in the Kimmeridgian (Late Cimmerian event) resulted in the destruction of the carbonate system, which was covered by Kimmeridgian– Tithonian limestone conglomerates, red beds and evaporites (Garedu Subgroup or Ravar Formation). Virtually the same pattern of relative sea-level change, facies development and succession of geodynamic events is recorded from the Late Triassic–Jurassic of northern Iran (Alborz Mountains), suggesting that the Iran Plate behaved as a single structural unit at that time.

Central Iran occupies a key position for unravelling the post-Eo-Cimmerian Mesozoic (Jurassic – Cretaceous) history of the Iran Plate. From the numerous structural units of central Iran, the Tabas Block of east-central Iran (Fig. 1) shows the thickest, most complete and best-exposed sequence of Jurassic rocks of the region. All of the Lower and most of the Middle Jurassic are characterized by thick siliciclastic sequences, whereas the Callovian –Upper Jurassic rocks are predominantly calcareous. However, apart from the early mapping surveys (e.g. Huckriede et al. 1962; Sto¨cklin et al. 1965; Ruttner et al. 1968; Aghanabati 1977; Kluyver et al. 1983a, b), no

comprehensive description and documentation of the Jurassic lithostratigraphic units exists. The objective of this study is, therefore, to describe the stratigraphy and facies development of the Jurassic System on the Tabas Block, and to integrate these data into new geodynamic models proposed for the Jurassic tectonic history of the Iran Plate.

Geological setting The study area is located in the central part of the Central-East Iranian Microcontinent (CEIM; Takin 1972) (Fig. 1). The CEIM, together with central Iran and the Alborz Mountains, forms the

From: BRUNET , M.-F., WILMSEN , M. & GRANATH , J. W. (eds) South Caspian to Central Iran Basins. The Geological Society, London, Special Publications, 312, 323–343. DOI: 10.1144/SP312.15 0305-8719/09/$15.00 # The Geological Society of London 2009.

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Fig. 1. Structural and geographic framework of Iran showing the main sutures, structural units and geographic areas.

Iran Plate, which occupies a structural key position in the Middle Eastern Tethysides (Fig. 2) (Sengo¨r et al. 1988). As an element of the Cimmerian microplate assemblage, it became detached from Gondwana during the (Late) Permian and collided with Eurasia (Turan Plate) in the late Middle– early Late Triassic, thereby closing the Palaeotethys (e.g. Berberian & King 1981; Boulin 1988; Sengo¨r et al. 1988; Sengo¨r 1990; Alavi et al. 1997; Saidi et al. 1997; Stampfli & Borel 2002; Fu¨rsich et al. 2009a). The CEIM consists of three north –southoriented structural units, called the Lut, Tabas and Yazd blocks, which are today aligned from east to west, respectively. The lateral relationships of the three blocks during the Jurassic Period, however, are largely unknown owing to potential anticlockwise rotation of the CEIM of about 1358 since the Triassic and the Cretaceous opening and succeeding closure of small oceanic basins around the microcontinent (e.g. Davoudzadeh et al. 1981; Lindenberg et al. 1983; Soffel & Fo¨rster 1984; Soffel et al. 1996; Alavi et al. 1997; Besse et al. 1998). During the Jurassic Period, however, the blocks were most

probably oriented east –west, with the Lut block facing the Neotethys (Fig. 2). The focus of this contribution is the middle of the three blocks, the Tabas Block. Palaeogeographic reconstructions for the Middle–Late Jurassic (e.g. Enay et al. 1993; Thierry 2000) (Fig. 2) place the Iran Plate at the northern margin of the Neotethys at a subtropical palaeo-latitude of approximately 20– 308N. Sedimentological and stratigraphical analyses indicate that the Tabas and Lut blocks were mostly covered by the sea during the Jurassic Period, whereas the Yazd Block remained emergent.

Previous work The reconnaissance of the Tabas Block (Fig. 3) started with mapping surveys in the 1960s and 1970s (e.g. Huckriede et al. 1962; Sto¨cklin et al. 1965; Ruttner et al. 1968; Aghanabati 1977; Kluyver et al. 1983a, b). The thick Upper Triassic –Jurassic succession of this hitherto geologically largely unknown area was subdivided into a number

JURASSIC OF THE TABAS BLOCK, IRAN

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Fig. 2. (A) Palaeogeography and plate tectonic situation of the western Neotethys during the Callovian (modified from Thierry 2000). Map area of B is indicated. (B) Palaeogeographic conclusion for the Iran Plate during the Middle and Late Jurassic with CEIM in a pre-rotational position. (C) North– south cross-section as indicated in B. Abbreviations: GCB, Greater Caucasus Basin; SCB, South Caspian Basin; KD, Koppeh Dagh; Bi, Binalud Mountains; Ab, Alborz; CI, Central Iran; Ya, Yazd Block; Tb, Tabas Block; Lut, Lut Block; CEIM, Central-East Iranian Microcontinent. For further explanations see text.

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Fig. 3. Locality map of the Tabas Block in east-central Iran. Asterisks labelled A– D refer to the approximate positions of lithostratigraphic successions shown in Figure 4. Block-bounding faults are approximately shown.

of formations (see the compilation by Sto¨cklin 1971) that were subsequently combined into two groups: the Upper Triassic– lower Middle Jurassic Shemshak Group and the Middle–Upper Jurassic Magu Group (northern Tabas Block) or Bidou Group (southern Tabas Block), respectively (Fig. 4) (Aghanabati 1977, 1998). Detailed studies of the Jurassic geology, however, were not undertaken until the 1990s.

The biostratigraphic studies of ammonites (Seyed-Emami 1967, 1971, 1988; Seyed-Emami et al. 1991, 1993, 1997, 1998, 2000, 2001, 2002, 2004b; Schairer et al. 2000, 2003) were crucial for dating the lithostratigraphic units and allowed a precise correlation with the European standard biozones. Furthermore, they formed the basis for the timing of geodynamic events recorded in the successions (Seyed-Emami et al. 2004a) and for

Fig. 4. Lithostratigraphy of the Jurassic System on the northern and southern Tabas Block, as well as the western Lut Block, east-central Iran. The key shows the predominant lithologies and characteristic faunal elements of the formations: 1, sandstone; 2, siltstone; 3, clay; 4, limestone; 5, marl; 6, conglomerate; 7, gypsum; 8, volcanics; 9, bioclasts; 10, platform debris; 11, ammonite data; 12, bivalves; 13, hiatus; 14, coal; See Figure 3 to locate the different regions A –D.

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sedimentological studies of individual formations such as the Sikhor Formation (Fu¨rsich et al. 2003a), the Qal-eh Doktar and Esfandiar formations (Fu¨rsich et al. 2003b), and the Korond Formation (Schairer et al. 2003). A detailed lithostratigraphic framework of the Magu Group of the northern Tabas Block was presented by Wilmsen et al. (2003) and later summarized by Seyed-Emami et al. (2006).

Lithostratigraphy and facies development The post-Eo-Cimmerian succession (i.e. postShotori Formation) of the Tabas Block was subdivided by Aghanabati (1977, 1998) into two major units, termed the Shemshak Group (Norian – Bajocian) and the Magu Group (Bajocian –Upper Jurassic), respectively (Fig. 4; on the southern Tabas Block, the Magu Group is termed Bidou Group). The formations of those two groups, their depositional environments and their geodynamic significance are described here, with the exception of those belonging to the Triassic which were the focus of previous work (see Fu¨rsich et al. 2005a). Co-ordinates of mentioned localities are given in the Appendix.

Shemshak Group (Norian – Bajocian) As elsewhere on the Iran Plate, there is a dramatic change from Middle Triassic platform carbonates (Shotori Formation) to siliciclastic rocks of the Shemshak Group (Norian –Bajocian: Seyed-Emami 2003; Fu¨rsich et al. 2005a, 2009a). Norian – Rhaetian sediments are predominantly marine and consist of predominantly fine-grained siliciclastic and calcareous sediments of the Nayband Formation (for a detailed description see Fu¨rsich et al. 2005a). The Nayband sediments are up to 3000 m thick and were deposited in extensional basins separated by horst structures (e.g. Shotori Swell at the eastern margin of the Tabas Block). At the Triassic –Jurassic boundary marine sedimentation was terminated in many places, followed by nondeposition or erosion, or replaced by non-marine sedimentation (Ab-e-Haji Formation; Fig. 4). Ab-e-Haji Formation. The Lower Jurassic – Aalenian Ab-e-Haji Formation overlies the marine Upper Triassic Nayband Formation. In the northern Tabas Block, the boundary is marked by a coarsegrained, quartzose sandstone unit (informally termed ‘Gravellite’ by the geologists of the Tabas Coal Company). The Ab-e-Haji Formation reaches a thickness of up to 500 m, but locally can be reduced to a few tens of metres (e.g. near Parvadeh mine, northern Tabas Block) or be completely

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missing, as in the eastern part of the Tabas Block (Shotori Mountains). The type area is located in the Kalmard area near Kuh-e-Rahdar, northwestern Tabas Block (Aghanabati 1977), where it attains a thickness of merely 86 m. The thickness of 480 m reported by Aghanabati (1977, p. 109ff.) also included the overlying Badamu Formation, which is very thick at this locality, consisting of more than 200 m of intercalated packages of oolitic limestone, sandstone and shale. The Ab-e-Haji Formation consists of greenish sand- and siltstones (Fig. 5A) and contains workable coal seams with proven coal resources of 282 billion (109) tonnes (Bassir 1985). The sandstones are fine- to medium-grained, often sharp-based and show ripple- or trough cross-bedding. Their thickness is in the range of 5–10 m. The siltstones occasionally show ripple lamination, but are mostly structureless and carbonaceous. Ferruginous concretions are common. Occasionally, thin beds of calcareous sandstone with bivalves and trace fossils are intercalated. The Ab-e-Haji Formation was not studied in detail, but appears to be largely non-marine. It was most probably deposited in fluvial– coastal plain environments. Contemporaneous deposits to the east, on the Lut Block, seem to have experienced a more marine influence (based on a brief survey of these deposits near Kuh-e-Shisui; see Seyed-Emami et al. 2004b), suggesting a continental-to-marine gradient from west to east (as also evident for later deposits). The Ab-e-Haji Formation is absent from the central Shotori Mountains (‘Shotori Swell’ of Sto¨cklin et al. 1965) (Figs 3 and 4). Rocks mapped as the Ab-e-Haji Formation in the remaining areas of the Shotori Mountains turned out to belong either to the Nayband Formation or the Hojedk Formation (pers. obs.). In the Zarand area (c. 70 km NW of Kerman; Fig. 3) of the southern Tabas Block, the Hettangian– Pliensbachian siliciclastics attributed to the Ab-e-Haji Formation may reach up to 3000 m in thickness (‘Zarander Trough’ of Huckriede et al. 1962). The contact to the underlying, predominantly marine Nayband Formation is marked by the appearance of coarse, quartzitic sandstones and, in some areas, coal seams (Fu¨rsich et al. 2005a). In terms of depositional setting and chronostratigraphy, the Ab-e-Haji Formation is comparable to the Alasht Formation of the Shemshak Group in the Alborz Mountains of northern Iran (see Fu¨rsich et al. 2009a). Similar to the situation in the Alborz Mountains (Fig. 1), the base of the formation is marked by an increase in grain size (Gravellite in the Tabas area, coarse arkosic sandstones of the lower Alasht Formation in the Alborz Mountains). In the southern Tabas Block too, fine-grained marine siliciclastics and

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Fig. 5. (A) Limestones of the Badamu Formation overlying greenish siltstones and sandstones of the Ab-e-Haji Formation, approximately 60 km west of Tabas north of the road to Yazd. (B) Limestones of the Parvadeh Formation unconformably overlying (white arrow) folded strata of the Hojedk Formation (black arrow) near Mazinu (SW of Tabas). (C) Fluvial, coal-bearing strata of the Hojedk Formation at Kalshaneh capped by a ridge formed by limestones of the Parvadeh Formation (arrow marks the boundary between the two formations). (D) Marine, fossiliferous silts and sandstones of the Hojedk Formation, sandwiched between limestones of the Badamu and Parvadeh formations west of Esfak. (E) Parvadeh Formation at Kalshaneh overlying siliciclastics of the Hojedk Formation. (F) Soft sediments of the Baghamshah Formation forming a north–south-trending valley within the Shotori Mountains; type section of the Baghamshah Formation near Tabas. For co-ordinates of mentioned localities see the Appendix.

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limestones of the Nayband Formation are abruptly replaced by coarse-grained siliciclastics at the Triassic –Jurassic boundary (Huckriede et al. 1962; Fu¨rsich et al. 2005a). Badamu Formation. During the Aalenian (in the southern Tabas Block and on the Lut Block already in the Toarcian; see Seyed-Emami 1967; Seyed-Emami et al. 2000, 2004b) a pronounced transgression initiated the deposition of comparatively condensed, ammonite-rich, and dark, often oolitic limestones, marls and siliciclastic rocks of the Badamu Formation (Toarcian–Lower Bajocian, 0–230 m; Fig. 5A). In the type area near Badamu, 24 km west of Kerman, the formation attains a thickness of only around 50 m. The type section is located near Titu village, 17 km NNE of Zarand, NW of Kerman (southern Tabas Block; Fig. 3) where the formation is thicker (163 m). It is an important marker formation in the expanded and predominantly siliciclastic Lower –lower Middle Jurassic successions of the Tabas Block (without the Badamu Formation it would be impossible to separate the Ab-e-Haji and the non-marine development of the Hojedk Formation in the southern Tabas Block and the western parts of the northern Tabas Block). The Badamu Formation consists of commonly dark-grey, often oolitic and/or bioclastic limestone beds of variable thickness, with intercalated sandy marls, shales and fine-grained calcareous sandstones. In the Mazinu area and at Kuh-e-Rahdar (northern Tabas Block) the formation consists of several 5– 20 m-thick packages of sandy, crossbedded oosparites alternating with marly or argillaceous silts with marine fauna (bivalves, brachiopods, belemnites and small solitary corals). At Kalshaneh (northernmost Tabas Block), coral– microbialite patch reefs are intercalated between a succession of marls, sandstones and bivalve shell beds. The Badamu Formation is fairly fossiliferous, and mainly yields ammonites, bivalves, corals and serpulids. More than 100 ammonite species have been described from the formation (‘Lias-Dogger Cephalopodenkalk’ of Huckriede et al. 1962), allowing a very precise biostratigraphic correlation and a recognition of all NW European standard zones (Huckriede et al. 1962; Seyed-Emami 1967, 1971; Seyed-Emami et al. 2000, 2004b). A contemporaneous transgressive phase is documented from the Lut Block where equivalents of the Badamu Formation are developed in relatively deep, offshore facies (Seyed-Emami et al. 2004b) and in northern Iran (Alborz Mountains), where a significant Toarcian– Aalenian deepening pulse was documented by Fu¨rsich et al. (2005b) from the upper part of the Shemshak Group (Fillzamin Formation; see Fu¨rsich et al. 2009a).

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Hojedk Formation. The Lower Bajocian Hojedk Formation is usually easily identified as the siliciclastic package between the carbonates of the underlying Badamu and the overlying Parvadeh Formation (see later, Fig. 5A–E). According to Sto¨cklin (1971) no type section has been designated for the Hojedk Formation, the type area being the hills north and west of the village of Hojedk in the southern Tabas Block (cf. Huckriede et al. 1962). The argillaceous silts and sandstones of the formation are well exposed in a discontinuous belt, especially in the western part of the Tabas Block, and reach their greatest thickness in the Mazinu and in the Hojedk–Lakar Kuh area (approximately 1000 m; Huckriede et al. 1962; Sto¨cklin 1971; Kluyver et al. 1983a). In the eastern part of the Tabas Block (and on the Lut Block) the formation is generally marine and can be dated using ammonites (e.g. Seyed-Emami et al. 2004b). Towards the west, its character changes to marginal marine and fluvial, including coal swamps (Fig. 5C). The Hojedk Formation is a nearly exclusively siliciclastic unit of dark-grey –brownish-grey and beige –dark-green colour. The grain size varies from argillaceous silt to fine- or medium-grained sandstone at most. The generally immature sediments are poorly to moderately sorted. In the eastern (i.e. marine) occurrences, sandstone units are either bioturbated or exhibit large-scale trough cross-bedding or small-scale ripple lamination. Levels with intraformational pebbles, trace fossils (Rhizocorallium irregulare, Zoophycos, Planolites, Chondrites), and/or shells and shell debris are common, and the different lithologies are often stacked in coarsening- and thickening-upwards cycles bounded by marine flooding surfaces (parasequences). The non-marine outcrops on the western and southern Tabas Block are dominated by trough cross-bedded sandstones alternating with coal seams and carbonaceous silt. The fluvial sandstones are laterally continuous, exhibit a sharp base and are characterized by lateral accretion surfaces, grading upwards into small-ripple lamination. Rootlet horizons are common, and the coal and carbonaceous silt are commonly associated with ferruginous concretions or crusts. Finingupwards (i.e. point bar) cycles dominate the stacking pattern of lithofacies, often topped by a root horizon and commonly overlain by coal. The Hojedk Formation is characterized by rapid lateral facies and thickness changes: in the northern Tabas Block at Parvadeh mine, the thickness of the formation is reduced within a few kilometres from a few hundred to a mere 30 m of marginal marine siliciclastic rocks. Also from the southern Tabas Block, rapid lateral facies and thickness changes are reported (Huckriede et al. 1962).

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This indicates tectonic instability during deposition, culminating in the mid-Bajocian MidCimmerian unconformity capping the Hojedk Formation (Seyed-Emami & Alavi-Naini 1990; Fu¨rsich et al. 2009b) (see Fig. 5B, E). This unconformity separates the Shemshak Group from the following Magu Group (Wilmsen et al. 2003; Seyed-Emami et al. 2004a). Shemshak Group – summary. The Jurassic formations of the Shemshak Group of the Tabas Block largely mirror the trend of facies development and geodynamic events seen in the Alborz Mountains of northern Iran (Fu¨rsich et al. 2005b, 2009a). A significant facies change occurs around the Triassic –Jurassic boundary when marine deposits of the Nayband Formation were replaced by Lower Jurassic non-marine and coal-bearing deposits of the Ab-e-Haji Formation (Fig. 4). As in northern Iran, this event is associated with source-area rejuvenation and indicates uplift and erosion. It is related to the late phase of the Eo-Cimmerian orogeny (Fu¨rsich et al. 2009a). A pronounced marine transgression is indicated by the condensed, ammonite-rich limestones of the Toarcian–Aalenian Badamu Formation. This transgression has its expression also in the Alborz area where it was related to extensional tectonics (Fu¨rsich et al. 2005b). This rifting is inferred to document the onset of Neotethys back-arc rifting in northern Iran (Fu¨rsich et al. 2009a). Rapid lateral facies and thickness variations of the succeeding Hojedk Formation (reaching in part more than 1000 m in thickness) indicate fast and laterally variable subsidence rates suggesting tectonic instability during the Early Bajocian, which culminated in the Bajocian Mid-Cimmerian event (Fu¨rsich et al. 2009b).

Magu and Bidou groups (Bajocian – Upper Jurassic) The Upper Bajocian– Upper Jurassic (Kimmeridgian–?Tithonian) Magu Group (Aghanabati 1977, 1998) is composed of numerous, often laterally equivalent formations (Fig. 4). This strong lithological variability reflects the increased activity of syn-sedimentary faults. The group was subdivided into three subgroups on the basis of (inter-) regional tectonic unconformities (Wilmsen et al. 2003), i.e. the Upper Bajocian– Lower Callovian Baghamshah Subgroup, the Callovian–Kimmeridgian Esfandiar Subgroup and the Kimmeridgian– ? Tithonian Garedu Subgroup. The Bidou Group of the southern Tabas Block is a lithostratigraphic unit contemporaneous with the Magu Group (Fig. 4).

Baghamshah Subgroup The Upper Bajocian –Lower Callovian Baghamshah Subgroup (Fig. 3) consists of mixed siliciclastic – calcareous lithologies, and comprises the Parvadeh, Baghamshah and Sikhor formations as well as the informal Qal’eh Dokhtar Sandstone formation (the ‘Sandstone Member’ of the former Qal’eh Dokhtar Formation of Sto¨cklin et al. 1965; Fig. 4). The subgroup was named after the eponymous Baghamshah Formation, the greenish silts of which are a very characteristic lithological unit of the Tabas area. Parvadeh Formation. The pronounced transgression succeeding the middle Bajocian MidCimmerian unconformity initiated deposition of the Upper Bajocian to Lower–Middle Bathonian condensed, oncolitic–microbial limestones and siliciclastic rocks of the Parvadeh Formation (30– 150 m). It was deposited across the Tabas Block, also onlapping the Shotori Swell (Aghanabati 1977, 1996), and is of Late Bajocian to Early – Middle Bathonian age, as constrained by numerous ammonites that predominantly occur towards the top of the formation (e.g. Seyed-Emami et al. 1991). The Parvadeh Formation is a predominantly calcareous unit with a siliciclastic lower part of variable thickness. The transgressive base of the Parvadeh Formation is sharp, commonly erosional, and often characterized by a decimetre- to metre-thick basal conglomerate of quartzite, milky quartz and sandstone pebbles. Above, it consists mainly of bedded, dark-coloured oncolitic and microbial, sometimes also oolitic limestone, commonly several tens of metres in thickness (Fig. 5B, E). At the type locality near the Parvadeh mine, the formation is fairly marly with intercalation of thin bioclastic, fossiliferous limestone beds. Calcareous sandstone intercalations of decametre thickness may occur in places. The Parvadeh Formation clearly is a transgressive unit deposited during a pronounced deepening interval when the siliciclastic depocentres were shifted far towards the source areas on the Yazd Block (a contemporaneous deepening pulse is recorded from northern Iran, i.e. the Alborz Mountains, succeeding the Mid-Cimmerian unconformity; Fu¨rsich et al. 2005b, 2009b). Near Sikhor, in the southern Shotori Mountains, the Parvadeh Formation onlaps Upper Palaeozoic formations of different ages with thick, cliff-related basal conglomerates (Fig. 6). In the eastern part of the Tabas Block, close to the boundary with the Lut Block, the Parvadeh Formation pinches out or grades into the as yet unformalized Qal’eh Dokhtar Sandstone formation (see below). Further east, on the Lut Block, the Parvadeh Formation is present again. In the southern part of the Tabas Block the

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Fig. 6. Onlap of the Parvadeh Formation with basal conglomerate onto Palaeozoic strata in the southern Shotori Mountains near Sikhor (for co-ordinates see the Appendix) reflecting the Late Bajocian–Bathonian submergence of the Shotori Swell.

thickness of the formation varies greatly. West of Ravar, at N318090 3100 , E568460 2600 , it is well developed and exhibits the characteristic features of the unit such as oncoidal–microbial limestones at the top (pers. obs.). Furthermore, a thin quartz microconglomerate occurs at its base, which most probably corresponds to the conglomerates associated with the Mid-Cimmerian unconformity elsewhere (Fu¨rsich et al. 2009b). Qal’eh Dokhtar Sandstone formation. The informal Qal’eh Dokhtar Sandstone formation was originally defined as the ‘sandstone member of the Qal’eh Dokhtar Formation’ by Sto¨cklin et al. (1965). The Qal’eh Dokhtar Sandstone formation is a 186 m-thick unit of cross-bedded, fine- to medium-grained marine sandstone (see Schairer et al. 2000, fig. 4.1, p. 41). It only occurs at the type locality of the former Qal’eh Dokhtar Formation in the easternmost part of the study area at the transition to the Lut Block (Sto¨cklin et al. 1965; Schairer et al. 2000) (Figs 4 and 7B). The Qal’eh Dokhtar Sandstone formation rests with sharp basal contact on siltstones of the Hojedk Formation and is overlain by a rather silty –fine-sandy Baghamshah Formation (see below). These stratigraphic relationships suggest that the Qal’eh Dokhtar Sandstone formation is a lateral equivalent

of the Parvadeh Formation or at least partly corresponds to the stratigraphic gap at the base of this formation, which is present at many places (for a discussion and a redefinition of the Qal’eh Dokhtar Formation see Wilmsen et al. 2003). However, in the absence of precise stratigraphic data and owing to the limited regional extent of the Qal’eh Dokhtar Sandstone formation its name is kept at an informal level (Wilmsen et al. 2003) until further information becomes available. Baghamshah Formation. The Late Bathonian was characterized by high rates of subsidence and the onset of widespread deposition of fine-grained shelf sediments of the Baghamshah Formation, the type locality of which is the valley of Baghamshah in the central Shotori Mountains ENE of Tabas (Sto¨cklin et al. 1965) (Fig. 5F). The thickness of the Baghamshah Formation is usually between 400 and 600 m, reaching more than 1000 m in the south of the study area (Kluyver et al. 1983a, b). Its age is (?Middle) Late Bathonian– Early Callovian (Seyed-Emami et al. 1997, 1998) and may reach locally up into the Middle Callovian (Seyed-Emami et al. 2002). The Baghamshah Formation consists mainly of soft, greenish, variably marly silts with some admixture of fine-sand or thin-sandstone intercalations

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Fig. 7. (A) Deltaic and fluvial siliciclastics of the Sikhor Formation capped by the lower part of the cliff-forming Esfandiar Limestone Formation at Kuh-e-Neygu near Khoda–Afarid. Arrow marks the boundary between the two formations. (B) Dark limestones of the Qal’eh Dokhtar Limestone Formation at the type section near Qal’eh Dokhtar. (C) Eastern slopes of the Shotori Mountains near Korond, formed by the steeply eastward-dipping Esfandiar Limestone Formation. (D) Greenish, silty marls of the Korond Formation sharply overlying (arrow) the Esfandiar Limestone Formation, approximately 1 km south of Korond. (E) Soft, rhythmically bedded marls and marly limestones of the Kamar-e-Mehdi Formation capped by the Nar Limestone Member of the Kamar-e-Mehdi Formation in the type area near Kuh-e-Qoleh Nar. (F) Detail of the tripartite Nar Limestone Member of the Kamar-e-Mehdi Formation in the type area near Kuh-e-Qoleh Nar (arrow marks the lower boundary of the member). (G) Reddish limestone conglomerates and siliciclastics of the Garedu Red Bed Formation erosionally overlying the Esfandiar Limestone Formation north of Honu in the northern Shotori Mountains (arrow marks the formational boundary). (H) Nar Limestone Member of the Kamar-e-Mehdi Formation at Kuh-e-Qoleh Nar unconformably overlain by red clays and gypsum of the Magu Gypsum Formation (arrow marks the boundary between the two formations). For co-ordinates of mentioned localities see the Appendix.

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(Fig. 5F). Likewise thin (at metre-scale) intercalations of limestone (oosparite, microbialites) occur locally. In the western parts of the Tabas Block, the formation is fairy marly and fine-grained, whereas towards the east, the silt and sand content increases. At the boundary with the Lut Block, it is represented by intercalations of silts and finegrained, often sharp-based, sandstone beds (the former ‘siltstone member of the Qal’eh Dokhtar Formation’; see redefinition of these units by Wilmsen et al. 2003). The Baghamshah Formation is generally poorly fossiliferous. However, some levels (mainly limestone intercalations) are fairly rich in fauna, yielding ammonites, bivalves, gastropods, sponges and crinoids. The formation was deposited in open-marine, middle –outer shelf to upper-slope settings. The latter is inferred from thick (.1000 m) turbidite units within the Baghamshah Formation in the Lakar Kuh area, which are intensely folded in large-scale slump units owing to synsedimentary deformation (Kluyver et al. 1983a). Sikhor Formation. At the Bathonian–Callovian boundary, asymmetric uplift of an east-vergent fault block at the eastern margin of the Tabas Block resulted in considerable erosion and concomitant NNE-ward progradation of the fluvio-deltaic clastic wedge of the Lower Callovian Sikhor Formation (0– 400 m; Fu¨rsich et al. 2003a) (Fig. 7A). The resulting north–south-trending topographic high (Shotori Swell) within the Baghamshah shelf system served as the nucleus for the development of the carbonate system of the succeeding Esfandiar Subgroup (see below). The Sikhor Formation is Early Callovian in age, suggesting that the tectonic instability at the eastern margin of the Tabas Block started around the Bathonian– Callovian boundary. Facies changes of the Sikhor Formation are pronounced in the east–west and less so in north– south direction, and the formation pinches out towards the east, west and north (there is no information from south of the Esfandiar Formation type section). The geodynamic significance of the unit (‘sandy base of the Esfandiar Limestone’ of Sto¨cklin et al. 1965) was recognized by Fu¨rsich et al. (2003a). The Sikhor Formation predominantly consists of siliciclastic (conglomerate, sandstone, siltstone) and mixed and/or interbedded siliciclastic– carbonate rocks. It was subdivided into an exclusively siliciclastic, fluvio-deltaic Kuh-e-Neygu Member, and a mixed siliciclastic –calcareous Majd Member of a marine ramp origin (Fu¨rsich et al. 2003a). At the type locality at Sikhor, southern Shotori Mountains, fluvial conglomerates of the Kuh-e-Neygu Member deeply cut into the Baghamshah Formation, followed by fluvial

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fining-upwards cycles and flood-plain fines with pedogenic overprint (caliche nodules, calcrete layers). Elsewhere, deltaic sediments rapidly prograded into the Baghamshah shelf, partly followed by fluvial deposits. With decreasing siliciclastic input, a juvenile carbonate system developed, resulting in the deposition of the mixed siliciclastic –carbonate strata of the Majd Member of the Sikhor Formation. The termination of siliciclastic input marks the change to the Esfandiar Subgroup. Bidou Formation. On the southern Tabas Block, the Bidou Formation (‘Bidou-Schichten’ of Huber & Sto¨cklin 1954; ‘Bidou-Fazies’ or ‘Bidou-Serie’ of Huckriede et al. 1962) is equivalent, in parts, to the Baghamshah Subgroup of the northern Tabas Block. The type area is the Bidou syncline approximately 70 km NNW of Kerman around the small village of Bidou (Hojedk coal district). However, the upper parts of the formation are not exposed there. The age of the Bidou Formation is poorly constrained by fossils but could be (Late Bajocian? –) Bathonian–Callovian (Seyed-Emami et al. 2001). The Bidou Formation, the thickness of which varies between 500 and more than 1000 m, consists of varicoloured (beige, brownish-red–violet, green) conglomerates, calcareous sandstones, pebbly sandstones, shales, gypsiferous marls and limestones with a marine fauna (mainly bivalves and gastropods). The lithofacies, as well as sedimentary structures such as large-scale trough cross-bedding, ripple marks and mud desiccation cracks, suggest a coastal plain and marginal–shallow-marine depositional environment for the Bidou Formation under (semi-)arid conditions. Coarse, polymict conglomerates with metamorphic and volcanic components, as well as a diverse suite of Palaeozoic–Triassic (meta-)sediment pebbles (Huckriede et al. 1962, pp. 98–99), occur preferentially at the base and near the top of the formation. At Kuh-e-Bande-Buhabad (200 km NW of Kerman), a few metres of fossiliferous, oolitic–oncolitic, darkcoloured limestones near the base of the formation (Huckriede et al. 1962: p. 91) may be regarded as a lateral equivalent of the Parvadeh Formation of the northern Tabas Block (see Fig. 4). The conglomerates at the base correspond to an erosional episode following the Mid-Cimmerian tectonic event (Fu¨rsich et al. 2009b). The conglomerates in the upper part of the formation may be an expression of the tectonic instability recorded from the Bathonian–Callovian boundary interval of the northern Tabas Block by Fu¨rsich et al. (2003a) and Seyed-Emami et al. (2004a). Baghamshah Subgroup – summary. The Baghamshah Subgroup documents the onlap of depositional

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units onto formerly partly emergent areas after the compressional phase of the middle Bajocian MidCimmerian tectonic event (Fu¨rsich et al. 2009b). Folding and erosion can be demonstrated for this interval from the Tabas Block (Mazinu area: Wilmsen et al. 2003; Seyed-Emami et al. 2004a) (Fig. 5B). The succeeding Late Bajocian– Bathonian onlap of the Parvadeh and Baghamshah formations was caused by increased subsidence of the Tabas Block, because the global eustatic sea level in fact dropped during this interval (Hardenbol et al. 1998; Hallam 2001), similar to the situation in the Alborz Mountains of northern Iran (Fu¨rsich et al. 2005b). This onlap also affected formerly permanent emergent areas in the Shotori Mountains at the eastern margin of the Tabas Block (Shotori Swell). The sediments of the Parvadeh Formation record a rapid landward shift of the shoreline and considerable condensation during the initial stages of this transgression (a basal conglomerate is often present and followed by microbial, ammonitebearing limestones). The succeeding widespread and uniform Baghamshah Formation is indicative of deeper marine deposition (mid-shelf–upper slope). The thick, slumped turbidite units within the Baghamshah Formation of the Lakar Kuh area (Kluyver et al. 1983a) suggest the presence of a considerable gradient within the basin and may also indicate some seismic activity that triggered slumping. The tectonic pulse that caused deposition of the Sikhor Formation was related to rotation of the Tabas Block around its long, i.e. north –south, axis and caused a renewed uplift of the Shotori Swell, which represents the crest of the rotated fault block (Fu¨rsich et al. 2003a). After erosion of the crest top (at some localities nearly the complete Baghamshah Formation was eroded) and concomitant deposition of the Sikhor Formation, the Shotori Swell served as the nucleus for the development of a shallow-marine carbonate platform.

Esfandiar Subgroup The Callovian– Upper Jurassic rocks show a predominance of calcareous sediments that were grouped into the Esfandiar Subgroup (Fig. 4). It consists of the Esfandiar Limestone, Qal’eh Dokhtar Limestone, and Korond and Kamar-e-Mehdi formations, and is named after the characteristic, cliff-forming Esfandiar Limestone Formation. The Esfandiar and Qal’eh Dokhtar limestone formations were described in detail by Fu¨rsich et al. (2003b), the Korond Formation by Schairer et al. (2003). The Esfandiar Subgroup represents a large-scale shelf lagoon– carbonate platform –slope-to-basin system. In the Ravar –Kerman area, the Esfandiar Subgroup is respresented by the Kamar-e-Mehdi Formation (Fig. 4).

Esfandiar Limestone Formation. The platform deposits of the Esfandiar Subgroup are represented by the Esfandiar Limestone Formation, occurring in a north–south-trending outcrop belt close to the eastern margin of the Tabas Block (Shotori Swell; Fig. 7A, C). The age of the Esfandiar Limestone Formation is Early Callovian –(?Early) Kimmeridgian (Sto¨cklin et al. 1965; Fu¨rsich et al. 2003b; Bagi & Tasli 2007). The type section is at Kuh-e-Esfandiar in the southern part of the Shotori Mountains (Sto¨cklin et al. 1965), where it reaches a thickness of 760 m. The Esfandiar Limestone Formation is underlain by the Baghamshah or the Sikhor Formation, and the contact is often gradational (Fig. 4). The upper contact is usually sharp, and the formation may be overlain by formations of very different ages (see below). To the east of the Shotori Mountains, the Esfandiar Limestone Formation interfingers with the Qal’eh Dokhtar Limestone Formation (Fig. 7B) (Fu¨rsich et al. 2003b) and to the west with the Kamar-e-Mehdi Formation (Fig. 7E; e.g. north of Tabas at Kuh-eBagh-e-Vang). The Esfandiar Limestone Formation entirely consists of medium-bedded to massive carbonates of light-grey –brownish-yellow colour, which were deposited on an extensive carbonate platform, the ‘Esfandiar Platform’ of Fu¨rsich et al. (2003b). The lower parts of the formation are characterized by slightly darker colours, whereas the middle and upper parts are often very-light grey to nearly white. At the eastern limit of the outcrop belt the upper part often shows yellow-brownish colours. Bioturbation is relatively rare; occasionally Thalassinoides burrows may occur. Primary sedimentary structures such as large-scale trough cross-bedding are likewise rare, and confined to bioclastic and/ or oolitic grainstones and rudstones, which mainly occur towards the eastern margin of the platform. These deposits are interpreted as shoal deposits of the platform margin. Most of the Esfandiar limestones, however, are fine-grained mudstones or wackestones with an impoverished fauna of the interior platform. Characteristic, albeit not common, faunal elements of the Esfandiar Limestone Formation are diceratid bivalves and sponges (Fu¨rsich et al. 2003b). The minute sclerosponge Neuropora sp. is rather abundant in fine-grained platform lagoonal sediments, often associated with Tubiphytes, whereas representatives of the genus Cladocoropsis are rare. Chaetetid sponges (?Ptychochaetetes sp.) occur mainly as hemispheroids up to a diameter of 50 cm, being more common towards the platform margin. Neither at the platform margin nor in the interior were true reefal structures observed (contrary to the characterization of the Esfandiar Limestone Formation as ‘reef limestone’ by Sto¨cklin et al. 1965). A detailed microfacies analysis was

JURASSIC OF THE TABAS BLOCK, IRAN

presented by Fu¨rsich et al. (2003b), revealing that the Esfandiar Limestone Formation represents a shallow- and warm-water depositional system. Qal’eh Dokhtar Formation. Towards the east, at the junction of the Tabas and Lut blocks, the platform sediments of the Esfandiar Limestone Formation interfinger with deeper-water sediments of the up to 400 m-thick Qal’eh Dokhtar Limestone Formation (the former Qal’eh Dokhtar Limestone Member of the Qal’eh Dokhtar Formation of Sto¨cklin et al. 1965). At the type locality at Qal’eh Dokhtar (Fig. 7B), the formation is 372 m thick. The age of the Qal’eh Dokhtar Formation is well constrained by ammonites, and comprises the Middle Callovian –Late Oxfordian (Schairer et al. 2000, 2003). The Qal’eh Dokhtar Limestone Formation consists of bedded, dark-grey and often sharp-based, oolitic –bioclastic limestone, between which finegrained marl and marly limestone are intercalated. Lenticular, metre- to decametre-scale carbonate bodies composed of siliceous sponges and microbial crusts, channelized limestone conglomerates and boulder beds with platform-derived biota may occur. The marly intercalations reach up to 10– 20 m in thickness, and contain abundant ammonites and trace fossils (Zoophycos). Fu¨rsich et al. (2003b) presented a detailed microfacies analysis. The Qal’eh Dokhtar Limestone Formation represents platform-derived allochthonites (graded allodapic limestones, debris- and mud-flow deposits, olistostromes) and autochthonous sediments such as periplatform muds settling from suspension and patchy microbial sponge reefs. The characteristic association of allodapic limestones with peri-platform muds and bioclastic debris-flow deposits and olistostromes was interpreted by Fu¨rsich et al. (2003b) as response to sea-level fluctuations. Korond Formation. The Korond Formation was erected by Schairer et al. (2003). It comprises a thick succession of light-green silty marl forming a north–south-trending outcrop belt bordering the eastern margin of the northern Tabas Block (Fig. 7D). Owing to its lithological similarity to the Baghamshah Formation, this unit has been mapped as the latter on geological maps of the area (e.g. Sto¨cklin et al. 1965; Ruttner et al. 1968). However, it overlies either the Esfandiar Limestone Formation or the Qal’eh Dokhtar Limestone Formation (Fig. 7D), and clearly represents a characteristic lithostratigraphic unit on its own, the recognition of which dramatically simplified the geological structures of the area. The thickness of this recessive formation is not known owing to the absence of complete sections, but it certainly exceeds 1000 m in the type area around the small

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village of Qassem-Abbad. The age of the Korond Formation, however, is well known due to the occurrence of numerous ammonites, which suggest a Middle Oxfordian– Kimmeridgian age. Its base is younging towards the west, displaying onlap onto the Qal’eh Dokhtar Limestone and the eastern occurrences of the Esfandiar Limestone Formation. The Korond Formation is characterized by greenish marl, silty marl and subordinate siltstone –very-fine-grained sandstone. The lower part is dominated by structureless and only weakly silty and poorly fossiliferous marl. Occasionally intercalated biorudstones show a sharp, erosional base and are rich in allochthonous platform-derived material (crinoid ossicles, echinoid spines, sponges, bivalves, brachiopods, coated grains, oncoids and intraclasts). Higher up, the siliciclastic content gradually increases, and silty marl, marly silt and ripple- or hummocky cross-stratified fine-grained sandstones become a characteristic feature. These lithologies are often bundled into 10 –30 m-thick coarsening- and thickening-upwards cycles (parasequences), and yield a moderately diverse assemblage of bivalves (e.g. Pinna, Arcomytilus, Pholadomya, Homomya, Integricardium, Rostroperna, Vaugonia), irregular echinoids and ammonites (Sowerbyceras, Orthosphinctes, Idoceras and Physodoceras). The Korond Formation is fairly monotonous and was deposited in a deeper, open-marine basinal environment. The occasional intercalations of limestones with shallow water components in the lower part indicate an upwards-diminishing influence of the Esfandiar Platform. Its onlap onto the eastern parts of the Esfandiar Platform indicates a drowning event. Towards the top of the formation the basin distinctly shallowed, as is shown by the presence of ripple lamination and hummocky crossstratification, suggesting that the basin floor was just within the reach of storm waves. Kamar-e-Mehdi Formation. The thick Kamare-Mehdi Formation represents the Esfandiar Subgroup in the central and western part of the Tabas Block. At the type locality of the formation between Kamar-e-Mehdi and Kuh-e-Qoleh Nar in the northern part of the Tabas Block (Fig. 3; see Aghanabati 1977 and Fig. 6E), it reaches a thickness of about 1300 m. The formation is very widespread, occurring also in the area around Ravar, in the southernmost part of the block (Fig. 3). In the northern part of the Tabas Block, two members can be recognized within the formation: a basal Echellon Limestone Member and a capping, cliffforming Nar Limestone Member (Wilmsen et al. 2003). The Kamar-e-Mehdi Formation was deposited in an extensive shelf lagoon on the western (i.e. landward) side of the Esfandiar Platform.

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Its age is Middle Callovian–Early Kimmeridgian, based on rare ammonite finds and benthic foraminifera. Except for the basal Echellon Limestone Member and the terminal Nar Limestone Member, the Kamar-e-Mehdi Formation is recessive, and predominantly consists of greyish marl, marly limestone and limestone with micritic fabric, typically stacked into metre-scale, asymmetric thickening-upwards cycles (Fig. 7E). Intercalated are occasional thin (1– 25 cm), sharp-based, finegrained sandstone and siltstone beds. From the type section, two gypsum– mudstone intercalations are recorded, each decametre-thick and thinbedded. The biofacies of the Kamar-e-Mehdi Formation is dominated by bivalves with large pectinids being the most conspicuous element (‘Pectinidenkalk-Gips-Fazies’ of Huckriede et al. 1962; ‘Calcaire a` Pectens’ of Aghanabati 1977). In the northern part of the Tabas Block, the lowermost part of the Kamar-e-Mehdi Formation is represented by the Kuh-e-Echellon Member (Wilmsen et al. 2003; ‘calcaire d’Echellon’ of Aghanabati 1977). The member is up to 200 m thick in the type area of Kuh-e-Echellon, and characterized by marly oncolites, bioclastic floatstones, oolites, microbialites, marls, shell beds and patch reefs dominated by corals and/or microbial crusts and sponges. The Nar Limestone Member (‘calcaire de Nar’ of Aghanabati 1977) is a cliffforming unit at the top of the Kamar-e-Mehdi Formation (Fig. 7F). It is 90 –150 m thick, and consists of monotonous, well bedded mudstones and peloid wackestones –packstones that commonly contain pseudomorphs after gypsum crystals. In the Echellon area, some tens of kilometres to the north of Kamar-e-Mehdi, an Early Kimmeridgian age for the Nar Limestone Member is indicated by agglutinated benthic foraminifera. The bio- and microfacies (a low-diversity fauna, abundance of peloids, gypsum pseudomorphs) of the Nar Limestone Member indicates a restricted (i.e. hypersaline) environment. Towards the southern Tabas Block, this hypersalinity increased as is shown by increasing gypsum precipitation (‘Pecten-KalkGips-Fazies’ of Huckriede et al. 1962). Esfandiar Subgroup – summary. The Esfandiar Subgroup documents the existence of a large-scale low-latitude carbonate system on the Tabas Block during the late Middle–Late Jurassic. The Esfandiar Limestone Formation represents a fault-block carbonate platform (sensu Leeder & Gawthorpe 1987) developing on the crest (‘Shotori Swell’) of a (today west-vergent) tilt block (Fu¨rsich et al. 2003b). As a north –south-trending barrier platform it sheltered a large-scale shelf-lagoon (Kamar-eMehdi Formation) on the hanging wall dip-slope,

in which fine-grained calcareous sediments accumulated. To the west, this shelf lagoon was bordered by the emergent Yazd Block. The eastern margin of the Esfandiar Platform was controlled by down-faulting at the footwall scarp (Nayband Fault separating the Tabas from the Lut Block), and the eastern slope and adjacent basinal area (Qal’eh Dokhtar Limestone and Korond formations) received large amounts of platform-derived material (gravitational sediments and peri-platform muds) during the persistence of the carbonate system of the Esfandiar Subgroup. Isolated occurrences of the Qal’eh Dokhtar and Esfandiar limestone formations are also known from the Lut Block (Kuh-e-Birg). The demise of this carbonate system started with the onlap of the basinal Korond Formation onto the platform around the Oxfordian – Kimmeridgian boundary and coincided with an apparent cut-off of the shelf lagoon from the open sea (Nar Limestone Member of the Kamar-e-Mehdi Formation). The final destruction occurred during the Late Cimmerian tectonic event (Kimmeridgian–earliest Cretaceous; see later).

Garedu Subgroup The Upper Jurassic (Kimmeridgian–Tithonian) Garedu Subgroup (Fig. 4) is named after the Garedu Red Bed Formation (Ruttner et al. 1968). The lithology of the subgroup is dominated by reddish limestone conglomerates, red siltstones and sandstones, and gypsum with siliciclastic intercalations. Its base is drawn at an intra-Kimmeridgian tectonic unconformity terminating the deposition of the Esfandiar Subgroup. Apart from the eponymous formation, the Garedu Subgroup only contains the Magu Gypsum Formation (Fig. 4). On the southern Tabas Block, the Ravar Formation forms a contemporaneous and facies-equivalent lithostratigraphic unit, which additionally is characterized by intercalations of volcanics (Huckriede et al. 1962). Garedu Red Bed Formation. This formation occurs in the Shotori Mountains at the eastern margin of the Tabas Block, where it unconformably overlies the Esfandiar Limestone Formation. The type locality of the Garedu Red Bed Formation is near the old mine of Garedu on the western flank of the northern Shotori Mountains (Ruttner et al. 1968). The maximum thickness of the formation in that area is not exactly known because no complete sections have been located. However, it certainly exceeds several hundred metres but varies considerably over short distances. The age of the Garedu Red Bed Formation is poorly constrained by biostratigraphic data. Owing to the superposition on the Callovian–Lower Kimmeridgian Esfandiar Limestone Formation (Fu¨rsich et al. 2003b; Schairer

JURASSIC OF THE TABAS BLOCK, IRAN

et al. 2003; Bagi & Tasli 2007), a post-Early Kimmeridgian age is indicated. Ruttner et al. (1968) reported rare calcareous algae of inferred Kimmeridgian–Tithonian age from the Garedu Red Bed Formation of the northern Shotori Mountains. The lower part of the Garedu Red Bed Formation of the northern Tabas Block consists of coarsely bedded and poorly sorted limestone conglomerates with a sandy, calcareous matrix, which unconformably or gradually overlie the Esfandiar Limestone Formation (Fig. 7G). The limestone pebbles and cobbles are exclusively derived from the Esfandiar Limestone Formation. There are occasional intercalations of limestones with oyster fragments, carbonate mudstones and microbialites, as well as reddish and brownish sandstones. In this lower part of the Garedu Formation there is a considerable lateral variation in thickness and facies (e.g. at the type section near the Garedu mine; Ruttner et al. 1968, p. 93, fig. 31) indicating faultcontrolled deposition in a closely spaced array of uplifted and downthrown blocks. Up-section, conglomerate intercalations become rarer and thinner, and pebbly sandstones and red siltstones of fluvial channel and flood-plain origin predominate. The onset of Garedu Red Bed deposition clearly indicates a prominent tectonic event being strongest at the eastern margin of the Tabas Block. Magu Gypsum Formation. The Magu Gypsum Formation (‘Gypse de Magou’ of Aghanabati 1977) crops out in the western part of the Tabas Block where it sharply overlies the Nar Limestone Member of the Kamar-e-Mehdi Formation (Fig. 7F). The thickness of the formation is about 600 m at the type section (Aghanabati 1977) at Kuh-e-Qoleh Nar (Fig. 7H). There are no biostratigraphic data from the formation itself, but considering that in the Kuh-e-Echellon area (northern Tabas Block) the upper part of the underlying Nar Limestone Member yielded Early Kimmeridgian benthic foraminifera, a post-Early Kimmeridgian age for the overlying Magu Gypsum Formation is confirmed. The Magu Gypsum Formation is probably Late Kimmeridgian–Tithonian in age and may even range into the Early Cretaceous (Aghanabati 1977). The Magu Gypsum Formation consists of softweathering, red, fine-grained sandstone, siltstone and clay, with intercalated gypsum beds and limestone and dolomite beds. These carbonates contain mono- to paucispecific bivalve concentrations and are cross-bedded oolitic grainstones or intragrainstones and rudstones. The gypsum beds are either massive units up to several tens of metres in thickness or are intercalated with reddish –greenish marl. The Magu Gypsum Formation was deposited in a

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sabkha-type environment that was subject to episodic marine ingressions. In the Echellon area, the contact between the Kamar-e Mehdi and Magu Gypsum formations is more gradual, and 12 m above the base a lenticular 6 –17 m-thick clast-supported conglomerate is present, which can be followed for several kilometres. It consists of rounded limestone pebbles up to 30 cm in diameter derived from the underlying Nar Limestone Member in a sandy matrix. This Nar Conglomerate Member indicates tectonic activity at the transition from the Kamar-e-Mehdi to the Magu Gypsum formations (Late Cimmerian event), corresponding to the onset of Garedu Red Bed deposition in the Shotori Mountains. The Ravar Formation of the Kerman-Ravar area is an equivalent formation (see Huckriede et al. 1962 for details). Ravar Formation. On the southern Tabas Block, the Ravar Formation (sensu Huckriede et al. 1962 non Sto¨cklin 1961) is the equivalent of the Garedu Subgroup (Fig. 4). Its type area is around Ravar, approximately 210 km NNW of Kerman (Fig. 3). The age of the Ravar Formation is Late Jurassic– early Early Cretaceous (‘Neocomian’) based on stratigraphic relationships with the underlying Kamar-e-Mehdi Formation and overlying Barremian –Aptian strata (Huckriede et al. 1962). In the literature, there is some confusion about the name of the Ravar Formation. Strictly speaking, the name was already in use when Huckriede et al. (1962) introduced it because the name had been given to an ‘Infra-Cambrian’ evaporitic formation by Sto¨cklin (1961) the year before. However, this much older formation is only poorly exposed in the Ravar area (in contrast to the younger one), and Huckriede et al. (1962) named the same ‘InfraCambrian’ evaporitic formation as the Desu Formation after the village Desu east of Zarand where those rocks are well exposed below the fossiliferous Cambrian strata. The name ‘Upper Jurassic Salt’ suggested by Sto¨cklin (1961, 1971) for the younger deposits is formally not acceptable. Thus, the consensus is that the ‘Infra-Cambrian’ evaporitic formation is called the Desu Formation, and the Upper Jurassic–Neocomian one, the Ravar Formation. The Ravar Formation consists of a varicoloured intercalation of evaporites (dark-coloured dolomites, gypsum, some rock salt), gypsiferous marls, marls, violet sandstones and clays, as well as some dark porphyritic–dioritic volcanics including quartz porphyries. Owing to the soft nature of the strata, the formation is internally commonly strongly deformed, so that thicknesses are difficult to assess. Seyed-Emami et al. (2001) estimated approximately 400 m for the Ravar –Kerman area. Similar to the

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Magu Gypsum Formation of the northern Tabas Block, a sabkha-type depositional environment is assumed. The intercalated volcanics may indicate synsedimentary tectonic activity. Garedu Subgroup – summary. Increased tectonic movements along the Nayband Fault during the Late Oxfordian –Early Kimmeridgian led to partial drowning of eastern areas of the Esfandiar Platform (deposition of the basinal Korond Formation) and a cut-off of the shelf lagoon from the open sea in the Early Kimmeridgian (gypsum unit and Nar Limestone Member of the Kamar-e-Mehdi Formation). This gradual decline of the Esfandiar Subgroup carbonate system was followed by pronounced syn-sedimentary tectonic activity starting in the Kimmeridgian (Late Cimmerian event), which resulted in its final destruction. Strata of the Esfandiar Subgroup were partly eroded, and subsequently covered by Kimmeridgian –Tithonian limestone conglomerates and red siliciclastic sediments and evaporites (Garedu Red Bed and Magu Gypsum Formation; Ravar Formation). However, the age of the Garedu Subgroup is poorly constrained and it may well range into the Lower Cretaceous. The geodynamic background for this Late Cimmerian tectonic event, which can be traced all across the Iran Plate, is still poorly understood. A possible but largely speculative explanation is the oblique subduction of the Neotethys mid-ocean ridge system below the Iran Plate, which, according to Stampfli & Borel (2002), could have taken place around the Jurassic – Cretaceous boundary. Large parts of the Tabas Block remained emergent until the Cenomanian–Turonian, when a shallow sea transgressed onto the Shotori Mountains (Wilmsen et al. 2005).

Discussion and conclusions The Central-East Iranian Microcontinent (CEIM), a part of the Iran Plate, consists of three, today north– south-oriented, structural units, called the Lut, Tabas and Yazd blocks. The Iran Plate as an element of the Cimmerian microplate assemblage became detached from Gondwana during the Permian, and collided with Eurasia (Turan Plate) during the Late Triassic thereby closing the Palaeotethys (Eo-Cimmerian event; e.g. Sto¨cklin 1974; Boulin 1988; Stampfli & Borel 2002; Fu¨rsich et al. 2009a). According to some geodynamic models (e.g. Soffel & Fo¨rster 1984; Soffel et al. 1996; Alavi et al. 1997; Besse et al. 1998), the CEIM experienced anticlockwise rotation by 1358 in post-Triassic times into its present-day position associated with considerable lateral movements along the block-bounding faults.

Facies analyses and stratigraphic relationships suggest that, during the Jurassic Period, the three blocks had the same lateral arrangement with respect to each other as today (Fig. 2), arguing against significant strike-slip movements along the block-bounding faults. The Yazd Block was emergent for most of the period, and there is a general trend for increasing marine influence from the Tabas to the Lut Block. The Tabas Block seems to have been rotated as a fault block around its long axis, its eastern margin acting as fault-block crest (the Shotori Swell of former authors). However, the blocks were most probably east –west oriented, and the Lut block faced the Neotethys, meaning any anticlockwise rotation of the CEIM should have been of post-Jurassic age (Fig. 2). This interpretation is supported by volcanic rocks (pillow lavas, palagonitic tuffs and keratophyre dykes) from a marine, shaly unit of Jurassic age from Kuh-e-Abdullahi, NW of the Deh Salm and the Jurassic Shah-Kuh granite in the present-day eastern Lut Block (Sto¨cklin et al. 1972; Esmaeily et al. 2007). These magmatic rocks may be related to the Neotethys subduction (volcanic arc), which we assume to have been present at the then southern margin of the Jurassic Lut Block (Fig. 2). This suggests a back-arc position for the CEIM during Jurassic times. The Upper Triassic –Jurassic succession of the Tabas Block is subdivided into major tectonostratigraphic units based on widespread unconformities related to the Cimmerian tectonic events (see Stille 1924; Ziegler 1975; Fu¨rsich et al. 2009b). The Eo-Cimmerian unconformity is a welldeveloped palaeokarst surface, separating Lower – Middle Triassic platform carbonates (Shotori Formation) from siliciclastic rocks of the Norian– Bajocian Shemshak Group (Fu¨rsich et al. 2005a). Norian–Rhaetian sediments are predominantly marine and consist of fine siliciclastic and calcareous sediments accommodated in the up to 3000 m-thick Nayband Formation that was deposited in extensional basins separated by horst structures (Fu¨rsich et al. 2005a). At the Triassic – Jurassic boundary, the main uplift phase of the Eo-Cimmerian orogeny in northern Iran (Fu¨rsich et al. 2009a) resulted in source-area rejuvenation (deposition of the non-marine Lower Jurassic– Aalenian Ab-e-Haji Formation; Fig. 8). During the Aalenian (in the southern Tabas Block and on the Lut Block already in the Toarcian) a pronounced transgression initiated the deposition of condensed, ammonite-rich limestones of the Badamu Formation (Toarcian– lowermost Bajocian), overlain by the thick marine– continental siliciclastics of the Lower Bajocian Hojedk Formation. Rapid lateral facies and thickness changes indicate tectonic instability, culminating in the mid-Bajocian

JURASSIC OF THE TABAS BLOCK, IRAN

Fig. 8. Geochronological summary diagram showing the main geological events and the second-order relative sea-level trend of the Jurassic System on the Tabas Block, east-central Iran. Absolute ages after Gradstein et al. (2004).

Mid-Cimmerian event, evident by stratigraphic gaps, quartz conglomerates and mild folding in some places (Seyed-Emami & Alavi-Naini 1990; Wilmsen et al. 2003; Seyed-Emami et al. 2004a; Fu¨rsich et al. 2009b). This widespread unconformity separates the Shemshak Group from the following Magu – Bidou groups (Fig. 8). The Magu Group (Upper Bajocian –Upper Jurassic) is subdivided into three subgroups. The Baghamshah Subgroup documents the onlap of marine depositional units (Parvadeh and Baghamshah formations) onto formerly emergent areas after the compressional phase of the Mid-Cimmerian tectonic event (Fu¨rsich et al. 2009b), caused by increased subsidence of the Tabas Block. This development is comparable to the pronounced Late Bajocian –Bathonian deepening phase in the Alborz Mountains of northern Iran, which is related to post-rift subsidence of the South Caspian Basin (Brunet et al. 2003; Fu¨rsich et al. 2005b,

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2009b). Increased rotation of the Tabas Block around its long axis caused the uplift and erosion of its (present-day) eastern margin (the Shotori Swell) and the deposition of the uppermost Bathonian(?) – Lower Callovian Sikhor Formation (Fu¨rsich et al. 2003a). This Shotori Swell served as the nucleus for the development of a large-scale carbonate system, consisting of, from west to east, shelf-lagoonal (Kamar-e-Mehdi Formation), platform (Esfandiar Limestone Formation) and slope deposits (Qal’eh Dokhtar Limestone Formation), combined in the Callovian– Kimmeridgian Esfandiar Subgroup (Fu¨rsich et al. 2003b; Wilmsen et al. 2003) (see Fig. 8). This carbonate system flourished until the Oxfordian– Kimmeridgian boundary, when increasing tectonic activity occurred, especially at the eastern margin of the Tabas Block (Nayband Fault). The tectonic instability culminated in intense block faulting during the Kimmeridgian related to the Late Cimmerian event (Fig. 8). This tectonic event resulted in the final destruction of the Esfandiar Subgroup carbonate system, which was subsequently eroded and partly covered by Kimmeridgian –Tithonian limestone conglomerates, red beds and evaporites of the Garedu Subgroup. A possible explanation for this widespread, but hitherto poorly understood, tectonic event may be the subduction of the Neotethys mid-ocean ridge system below the Iran Plate. The long-term relative sea-level trend for the Tabas Block (Fig. 8) is clearly dominated by tectonic signals (regional uplift and subsidence phases), as the correlation to Jurassic eustatic sealevel charts (Hardenbol et al. 1998; Hallam 2001) is generally poor (see Fu¨rsich et al. 2005b for discussion). However, the same pattern of relative sealevel change, facies development and succession of geodynamic events is recorded from the Upper Triassic –Jurassic of northern Iran (Alborz Mountains). In this area, the plate tectonic processes manifested in the Eo-, Main and Mid-Cimmerian events took place (Fu¨rsich et al. 2009a, b). The common tectono-sedimentary history of the two palaeogeographic domains implies that the Iran Plate behaved as a single structural unit during Late Triassic –Jurassic times. We thank D. Bosence (London) and A. Saidi (Tehran) for constructive reviews. The present study is part of a joint research programme of the Institute of Geology and Palaeontology of Wu¨rzburg University and Tehran University (Centre of Intelligence). We greatly appreciate the logistic support in the field offered to us by the Geological Survey of Iran (GSI). Most of the fieldwork was financially supported by the National Geographic Society (grant No. 5888–97). We also thank the Tabas Coal Company for its hospitality.

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Appendix

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Co-ordinates of mentioned localities

Name

Latitude N

Longitude E

Remarks

Badamu village Baghamshah valley Bidou village Desu village Esfak village (west of) Garedu mine Hojedk village Honu section

308190 2500 338400 2200 308480 0500 308470 3000 348020 3000 348130 5500 308450 4100 348290 3000

568470 5000 578100 5400 568590 4500 568410 4500 578060 1200 578070 2800 568590 3100 578150 4800

Kalshaneh

348080 5400

568410 1600

Kamar-e-Mehdi area

338010 1000

568260 3600

Korond Kuh-e-Abdullahi, NW of Deh Salm Kuh-e-Bagh-e-Vang

338550 0200 318330

578090 2000 588480

Type area of the Badamu Formation Type section of the Baghamshah Formation Type section and type area of the Bidou Formation Type area of (Cambrian) the Desu Formation Reference section for the marine Hojedk Formation Type section of the Garedu Red Bed Formation Type area of the Hojedk Formation Unconformable contact of the Esfandiar and Garedu Red Bed formations Reference section for the non-marine Hojedk Formation Type section and type area of the Kamar-e-Mehdi Formation Type section of the (lower) Korond Formation (Submarine) Jurassic volcanics on the Lut Block

338560 5400

568470 0500

Kuh-e-Birg

338170 1500

588190 2900

Kuh-e-Echellon (Echellon area) Kuh-e-Esfandiar Kuh-e-Neygu near Khoda-Afarid Kuh-e-Qoleh Nar

338490 0400

568360 1900

338100 5100 338440 4700

578270 3500 578120 2800

338020 0700

568260 5000

Kuh-e-Rahdar Kuh-e-Shisui

338380 4100 338370 1100

568200 4400 588000 0800

Lakar Kuh area

328110

758030

Mazinu area

338140 3900

568140 2400

Parvadeh mine Qal’eh Dokhtar

338000 0300 338550 4000

568460 5800 578170 4700

Qassem-Abbad Sikhor area Titu village

338540 3900 338230 3000 308540 4000

578150 2800 578220 4000 568370 1200

Kuh-e-Band-e-Buhabad

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Index Note: Figures are indicated in italic font, tables in bold. Ab-e-Haji Formation 133, 327 Absheron Ridge 241, 242, 243, 245–247, 258 accretion 76 accretionary prism 176, 244, 258 accretionary wedge central Iran 261, 263, 272, 273, 281, 316 northern Iran 32, 50, 51, 52 unroofing 280 acritarchs 40, 168, 169 age see radiometric age age spectrum diagram 73, 74 age, opening of South Caspian Basin 243 age, stratigraphic, Alborz Mountains 85, 91, 95, 102, 109– 110, 113 Aghdarband succession, comparison with Nakhlak 288, 316– 319 Aghounj Formation 184 –185, 186 Akchagyl Formation 250 Alam Formation 289, 291 palaeomagnetic interpretation 13, 15–17 petrography 269, 270, 271 stratigraphy 293–308, 310, 312, 313– 314, 318 structure 266, 268 Alasht Formation 186, 327 depositional environment 143– 144, 147 stratigraphy 146–149 Alborz belt 31–35, 59 provenance 49– 51 stratigraphy 35–36 structures 37–48 tectonic setting 37–38 Alborz Mountains 79–122 age 85, 91, 95, 102, 109– 110, 113 correlation 82, 120 depositional environment Pennsylvanian– Permian 87, 91 Permian 96, 102, 110, 113 Triassic 117 geodynamic interpretation 117, 118–120 previous research 80– 81 sedimentation 4 stratigraphy 82–117 Alborz thrust stack 3, 4 Alborz, palaeomagnetism 19, 24 algae 166, 169, 170, 172, 173 Permian 103, 104 alluvial fan 201, 266, 280 Kashafrud Basin 208 alluvial plain deposits 181 Alpine–Himalayan Collision Zone 244 ammonite studies 326, 329, 335 ammonites Jurassic 195, 197, 198, 199, 208, 214 Upper Triassic to Middle Jurassic 133, 138, 147, 149, 150, 154 ammonoid preservation 297 ammonoids, Triassic 293, 296, 298, 299, 306, 308, 311, 315

amphibole 69 mineral chemistry 276, 277 amphibolite facies 37, 59, 61, 62, 63, 74 Anarak Metamorphic Complex 5, 263, 265, 267, 272–279 geochemistry 276, 277, 280 petrography 275–279 Anarak Seamount 273 apatite 69 apparent polar wander path 19, 22, 26 Apsheron Sill 220, 222, 225, 230, 236 Arabian Plate 190 Ar– Ar radiometric age 45, 58, 74– 75, 265, 273, 280, 282 archaeocyathids 272 Arefi Formation 178– 181, 185, 186 Aruh laterite 9 Ashin Formation in palaeomagnetic interpretation 13, 15–17 petrography 271, 281 stratigraphy 308 –310, 311, 313 structure 266, 268, 269 Azerbaijan, South Caspian Basin 241– 259 depositional environment 248 –250 geodynamics 242–243 stratigraphy 244 –248 subsidence modelling 251– 258 tectonic evolution 222 –224, 243–244 cross-section 223 back-arc basin 52, 186, 221, 224, 251 back-arc extension 237 back-arc rifting, Neotethys 3, 156, 216, 330 backstripping 252, 253, 254, 257 bacteria 166, 172, 173 Badamu Formation 328, 329, 338 Baghamshah Formation 328, 331, 338 Baghamshah Subgroup 330–334, 338 bajada 186 Ba¯qoroq Formation petrography 310–313 provenance 270, 271, 280 siliclastic sediments 291, 315, 318 stratigraphy 266, 268, 269, 308, 312 Base Maykop event 245 basement, Greater Caucasus 245 basement, South Caspian Basin 229 basin-plain facies 212, 214 bastite 279 bathymetry, restored 254–255, 257 bauxite 140, 143, 145 Bazehowz Formation 181– 184, 185, 186 belemnites Early Jurassic 147, 149 Middle Jurassic 150, 152, 193, 195, 197 Benioff Zone 241 Bidou Group 326, 330, 333, 339 Binalud Basin 186

346 Binalud Mountains 175 –187 biochronology, Alborz Mountains 89 Permian 96, 98, 101, 113 Nesen Formation 102–107 Pennsylvanian 83– 87 Triassic 117 bioclastic limestone 302, 304 biolithites 115 bioturbated beds 165, 209, 213, 214 bivalves 162, 214, 335 Jurassic, Alborz 147, 149, 150, 154 Mid Jurassic Dalichai Formation 197 Dansirit Formation 193, 195 Permian 110 Triassic 117 Alam Formation 299, 303, 308, 310 Alborz 138, 140, 143, 144 black shale 212 Shemshak Group 138, 140, 141, 169, 170, 172 blueschist, Anarak Metamorphic Complex 272, 273, 274, 275 –277 mineral chemistry 276, 277, 280 blueschist, Shanderman Complex 63, 69, 70 Boghrov Dagh thrust sheet 50 boulder beds 180, 186, 208 brachiopods Jurassic 149, 195, 197 Pennsylvanian 85, 91 Permian 95, 98–99, 102–104, 107 Triassic 138, 306 brachiopods, extinction 116 braided fan, Kashafrud Basin 208, 209 braided river environment 182, 184 break-up unconformity 201, 202, 216 breccia 180 burial curves 251 burial history modelling 225 –228 burial temperature 4, 172– 173 burrows 100 Calcaires de Dorud 91 calc-alkaline magma 245 carbonate mounds 301, 304, 305, 314, 317 carbonate platform complex 130 Jurassic 205 Alborz 192, 195, 198 Tabas Block 334, 336, 338, 339 carbonates, Greater Caucasus 245 Central Caspian Basin 229 Central-East Iranian Microcontinent 323, 324, 325, 338 cerussite 268 Cha¯h Gorbeh unit 275–277, 278, 279 chamosite 112 channel fill 193 channel sandstone 148, 182, 184 chert 104, 121, 180, 184 Chitinozoa 39, 40 chlorite 69, 110, 275, 278 Cimmeria/continent/blocks 2, 3, 7, 31, 118 Cimmerian event 339

INDEX Cimmerian evolution, Nakhlak–Anarak area 261–284 metamorphism 265, 272–279 petrography 269 –271, 275–279 provenance 270, 271– 272 stratigraphy 266– 272 structure 263–268, 273, 282 Cimmerian orogeny 6, 31, 268 Early 130, 156 Cimmerian terranes 129 cleavage 37, 37, 223, 267, 270, 275 climate 4, 119 clinopyroxene 62, 71, 279, 281 clinozoisite 62 closure 57 coal seam 148 coal swamp 156 coal Jurassic northern Iran 193, 195 northeast Iran 182, 183, 207, 215 Tabas Block 327, 328, 329, 330 Upper Triassic to Jurassic 130, 145, 147, 154 coal, organic content of 166 coastal bar 193 coastal marsh 193 coastal plain 197 collision 262 Late Triassic 27, 51, 52, 57, 192, 205–206 Mid Jurassic 199 collision, Arabian-Eurasian plates 1, 3, 37, 48, 244 collision, Greater Caucasus 223, 224 Coloured Me´lange 263 compression 233, 236 Cretaceous–Miocene 248 compression tectonics 202, 245 condensed bed 197 conglomerate 223 Alam Formation 271 base Cretaceous 290, 291 conglomerate, analyses of 313 conglomerate, Jurassic 193, 194, 195 Binalud Mountains 179, 180, 181, 182, 183, 184 Kashafrud Formation 206, 208, 209, 210 Tabas Block 330, 332, 334, 336 conglomerate, Triassic, Nakhlak Group 296, 306, 308, 312, 315 conglomerate, Upper Triassic–Jurassic, Shemshak Group 137, 153, 154, 155 conodont analyses 310 conodonts Alam Formation 293, 299, 301, 302, 308 Alborz 109, 114, 117 Sina Formation 317, 318 continental facies 83 convolute bedding 208, 211 coral 147, 153, 193, 195 correlation, Binalud Mountains 185 correlation, Shemshak Group 191 crinoids 149 cross-bedded sandstone 194 cross-bedding, trough 183 cross-section, Alborz Range 172 crust density 231, 235, 237 crust, subducting 222

INDEX crustacea 150 crustal deformation 233, 234, 235, 236 crustal structure 241–242 crustal thickness 255, 256 cryptospore 40 Dalichai Formation 196, 198, 198, 208 Dansirit Formation depositional environment 152– 154, 191, 192–195, 196, 198, 200 lithology 152 sequence stratigraphy 198 –199, 201 dating see radiometric age debris flow 180, 208, 335 deformation, Shanderman eclogites 60– 62 delta facies 165, 210 delta front 193, 198 delta plain 192 deltaic sandstone 148 density, crust and mantle 231, 235, 237 depositional environment Alasht 141–142, 147 Alborz Mountains Pennsylvanian– Permian 86– 87, 91 Permian 96, 102, 109, 113 Triassic 117 Dansirit Formation 152– 154, 191, 192–195, 196, 198, 200 Kashafrud Formation 216–217 Nakhlak Group 313–314 South Caspian Basin 220, 248 –250 Tabas Block 327 see also Shemshak Group, depositional environment depositional history, Shanderman eclogites 59–61 depth maps 225 depth to basement map 246 Derekhtoot Member 180, 185– 186 desiccation cracks 333 desiccation crust 117 detachment zone 233, 234 dinoflagellate cysts 133, 138, 168, 169 docking, Iranian block 32 dolomite 245, 294, 337 dolostone 117 Dorud Group 82, 87– 98, 109, 118, 120 provenance 98 Dozdehband Formation 82, 84, 118 drift history of Iran 17–26 Early Cimmerian orogeny 189, 191, 207 Early Cimmerian unconformity 191 earthquake zone 37 Apsheron Sill 222 earthquake, Absheron Ridge 242 earthquake, Toram 59 Echellon Limestone Member 335 echinoid 335 eclogite 44, 45, 50 see also Shanderman eclogites effective elastic thickness (Te) 231, 252

Ekrasar Formation 134 –139 lithological logs 135, 136 Elikah Formation 82, 105, 111, 114, 115–117 carbonates 140 isotope curves 114, 116 Emarat Formation 87, 91– 96, 121 emersion surface 119–121 Eo-Cimmerian deformation 265, 266, 282 Eo-Cimmerian event 156, 156, 272, 330, 338 Eo-Cimmerian orogeny 3, 9, 27, 31–48, 57, 58, 59 Eo-Cimmerian suture 186– 187 epidote 67 erosion surface 118 Esfandiar Limestone Formation 332, 334 Esfandiar Subgroup 334, 337, 339 Euler poles 25, 26 evaporite 138, 337, 250 extension 229, 232, 234, 237 extension mechanism 242– 243 extensional tectonics 201 facies associations 208– 214 Greater Caucasus 245, 247, 248– 254 facies development, Tabas Block 327– 338 fan delta 209 faunal comparison, Aghdarband and Nakhlak 318 ferricrete 147 ferruginous crust 329 Fillzamin Formation 150–152, 191 fauna 195 lithology 192, 194, 196 flame structures 209, 211 flood plain deposits 182 fluvial cycles 333 fluvial environment 314, 315 fluvio-deltaic sequence, South Caspian 248 foraminifera 96, 100– 102, 107, 117 see also fusulinids forearc basin 130, 156, 230, 271– 272, 279 forearc basin, Nakhlak succession 316 foreland bulge 48 foreshore deposit 193 fungal filaments 166 fusulina-bearing limestone 59 fusulinid, cold water 265 fusulinids 92–93, 100 –102, 103, 109, 120, 121 Pennsylvanian 83– 87, 91, 91 gabbro, geochemistry 63 Galanderud Member 137, 138, 140 galena 268 Garedu Red Bed Formation 332, 336–337 Garedu Subgroup 336– 338, 339 garnet 65, 66, 67 gas pipeline section, NE Iran 213, 214 Gasht Complex 37, 51, 59 gastropod 138, 147, 150, 193, 197 geochemistry 276, 281 of organic matter 167 Shanderman eclogites 63–69 geochronology, Shanderman eclogites 73– 74 geothermal gradient and deformation 228

347

348 Ghalimoran Formation 41, 49 Ghosnavi Formation 84, 87, 93, 95– 96, 111, 121 glaciation, Gondwanan 117 Gondwana 3, 7, 19, 26, 31 Gondwana microcontinents 324, 338 Gondwana microplates 129, 261, 262, 263, 265, 283 Gondwana reconstruction 57, 75 Gondwanan terrane 118 Gorgan Schists 37–41, 49, 50 graben 48 granitoids 37 Greater Caucasus orogenic belt 221, 222– 224, 236 greenschist 272, 273, 275 –277 greenschist facies 59, 63, 74 Gre`s de Dorud, Formation des 87 gypsum 37, 333, 336, 337, 338 hardground, iron– rich 314 heavy minerals 83, 89, 98 Hemigordiopsidae 114 highstand system tract 198, 250 Hojedk Formation age 133 coal 328 facies 199, 202, 329, 338 hummocky cross-stratification 148, 192, 193, 195, 212 hydrocarbon reservoir 5, 161 hydrocarbon source rock 4, 214, 221, 247 Shemshak Group 161, 162, 170, 173 hydrogen index values 166, 168, 170, 171, 172–173 ilmenite 69 intermontane basin 186, 187 inversion 224, 237, 266 Iran microplate 118, 120 Iran Plate 3, 5, 59, 76, 205, 215, 323 active margin 266 early Mesozoic stratigraphy 129 palaeogeographic reconstruction 156, 325, 338 subsidence rate 201 Triassic collision 192 unconformity 191 iron crust 197 ironstone 98 isopach map, South Caspian 245, 246, 247, 249, 250 isostatic adjustment 228– 229, 235 isotope age see radiometric age isotope curves d18O and d13C 114– 117 Javaherdeh Formation 154, 186 Kalariz Formation 145– 146 Kamar-e-Mehdi Formation 332, 334, 335– 336 K–Ar radiometric age 38, 74–75, 265, 272, 282 Karaj Formation 37 karst 191 karstification 119 Kashaf Rud Formation 32 Kashafrud Basin 5, 186, 205

INDEX Kashafrud Formation 205 –208 age 200 depositional environment 216– 217 facies associations 208 –214 map 206 stratigraphy 208 lithological log 207, 210, 211, 214 Kashafrud rift-basin model 214–216 keratophyre dykes 338 kerogens 170, 172, 173 kinematic modelling 228 Kopet Dagh 57, 243– 245 Koppeh Dagh 1, 3, 5, 6, 175– 177, 189, 200, 205, 207 Korond Formation 332, 334, 335 Kura Basin, cross-section 222 Kurtian Member 180– 181, 186 lacustrine deposits 146, 154 South Caspian 248, 250 lacustrine environment 162, 165, 166, 168, 170, 172, 173 La¯k Marble 272, 273 Laleband Formation 140 –141 Lar Limestone 41 Late Cimmerian event 6, 336, 337, 339 laterite 9, 120 Permian 19, 24, 110, 112 Triassic 145 Laurasia 75, 75 lavas and palaeomagnetism 19, 26 lead mine 290 lignin 166 limestone Alborz Mountains 84, 95, 100, 114, 115, 117, 121 Kashafrud Formation 213, 214 Nakhlak Group 294, 296, 308 Shemshak Group 138, 139, 140, 141, 150 Tabas Block 328, 329– 338 Upper Cretaceous to Palaeogene 266 lithosphere extension 228, 230 lithostratigraphy, Mashad area 207 lithostratigraphy, Shemshak Group 133 –154 lowstand system tract 199, 250 Lut Block 324, 326, 336 Lutian (also Mid–Cimmerian) unconformity 4 –5 lydite 195 magmatic arc 265, 266 magnetic experiments 10– 11 magnetite 15 Magu Group 326, 330, 338 Magu Gypsum Formation 332, 337, 339 Maikop Suite 221 mantle lithosphere 228, 229 Masuleh-Shah Rud Unit 47, 50 Mega-Lhasa 31, 118 Messinian Salinity Crisis 222, 248 Messinian unconformity 250, 258 metacarbonates 272 metamorphic complex 37–38, 41 metamorphic event 3, 4

INDEX metamorphic rocks 265, 272–279 Shanderman eclogites 70– 73 Miankuhi Formation 207 microbialites 114 microfossils 40, 103 microplate collision 51 microprobe analyses 68 microstructures, garnet 62 Mid-Cimmerian tectonic event 130, 186 Alborz Mountains 189–199 east-central Iran 199– 200 geodynamic interpretation 192 Koppeh Dagh 200 Tabas Block 329, 331, 338, 339 Mid-Cimmerian unconformity 4– 5, 6 in Alborz Mountains 192, 195, 197 Lower 195– 197, 200 Upper 197– 198, 200 and compression deformation 236 Middle East Basins Evolution Programme (MEBE) 1, 80 mine, coal 327 mineral analyses 64– 69 mineralization, lead 290 Mobarak Formation 82 modelling, geodynamic, South Caspian Basin 221– 222 burial 225–228 constraints and parameters 229– 236 crustal attenuation 232– 235 lithosphere deformation 231 –232 oceanic crust 235 subduction of crust 235– 236 Moho, depth to 222 molasse 181 Cimmerian 156, 186, 192, 207 Early Jurassic 201–202 Eo-Cimmerian 50 Nakhlak Group, central Iran 287–319 age 297, 299, 301, 308 comparison with Aghdarband/Koppeh Dag 316–318 depositional environment 313– 314 detrital composition 311, 313 geodynamic interpretation 315– 316 palaeobathymetry 309 stratigraphy 288–289, 291– 311 lithological logs 292, 294, 295, 300, 301, 305, 307 Nakhlak-Anarak basin evolution 5– 6 map 267 nappe stack, 76 nappes 45– 48, 51 Nar Limestone Member 332, 336 nautiloid 138 Nayband Formation 133 Nazarkardeh Formation 318 necking depth 233 Neo-Cimmerian event 8, 266, 272 Neotethys 2, 3, 37 opening 26, 31, 34, 119, 156, 261, 266 subduction 58, 119, 338, 339 Nesen Formation 100, 102, 104–109, 113, 114, 119– 120 isotope curves 116, 121

oil and gas seepage 161 olistolith 223– 224, 237, 245 olistostrome 335 olivine relicts 279 oncolite 94 ophiolite 32, 51, 74 central Iran 266, 267, 272, 273 ophiolitic ring 263, 282 ore deposits 268 organic matter, Shemshak Group 161– 173 analysis 165–166, 170 classification 168 localities studied 162 palynofacies 166, 168 –170 Rock-Eval data 165– 166 shaly units 162– 165 ostracodes 109 pahoehoe 139 palaeobiogeography 319 palaeoenvironment see depositional environment palaeogeographic reconstruction 25, 325 Cimmeria 131, 156 Iran Plate 216 Pennsylvanian, Permian, Triassic 81 Permian, Middle 102 Permian, Upper 106 Permian-Triassic boundary 114 palaeokarst 139, 338 palaeolatitude 7, 19, 24–25, 26 palaeomagnetic curve 23 palaeomagnetic data 3, 7– 17, 263, 283, 284 Cimmeria 7– 17, 18 method 7– 8 palaeomagnetism in Iranian blocks drift history of Iran 17–26 mean directions 13 Palaeozoic 8– 9, 19 rotation 27 reference poles 20–21 Triassic 9 –11, 13, 14, 15– 17, 19–26 palaeomagnetism Sorkh Shale, 9– 11, 13, 14, 24–25 Palaeotethys 2, 3, 7, 26, 221 northern Iran 131, 161, 192, 203 subduction 119 Palaeotethys suture central Iran 261, 263, 265 northern Iran 32, 50– 52 north-eastern Iran 175, 176 Shanderman eclogite 57, 58 Palaeotethys, closure 31, 34, 51 palynofacies 166, 168–170, 172 palynology 37, 85, 107 palynomorphs 108, 113, 133, 169 Pan-African orogeny 31 Pangaea 25, 26, 27 reconstruction 75 paragenesis (eclogite) 63 Paratethys 247 Parvadeh Formation 197, 199–200 lithostratigraphy 328, 330– 331, 338 Parvar Member 139, 143– 145

349

350 passive margin 118– 119, 245, 248 Scythian Plate 222 patch reef 195 pedogenic 193 Permian-Triassic boundary 111, 113–115 petrography Alborz sandstone 83, 88– 89, 91, 96, 111, 112 Anarak Metamorphic Complex 275–279 Nakhlak Group 311–313 Nakhlak sandstones 269– 271 Shanderman eclogites 62, 63 Shemshak sandstones 48– 49, 53 petroleum see hydrocarbon plagioclase 69, 275, 278 plant fossil 133, 137 Jurassic, Lower 181– 182 Jurassic, Lower to Middle 149, 154 Jurassic, Middle 193, 209 Late Triassic 141, 146, 147 plate movement, speed of 3, 26–27 platform limestone 208 polar wander 19, 22, 26 pressure-temperature conditions 277, 280, 281 pressure-temperature estimate 70, 71, 72, 74 Productive Series 248, 249, 258 reservoir unit 221 provenance Shemshak sandstones 48– 49, 50 Nakhlak Group 311–313, 314, 315 Qal’eh Dokhtar Sandstone Formation 331, 334– 335 Qeshlaq Formation 88, 92, 110–113 Qezelqaleh Formation 83–87, 120, 121 radiometric age Ar– Ar 45, 51, 74, 265, 273, 280, 282 K –Ar 38, 74– 75, 265, 272, 282 Rb –Sr 37, 59, 265, 272, 282 U –Pb 34, 262, 272, 282, 283 rare earth elements, eclogite 63, 64 rate of sedimentation 151, 228, 258 South Caspian Basin 221–222 Ravar Formation 337 Rb–Sr radiometric age 37, 59, 265, 272, 282 red beds 336, 338 resinite 169, 171 rift basin model 214, 215–216 rift phase 231, 233, 234 Rock-Eval data 165–166, 167, 170 –173 rootlet 145, 146, 147, 193, 197, 329 rotation 334, 338, 339 central Iran 27, 287, 316, 318 Iran blocks 6, 263, 283 Ruteh Limestone 81, 98–102, 104, 111, 118, 120 rutile 69, 71 sabkha 337 Sanandaj Sirjan Zone 265, 266 scolecodonts 38, 40 sea-floor spreading 192, 202 sea-level change 118

INDEX sediment decompaction 254, 257 sedimentary fill, South Caspian Basin 221, 229 sedimentological analyses, South Caspian Basin 220 Sefid Kud Limestone 316 seismic data 225 seismic studies, South Caspian 241, 243– 244, 245 –247, 248 sequence stratigraphy, Dansirit Formation 198–199 serpentinized peridotites 277, 279 serpulids 147, 149, 193, 195, 197 Shah Zeid Formation 87, 93, 96– 98, 121 shale, Shemshak Group 162– 165 Shahmirzad Formation 141– 145 Shanderman Complex 3, 4, 37, 41– 45, 50 Shanderman eclogites (Complex) 57– 76 deformation 61–62 depositional history 58–59 geochemistry 58 geochronology 73– 74 metamorphism 64–74 petrography 61 tectonic evolution 58–59 shear 228, 231 shelf carbonates 214 shelf facies 212 shelf-slope facies 245 Shemshak Basin 201 Shemshak Group 4, 129 –158, 207, 262, 272 age 138, 143, 144, 146, 147, 150, 154 biostratigraphy 133 chronostratigraphy 134 depositional environment 154– 156 and see also separate entry below geodynamics 156 lithological logs 163, 164 location of sections 156–158 nomenclature 132, 153 organic matter in 161– 173 palaeoenvironment 154– 156 siliclastic sediments 161, 162 stratigraphy 133– 154 tectonic setting 34, 37, 50, 51 Shemshak Group, Binalud Mountains 175–187 Aghounj Formation 184– 185, 186 Arefi Formation 178–181, 185, 186 Bazehowz Formation 181–184, 185, 186 correlation 185 depositional environment 177, 180–182, 184–185 lithological logs 177, 178, 182, 184 map 176 Shemshak Group, depositional environment Jurassic 177 Aghounj Formation 184 Alasht Formation 142 –143, 147 Arefi Formation 180– 181 Bazehowz Formation 182 Fillzamin Formation 150, 152 Javaherdeh Formation 154 Shirindasht Formation 149, 150 Pennsylvanian–Permian 86–87, 91 Permian 96, 98, 102, 109– 110, 113 Triassic 116– 117 Ekrasar Formation 138, 140 Kalariz Formation 146

INDEX Laleband Formation 141 Shahmirzad Formation 142 –143, 145 Shemshak Group, Tabas Block 324, 327–330, 338, 339 Shemshak sandstone, petrography 48– 49, 53 Shirgesht Formation, palaeomegnetism 8–9, 12 Shirindasht Formation 149–150 Shotori Swell 327, 334, 339 Sikhor Formation 332, 333 siliclastic rocks Alborz Mountains 83, 84, 89, 96 Kashafrud Formation 207–214 Nakhlak Group 288–297, 299– 301, 306, 308, 311 Shemshak Group northeastern Iran 161, 162, 185 –186, 199, 207–214 northern Iran 134, 140, 142, 145, 153, 154 Tabas Block 323, 327, 329– 338 Sina Formation 207, 316, 317 slab break-off 187, 265 slope facies 212 slump 212, 223, 334 Sorkh Shale, palaeomagnetism 9 –11, 13, 14, 24– 25 South Boghrov Dagh thrust sheet 47–48, 50 South Caspian Basin 58, 156 depositional environment 220 evolution 1 –3, 221– 228, 229 extension 186, 202, 215, 216, 217 post-rift subsidence 339 research summary 3 –6 stratigraphic sequence 223, 224 subsidence mechanisms 228– 236 subsidence modelling 5, 251– 258 uplift 219– 238 South Caspian Basin, sedimentation and subsidence 241– 260 see also under Azerbaijan South Caspian to Central Iran Working Group 1 spinel 279, 280 sponges 334 staurolite 59, 61 storm influence 192 stratigraphy Jurassic see under Tabas Block Jurassic, Lower to Middle see under Shemshak Group Jurassic, Middle see under Kashafrud Formation Late Triassic to Mid Jurassic see under Shemshak Group Mesozoic to Tertiary see under Azerbaijan Pennsylvanian, Permian, Early Triassic see under Alborz Mountains Triassic see under Nakhlak Group strike-slip fault 37, 59, 270 stromatolites 115, 118 structural map Alborz belt 33, 36, 43 central Iran 264, 267, 288 Nakhlak 290 Iran 8, 32, 81, 190, 262, 282, 324 South Caspian Basin 58 subduction 2, 3, 32, 33, 225 Tethys Plate 219, 224, 237 subduction depth 74– 75 subduction zone, Alpine–Himalayan 244 subsidence mechanisms, South Caspian Basin 228–238

subsidence modelling, South Caspian 5, 251 –258 subsidence rate, Kashafrud Basin 215, 216 suture, Palaeotethys 32, 205, 216 syn-depositional faulting 200 syn-rift subsidence 252, 255, 256 syn-sedimentary faulting 214 Tabas Block 6, 199, 200, 323–340 age 330, 331, 333, 334, 335 depositional environment 327 geodynamics 338 locality details 339–340 map 326 research history 324, 326–327 stratigraphy 327 –338 Talesh Mountains 50 nappes 45– 48 stratigraphy 34 tectonic balanced cross-section 243 tectonic evolution, Greater Caucasus 222– 224 tectonic evolution, Shanderman eclogites 58–59 tectonic setting, Alborz belt 34– 48 temperature estimates 72 tempestites 113, 212 Tertiary deformation 46 thermal decay 9, 10–11, 15 thermal maturity 165 –166, 168, 170– 173 thermal recovery 228 thermal subsidence 227–228, 254, 256, 257, 258 thrust belt 230 thrust sheet 47, 50 thrust stack 51 thrusting and crustal thickening 228 titanite 69, 275, 277, 278, 277 toe-of-slope facies 212 Top Cretaceous event 245 Toyeh Formation 87–91, 121 trace fossils 165, 193, 195, 214, 329, 311, 334 trace fossils, Shemshak Group Jurassic 147, 149, 154 Triassic 138, 141–144, 146 Turan Plate 3, 177, 205, 318 collision 192, 200, 201 reconstruction 190 turbidites 212– 214, 315, 333 Ashin Formation 266 ultramafic cumulate 44– 45, 50, 61, 63 ultramafic rocks 272, 281 unconformity 120, 176, 223, 224, 316 Mesozoic 37, 39, 42 base Upper Cretaceous 269, 270, 290, 291 Early Cimmerian 137, 141 Eo-Cimmerian 34, 39, 41, 44, 47 Jurassic 207 Late Triassic 262 Triassic–Jurassic 155 Mid-Cimmerian 130, 152, 154, 156 Miocene 248 Permian–Triassic 179 unconformity and tectonism 189– 192, 338 unconformity, break-up 201, 202

351

352 unconformity, drowning 314 U–Pb radiometric age 34, 262, 272, 282, 283 uplift mechanisms, South Caspian Basin 219 –238 Variscan deformation 75 central Iran 3 –4, 261, 265, 272, 280, 281, 283 Variscan units 282 volcanic arc 37, 63– 64 volcanic arenites 269, 312, 313 volcanic rocks 245, 333, 337 Kashafrud Basin 206 Permian–Triassic 101, 102, 117 –119 Shemshak Group 138, 139, 140, 143, 145, 154 Talesh 34–35, 59, 61 volcaniclastic rocks 3, 266, 271

INDEX Nakhlak Group 288– 289, 294, 299 palaeomagnetism in 15 West Gondwana 19, 24, 25, 26 palaeomagnetic reference poles 20–21 white mica 67, 69, 275, 277, 278 whole-rock analyses 64 wood 194 working groups 1 Yazd Block 267, 324, 326 Zagros suture 31